Quantitative structural analyses and numerical modelling of ...

“It strikes me that all our knowledge about the structure of our Earth is very muchlike what an old hen would know of the hundred-acre field in a corner of which she isscratching.”Charles Darwin

AcknowledgementsThis habilitation thesis is the result of several years of research, and many people have, inone way or another, influenced this work. First and foremost, I would like to express mydeepest thanks to Prof. Karel Schulmann at whose department I have been studyingand working since 1995, and whose support and encouragement in all phases of myacademic career have been invaluable to me. I have always profited greatly from thestimulating discussions with Prof. Schulmann, and his ideas, advice, and critical readinghave boosted the present work enormously.Secondly, sincere thanks go to Drs. P. Štípská, J. Konopásek, S. Ulrich, P. Jeřábek,J. Ježek, F. Hrouda and V. Janoušek, with whom I had, or have, the great pleasureto collaborate on the subjects presented in this thesis. I also thank my former Ph.D.students and collegues P. Hasalová, M. Racek, P. Závada, J. Franěk, L. Baratoux forperforming joint research with me on their interesting research topics.I feel very much enriched by all my visits in Strasbourg. I would like to thank J.-B.Edel, G. Manatschall, J. Lehmann, E. Skrzypek, F. Chopin, and all the other people atthe CGS/EOST University of Strasbourg, for the countless and motivative discussions.Important part of this thesis was written while I was a distinguished postdoctoral fellowat this institute.It was furthermore a great pleasure to organize an international conference Granulites &granulites 2009 on the field of research partly covered by this thesis, which was pursuedtogether with Profs. K. Schulmann, R. White, M. Brown, and P. O’Brien. Thanks goto all of them for their great enthusiasm, help and advice.I wish to express my gratitude to the directors, to the staff, and to all the other membersat the Institute of Petrology and Structural Geology in Prague for providing excellentworking conditions and a fantastic working atmosphere as well as to Mrs. M. Wontrobováfor her kind and efficient administrative support.So here I am at the end. On my personal treasures. I would never have achieved thishabilitation without the support of my wife Markéta and her infinite understanding formy staying late at work, disappearing for fieldwork, traveling on a conferences, beingabroad. . . And I would never be where I am without my family, and without my friends. . .v

Contentsviii4.7 Schulmann, Konopásek, Janoušek, Lexa, Lardeaux, Edel, Štípská, andUlrich 2009b . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1594.8 Lexa, Schulmann, Janoušek, Štípská, Guy, and Racek 2011 . . . . . . . . 1814.9 Franěk, Schulmann, Lexa, Tomek, and Edel 2011a . . . . . . . . . . . . . 2054.10 Lexa, Štípská, Schulmann, Baratoux, and Kröner 2005 . . . . . . . . . . . 2314.11 Baratoux, Schulmann, Ulrich, and Lexa 2005b . . . . . . . . . . . . . . . . 2494.12 Závada, Schulmann, Konopásek, Ulrich, and Lexa 2007 . . . . . . . . . . . 2794.13 Schulmann, Martelat, Ulrich, Lexa, Štípská, and Becker 2008b . . . . . . 2954.14 Hasalová, Schulmann, Lexa, Štípská, Hrouda, Ulrich, Haloda, and Týcová2008a . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3154.15 Franěk, Schulmann, Lexa, Ulrich, Štípská, Haloda, and Týcová 2011b . . 341

List of Figures1.1 Strain evolution modelling . . . . . . . . . . . . . . . . . . . . . . . . . . . 61.2 Mechanical and rheological evolution of orthogneisses and migmatites . . 71.3 Parameter space for homogeneous transpression . . . . . . . . . . . . . . . 81.4 Schematic 3d structure of Gemer unit . . . . . . . . . . . . . . . . . . . . 101.5 Numerical model for continental indentation . . . . . . . . . . . . . . . . . 101.6 Oblique fold sections . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 111.7 Schematic 3d structure of Jeseníky mountains . . . . . . . . . . . . . . . . 121.8 Fold analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 121.9 AMS data of Třebíč massif . . . . . . . . . . . . . . . . . . . . . . . . . . 141.10 Numerical simulation of AMS fabric development . . . . . . . . . . . . . . 142.1 Schematic profile across Central European Variscides . . . . . . . . . . . . 202.2 Kinematic model of synchronous folding and erosion . . . . . . . . . . . . 212.3 N-S profile of eastern margin . . . . . . . . . . . . . . . . . . . . . . . . . 212.4 Numerical modelling of Rayleigh-Taylor instability within orogenic root . 222.5 Example of gravity inversion modelling . . . . . . . . . . . . . . . . . . . . 232.6 Model of thickening from above process . . . . . . . . . . . . . . . . . . . 242.7 Geodynamic model of crustal indentation . . . . . . . . . . . . . . . . . . 253.1 Crystal size distribution development . . . . . . . . . . . . . . . . . . . . . 313.2 Contrasting microstructures of amphibolites . . . . . . . . . . . . . . . . . 333.3 Melt pockets preferred orientation . . . . . . . . . . . . . . . . . . . . . . 343.4 Grain boundary statistics evolution . . . . . . . . . . . . . . . . . . . . . . 343.5 CSD in channel flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 353.6 Example of digitized microstructures . . . . . . . . . . . . . . . . . . . . . 363.7 Coupled analysis of textures and microstructures . . . . . . . . . . . . . . 37ix

Dedicated to my wife Markéta, daughter Eliška and sonJáchym.... . .xi

ForewordThis habilitation thesis consists of three main parts, which aim to underline the principaldirections and features of my scientific career and gives the overview of new scientificdirections, which I would like to follow as an associate professor of the Charles University.My research career is based on collaborative work of a larger research team, whichworked at the Institute of Petrology and Structural Geology from 1995 to 2011 and myresearch interest and focus was necessarily influenced by the overall aim of that group.This was particularly expressed in mathematical modelling of continental deformationin the micro-, meso - and macro-scale and the development of the necessary softwaretools for specific geologically oriented quantitative studies, which were published underthe Institute of Petrology and Structural Geology mostly in collaboration with K. Schulmann,J. Konopásek, P. Jeřábek, P. Štípská, S. Ulrich and others, who initiate manysubjects further discussed in this text. My skills in computer programs development,computer-based processing and statistical evaluation of structural data, implementationof database and geographic information systems and ability to model processes numericallyallowed me substantially contribute to several manuscripts of different authors thatare cited in this habilitation thesis. As these techniques were progressively more andmore used by IPSG research team, it became the driving force for three major researchdirections of my career presented here. These are following themes: 1) quantitativeanalysis of deformation structures and numerical modelling of deformation,2) mechanisms of lower crustal flow and its thermal and mechanical implicationsand 3) quantitative analysis of metamorphic microstructures, their visualizationand statistical processing. At this point I would like to stress that I highlyappreciated multidisciplinary approach of collective scientific collaboration, which ruledin the last ten years of my professional career in the Institute of Petrology and StructuralGeology and that helped me to identify the essentials of my personal approach, as wellas the professional qualities of my colleagues. It is my aim further transfer this spirit ofactive cooperation and stimulating discussions to students and cultivate this approachat Charles University in the future.1

Foreword 2First topic “Quantitative analysis of deformation structures and numerical modellingof deformation” includes mainly numerical modelling of progressive development of finitestrain and superposition of deformation events in collision zones. This part of myresearch was originally related to the simple kinematic modelling of deformation pathsin the Flinn space (Konopásek et al., 2001, 2003) and modelling of finite strain withintranspressional zones (Schulmann et al., 2003). Later, I focused on a “thin-sheet” numericalmodelling of deformation in collision zones characterized by complex geometriesand variable boundary conditions (Lexa et al., 2003). Part of my research activities wasdevoted to quantitative geometrical evaluation of so-called “extensional shear bands”,which are commonly used as tectonic transport indicators (Lexa et al., 2004) and possibleuse of fold shape analyses as a proxy of degree of mechanical anisotropy and bulkdeformation intensity (Baratoux et al., 2005a). Recently, my research activities areoriented towards understanding of deformation overprints in magmatic and solid staterocks, which are based on my early research (Konopásek et al., 2001) and whose resultsappear in some recent works of my co-workers (e.g. Kratinová et al., 2010, Lehmannet al., 2011, Schulmann et al., 2009b). This topic is crucial in understanding the symmetryand intensity of deformation in orogenic zones and it will be discussed in thechapter “Perspectives of numerical modelling of deformation in geology”. In all studies,where I am not listed as first author I have decisively contributed to the results by meansof creating software, own contribution in formulating the manuscript as well as field researchincluding the West Carpathians, Bohemian Massif, Central Asian Orogenic Beltand the Vosges Mountains.The second topic “Mechanism of lower crustal flow and its thermal and mechanicalimplications” is based on field study of two areas: the eastern margin of the BohemianMassif and central Western Carpathians. Research related to channel flow model alongthe eastern margin of the Bohemian Massif is the subject of numerous publicationswhich I co-authored with my colleagues and students, for example, Schulmann et al.(2005), Racek et al. (2006) and Franěk et al. (2006). Our research gained new stimulusby recently published description of kinematics of the channel flow model (Schulmannet al., 2008a), where the channel geometry, thermal structure and polyphase nature ofstructural evolution were firstly identified within Variscan orogenic system. New developmentsbased exclusively on my own research are further discussed in “Perspectivesfor the thermomechanical modelling of channel flow in the orogeny”, where I illustratethe plausibility of our model in terms of numerical modelling for different boundaryconditions. However, there is a second type of channel flow, which is linked to influx ofcrustal material into the orogenic root, which is subsequently extruded during exhumationstage. This model was firstly proposed on the example of the Western Carpathians,where some aspects of the lower crustal flow were numerically simulated (Jeřábek et al.,

Foreword 32008, 2007) and subsequently applied to describe evolution of orogenic root in BohemianMassif. Here we used available geochemical and geochronological data to argue that thefelsic lower crust is an allochthonous body emplaced underneath pre-existing mafic lowercrust during continental subduction (Schulmann et al., 2009b). Similar scenario was recentlyproposed for evolution Orlica-Śnieżnik dome in Western Sudetes (Chopin et al.,2011a). It appears that influx of crustal material to the orogenic root zone have a majorimpact on his future thermal and rheological evolution. Detailed study of geodynamicconsequences caused by subduction of felsic and radioactively productive material underthe orogenic system is described in Lexa et al. (2011) and this aspect will be furtherelaborated in the chapter “Perspectives for the thermomechanical modelling of channelflow in the orogeny”.The last topic “Quantitative analysis of metamorphic microstructures, its visualizationand statistical description” introduce a comprehensive set of quantitative methodsfor analysis of deformation microstructures and it is based on PolyLX software developedby myself. This chapter will deal with techniques of microstructure digitalization andsubsequent statistical processing, including evaluation of the grain size, grain shapes andstatistics of their mutual contacts. The PolyLX covers several methods of quantificationof preferred orientation of grains and boundaries as well as their statistical processing.This method allows us not only visualize microstructure characteristics and their statisticaltreatment, but also address some kinetic problems of metamorphic microstructures.This method was originally described in my Ph.D. thesis and firstly applied in the workof Lexa et al. (2005). Subsequently, the PolyLX software was successfully used in numerouspublications of our research group (Baratoux et al., 2005a,b, Hasalová et al., 2008a,Jeřábek et al., 2007, Kratinová et al., 2010, Lexa et al., 2005, Machek et al., 2007, Schulmannet al., 2008b, Závada et al., 2007, 2009). In addition, there is a number of laboratoriesin France, where I worked during my post-doctoral fellowship and I led specializedcourses in this subject, which are routinely using PolyLX software in complex microstructuralproblems (Barraud, 2006, Chopin et al., 2011b, Martelat et al., 2011, Oliot et al.,2011). It appears, that PolyLX becomes valuable tool for projects, where quantificationof the spatial relationships of minerals is needed. Recently I leed three specialized workshops(Granulites & Granulites 2009, Prague; 2010 Tromsø; Geomaterials 2011, Prague)focused on quantitative microstructure analyses. Recently, PolyLX has interfaces toexchange data with softwares like ELLE (http://www.materialsknowledge.org/elle) orMTEX (http://code.google.com/p/mtex) that allows simulating the evolution of deformationmicrostructures or quantitative textural analyses of natural samples. Futherideas as some of the new kinetic and petrogenetic outputs that could be depicted bythis method (Franěk et al., 2011b) are given in the chapter “Perspectives for numericalmodelling of deformation microstructures and their quantification”. As PolyLX is open

Foreword 4source and freely available on the internet (http://petrol.natur.cuni.cz/˜ondro) we hopethat it will become a widely used method in petrology and structural geology.Ondrej Lexa: lexa@natur.cuni.cz

Chapter 1Quantitative analysis ofdeformation structures and theirnumerical modellingFinite deformation modelling is an approach that is used in structural geology for decadesas a proxy to study distribution and nature of deformation within crustal collision zones,zones of transpression or transtension and lithospheric extension (Sanderson and Marchini,1984). The main question arises in situations, where rock masses do not behaveaccording to “simple rules” of homogeneous deformation, i.e. where considerablevariation in lithology and rheology introduce heterogeneities, which are likely to controltheir complex deformation behaviour accompanied with strain superposition (Burg,1999). Konopásek et al. (2001) described in western part of Erzgebirge steeply dippinglineations contained by steeply dipping eclogite foliations and horizontal lineations insurrounding structurally conformable orthogneisses. This conflict has led us to modelevolution of deformation pattern within “soft” incompetent orthogneiss during verticalflattening. In our model, these orthogneisses initially showed a similar orientation offinite strain axes as eclogites but after subsequent vertical shortening suffered viscousdeformations, which does not affected competent eclogites. Our numerical simulationshave shown conclusively that the superposition of deformation stages is a powerful mechanismthat allows us to understand some strain paradoxes as one described above. Theresult of our modelling (Fig. 1.1) shows that the contrast in deformation record of eclogitesand orthogneisses results from common deformation history, which is selectively andto different extent recorded in individual lithologies.Our results were challenged by German colleagues Krohe and Willner (2003) and wewere obliged to defend our hypothesis in the manuscript Konopásek et al. (2003) using5

1. Quantitative analysis of deformation structures and their numerical modelling 6Figure 1.1: Calculated evolution of the finite strain as a result of homogeneous verticalshortening of previous fabric ellipsoid marked by vertical X-axis. Black and grey linesshow the progressive strain being labelled by elongation of former X-axis. Results ofcalculations is shown for deformation fabrics marked by contrasting strain symmetriesand intensities as observed in the field. a) low strain domain with plane strain symmetryb) high strain domain with prolate shape.our numerical modelling as a backbone of the discussion. More recently, the problemof strain superimposition was elaborated again in the study of orthogneiss migmatitesdeformation in the French Vosges (Schulmann et al., 2009a). Here, the originally verticalstructures of orthogneisses embedded by metasedimentary migmatites are overprintedby vertical shortening due to larger scale crustal extension, which resulted in developmentof prolate strain fabrics in the so-called “internal margin” and oblate strainfabrics in the “external margin” of the crustal scale orthogneiss inclusion. Our detailedstructural and AMS study revealed that these changes correlate well with melt fractionand orientation of mechanical anisotropy relative to stress axes applied to orthogneissbody by flow of surrounding migmatites (Fig. 1.2). It turns out that in order to explainprolate finite strain ellipsoids within orogenic zones, one must account to the model of

1. Quantitative analysis of deformation structures and their numerical modelling 7Figure 1.2: Three stages of the mechanical and rheological evolution of orthogneissbodies and host metasedimentary migmatites: (a) Early stage of layer parallel shortening.(b) Later stage in the deformation when the margin of the orthogneiss body hasbeen transformed to an orthogneiss/granite multilayer. (c) In the final stage furtherhomogeneous deformation of host diatexites is accompanied by flattening and intensificationof the orthogneiss fabric.deformation overprints, otherwise we are exposed to the problem of strain compatibility(formation of open spaces within rock masses is not possible for obvious reasons). In thiswork, numerical models were set-up to reproduce the variation in the finite strain andorientation of the original mechanical anisotropy with respect to external stress field.Although the model successfully explained the structural zonality of rigid domains inthe migmatites, it was not robust enough to understand general dependencies of thefinite strain intensity and symmetry to the degree and symmetry of initial mechanicalanisotropy. In the following I will show some other perspectives of the problem.Relative motion of lithospheric plates on a spherical surface is such that the plateconvergence vectors are often not orthogonal to plate boundaries (Dewey, 1975, McKenzieand Parker, 1967). These plate boundaries experience combined transcurrent andconvergent displacements associated with development of deformation zones of differentsize. Within continental blocks, the deformation is not only restricted to active plate

1. Quantitative analysis of deformation structures and their numerical modelling 8Figure 1.3: Superposition of strain intensity (D), strain symmetry (K), and sampleelevation for various values of rigid floor depth and initial sample depth z 0 , plotted interms of angle of convergence (α) and time parameter (k t ). The values of D show astrong dependence on ratio between convergence velocity and width of deformed zone(R vd ) and time for α = 20 ◦ − 90 ◦ ; for α < 20 ◦ the dependence of D on R vd increases.K is strongly dependent on a but not much on R vd and time.boundaries but occurs within zones of weakness inside rigid continental domains andcan be approximately described as a deformation of a weak tabular zone bounded byrigid blocks with steep parallel walls. All the mentioned types of deformation zonescan be more or less described by a model called “transpression”, introduced first byHarland (1971), developed by Sanderson and Marchini (1984), and elaborated by manyothers. In the paper Schulmann et al. (2003), we summarize our experience with thisissue and we have modelled strain evolution in such zones including various deformationpartitioning schemes, namely 1) discrete partitioning (lateral component of motion isaccommodated by discrete faults), 2) ductile partitioning (lateral component of movementis accommodated by narrow zones of ductile flow) and 3) viscous partitioning(deformation zone is divided to domains where deformation is accumulated at differentstrain rate). In this model, we also included the concept of a rigid floor, which controls

1. Quantitative analysis of deformation structures and their numerical modelling 9elevation and exhumation of rocks. The results of our model (Fig. 1.3) shows systematicvariability in strain pattern, depending on the convergence angle and strain-rate withinthe transpressive zones. This fundamental behaviour could be used to inverse field datainto controlling parameters in both active and fossil deformation zone.One of the most important published contributions to this topic (Lexa et al., 2003)is based on our detailed structural analysis of polyphase deformation within the Gemerunit in the Western Carpathians. Here, the Paleozoic rocks deformed during LowerCretaceous convergence form steep cleavage fan between the more rigid blocks of easternand western part of the Vepor crystalline basement. Moreover, the so-called Trans-Gemeric shear zone is developed along eastern margin of the Vepor promontory andfurther east separate Gemer unit into two domains. By contrast, the eastern part ofthe Vepor basement suffered more frontal convergence with Gemer unit documented bydevelopment of Eastern Gemer thrust (Fig. 1.4). We tried to justify this complicatedtectonic model using viscous sheet formulation modelled by finite element method, wherewe simulate deformation in front of the moving indenter, which progressively deformedviscous domain in between the stationary blocks of eastern and western Vepor basementand which was free to thicken and outflow among these blocks.The result of numerical modelling (Fig. 1.5) was more than surprising, since both thesymmetry and distribution of viscous deformation of thin viscous sheet corresponds withsufficient precision to our observations. We consider this model as particularly valuablecontribution to the simulation of continental deformation. This example suggests thatthe concept developed since the eighties of last century (England and Houseman, 1986,Houseman and England, 1986) for deformation lithospheric scale deformations (deformationof Tibet in front of India indenter) could be successfully applied to significantlysmaller scale of geological units.Our simulations showed that our model for transpression was satisfying startingestimate for the deformation evolution in areas with complex kinematic framework. Webelieve that the deformation of continental crust is dominantly controlled by the presenceof lateral inhomogeneities of different shape and that the deformation is subsequentlypartitioned and channelized into narrow zones. Understanding polyphase deformationas consequence of principles of continuum mechanics is a challenge that will be testedin the near future theoretically as well as on field examples.Our field studies and observations rise the question of understanding the flow kinematicsin anisotropic rocks. The most common kinematic indicators used by field geologistsare “extensional shear bands”, which form basis of argumentation in numberof kinematic orogenic models (e.g. Behrmann, 1988, Neubauer et al., 1999, Platt and

1. Quantitative analysis of deformation structures and their numerical modelling 10Figure 1.4: Geometry of main tectonic units of Central Western Carpathiansand schematic evolution of major structures during progressive deformation used toparametrize subsequent numerical model.Figure 1.5: Finite strain pattern developed in Gemer unit after 5 Myr of shortening.(a) Distribution of strain intensity (D) and orientation of finite strain ellipsoid. (b)Distribution of probability of possible simple shear reactivation of evolved anisotropy.(c) Distribution of finite strain symmetry (K). (d) Distribution of finite topography (inmeters) in front of the indenter.

1. Quantitative analysis of deformation structures and their numerical modelling 11Figure 1.6: Stereographic projection showing influence of inter-limb angle and anglebetween fold axial plane and main anisotropy. The range of section planes showingshear-band geometry increases with inter-limb angle of folds affecting main anisotropy.Vissers, 1989). Some field observations, however, were cautionary: Are geologists alwayssure about mutual orientation of observed plane and kinematic frame? Are all the“extensional shear bands” really extensional structures? I tried to answer these questionsin a simple fold section model (Lexa et al., 2004) which showed that from all thepossible sections of 3D fold structure, only a small percentage allows us to observe theactual fold shape. In fact, in most of the sections, asymmetric fold structure appears as“extensional shear bands”. Our own observation from the central Western Carpathiansand from eastern margin of Bohemian Massif revealed that most of “extensional shearbands” represent only the oblique sections across compressive folds (Fig. 1.6).This important observation seriously modifies the today view on origin or core complexes,which can be only the oblique sections across the crustal scale folds. For instance,the discussion of origin of Montagne Noire gneiss dome as a core complex or steep crustalmegafold is debated for more than twenty years by French geologists (e.g. Aerden, 1998,Matte et al., 1998, Van den Driessche and Brun, 1992). We believe that the analysisof kinematics of such orogenic systems requires a thorough revision in the light of theseobservations. Similar conclusion could be made about the origin of Tauren Window as itis shown by contrasting interpretations (e.g. Ratschbacher et al., 1989, Rosenberg et al.,November, 2004).

1. Quantitative analysis of deformation structures and their numerical modelling 12Figure 1.7: 3d blockdiagram of general structure of northern termination of Silesiandomain. Figure shows progressive increase of late deformation and folding superimposedon previous flat fabric in conjunction with increase in metamorphic grade toward west.Figure 1.8: Selected fold assemblages from the garnet, staurolite and sillimanite zones,respectively. (a) Dip isogon patterns. (b) b1 (fold shape) vs. F (degree of fold flattening)plots constructed for the presented folds showing the relative importance of foldflattening and active fold amplification.

1. Quantitative analysis of deformation structures and their numerical modelling 13Some models of viscous deformation of thin plates show the characteristic deformationgradient, which is generally difficult to identify. Are we able to recognise sucha gradient within anisotropic rock masses? And what is the effect of lateral temperaturegradient on deformation of the crust? We have tried to answer those interestingquestions in the paper Baratoux et al. (2005b), where we observed structural variationsin shape and flattening of mesoscopic folds across 40 km long section (Fig. 1.7). Thisprofile was characterized by buckle and chevron folds in the east, flattened buckle foldsin the centre and flow folds in the west. Quantitative analysis of fold shape in termsof Fourier harmonic analysis method and the Ramsay’s method of dip isogons pointedto systematic decay in degree of mechanical anisotropy from east to west. This trendwas correlated with PolyLX analyses of microstructures of amphibolites, which revealexistence of gradient in shape anisotropy on grain scale. This detailed quantitative workconcludes that the character of folding is directly controlled by degree of anisotropy ofmicrostructure, so that the grain shape anisotropy governed mechanical properties offolded macroscopic layers. In another words, lateral variations in fold style correlatewith lateral variations in the mechanical anisotropy which reflect metamorphic zonalityi.e. thermal history of previous tectonometamorphic event.1.1 Perspectives of numerical modelling of deformation ingeologyUse of numerical modelling in structural geology introduced new possibilities especiallyin the area of deformation overprints. The range of deformation structures, which aretraditionally explained as the result of “magic” deformation partitioning into zones ofpure shear (prolate and oblate) and simple shear dominated deformation can not explainfabric variations encountered in orogens. This concerns the problem of strain compatibilityand second the problem of significance of deformation intensities. Therefore wedecided to elaborate methods to quantify and numerically simulate these deformationprocesses (e.g. Kratinová et al., 2010). For example, long-standing study of Třebíč syenitemassif led us to conclude is that the magma body suffered multiple deformationsduring magmatic stage, which led to the distinctive variance of the magmatic fabricsthat are not explainable in terms of standard approaches (Lexa et al., in prep.). OurAMS study shows that northern part of the body exhibit NW-SE trending lineation accompaniedwith oblate shape of AMS ellipsoid, while the southern part is characterizedby NE-SW trending lineation and prolate shape of AMS ellipsoid (Fig. 1.9).Our numerical modelling revealed that the variance in the symmetry and direction oflineations could be successfully explained by the superposition of horizontal flow (within

1. Quantitative analysis of deformation structures and their numerical modelling 14Figure 1.9: Map of P and T parameters of AMS fabrics with selected homogeneousAMS data showing two domains within Třebíč massif.Figure 1.10: Numerical simulation of progressive deformation of uniformly dispersedbiotite grains for horizontal shortening (green frame) superimposed by vertical shorteningand perpendicular elongation (orange frame). Strain paths in PT space show highvariability of AMS fabric shape (T) for similar values of degree of AMS (P).

1. Quantitative analysis of deformation structures and their numerical modelling 15crustal channel) on older vertical magmatic fabrics that are folded and therefore rotatedduring early stages of channel flow. This numerical model shown that routine utilisationof methods of AMS fabric interpretation could easily fail to correctly interpret plutonsemplacement histories. It appears that the superposition of two distinct plane straindeformations with various mutual geometrical relations could produce broad spectrumof finite deformation symmetries as well as finite deformation intensities which are commonlyobserved in real rocks (Fig. 1.10). Recently we also investigated intrusions ofmagmatic sheets related to exhumation of crustal scale gneiss dome Lehmann et al.(2011). The magmatic sheets are emplaced along boundary between core of gneiss domeand its mantle in contrasting tectonic regimes. Our modelling showing that systematicfabric transitions are related to superpositions of strain in magmatic state. It should bepointed out that without these experiences and deep understanding of strain accumulationin magmatic rocks, geologists can easily commit wrong geological models.

1. Quantitative analysis of deformation structures and their numerical modelling 161.2 Accompanying publicationsKonopásek, J., Schulmann, K., Lexa, O., 2001. Structural evolution of the centralpart of the Krušné Hory (Erzgebirge) Mountains in the Czech Republic; evidencefor changing stress regime during Variscan compression. Journal of Structural Geology23 (9), 1373–1392. 49Schulmann, K., Thompson, A. B., Lexa, O., Ježek, J., 2003. Strain distributionand fabric development modeled in active and ancient transpressive zones. Journalof Geophysical Research, B, Solid Earth and Planets 108(B1), 2023, doi:10.1029/2001JB000632. 69Lexa, O., Schulmann, K., Ježek, J., 2003. Cretaceous collision and indentationin the Western Carpathians: View based on structural analysis and numericalmodeling. Tectonics 22 (6), 1066, doi:10.1029/2002TC001472. 85Lexa, O., Cosgrove, J., Schulmann, K., 2004. Apparent shear-band geometry resultingfrom oblique fold sections. Journal of Structural Geology 26 (1), 155–161.101Baratoux, L., Lexa, O., Cosgrove, J. W., Schulmann, K., 2005a. The quantitativelink between fold geometry, mineral fabric and mechanical anisotropy; as exemplifiedby the deformation of amphibolites across a regional metamorphic gradient.Journal of Structural Geology 27 (4), 707–730. 1091.3 Related publicationsKonopásek, J., Schulmann, K., Lexa, O., 2003. Reply to comments by A. Krohe andA.P. Willner on ”Structural evolution of the central part of the Kru né Hory (Erzgebirge)Mountains in the Czech Republic—evidence for changing stress regime duringVariscan compression”. Journal of Structural Geology 25 (6), 1005–1007. Schulmann, K., Edel, J.-B., Hasalová, P., Cosgrove, J., Ježek, J., Lexa, O., 2009a.Influence of melt induced mechanical anisotropy on the magnetic fabrics and rheologyof deforming migmatites, Central Vosges, France. Journal of Structural Geology31 (10), 1223–1237.Lehmann, J., Schulmann, K., Edel, J.-B., Hrouda, F., Lexa, O., Štípská, P., Haloda,J., 2011. AMS and structural record of granitoid sills emplacement during growthof continental gneiss dome: implications for exhumation of deep crust. Journal ofStructural Geology, submitted.

1. Quantitative analysis of deformation structures and their numerical modelling 17 Kratinová, Z., Ježek, J., Schulmann, K., Hrouda, F., Shail, R. K., Lexa, O., 2010.Noncoaxial K-feldspar and AMS subfabrics in the Land’s End granite, Cornwall:Evidence of magmatic fabric decoupling during late deformation and matrix crystallization.Journal of Geophysical Research, B, Solid Earth and Planets 115,B09104, doi:10.1029/2009JB006714.

Chapter 2Mechanisms of lower crustal flowand its thermal and mechanicalimplicationsMechanisms of lower crustal flow are reflected in two different crustal levels linked bothto the processes of orogenic root building and exhumation of the orogenic lower crust.These processes have been evaluated on the basis of our own structural, petrologicaland geochronological studies along the eastern margin and from the central part of theBohemian Massif (Franěk et al., 2006, Racek et al., 2006, Schulmann et al., 2005). Theprinciple of this model is the large-scale folding of the bottom crust as a response tolateral forcing. Amplification of folds gradually passed to vertical extrusion, which wassubsequently transferred to horizontal flow just below the brittle-ductile transition. Thehorizontal flow is driven by two processes: 1) collapse of vertical extrusion related fabricsdue to excess of gravity potential energy of elevated orogenic lid and 2) horizontalsubsurface flow driven by indentation of Brunia basement generating hot and heterogeneousfold-nappe (length of 200 km). The initial folding and extrusion of lower crustis probably controlled by the ongoing subduction of Saxothuringian continental lithosphere(Schulmann et al., 2009b) while the horizontal channel flow is possibly controlledby change in lithospheric plates configuration (Fig. 2.1). The whole concept is clearlyexplained in the work of Schulmann et al. (2008a), where we report a full geochronological,structural and petrological inventory supporting the model of vertical extrusionand subsequent subsurface horizontal flow processes.The subsurface horizontal flow or horizontal spreading of exhumed orogenic lowercrust are generally interpreted as a result of ductile rebound of buoyant crustal root19

2. Mechanisms of lower crustal flow and its thermal and mechanical implications 20Figure 2.1: Model of laterally forced overturns related to Saxothuringian convergencein Bohemian Massif.(Koyi et al., 1999). In detail the process was described by Rey et al. (2001) or Vanderhaegheand Teyssier (2001) who discuss possible structural levels of horizontal flowresulting from rheological stratification and boundary conditions of their models. Thistype of orogenic collapse was recognized in two specific areas of Bohemian Massif: thesouth Bohemian granulites (Franěk et al., 2006, 2011a) and the Lugian domain of WesternSudetes (Dumicz, 1979, Štípská et al., 2004). In our new studies we have describedmetamorphic evolution related to burial of orogenic middle crust contemporaneous withextrusion of lower crustal omphacite bearing granulites (Skrzypek et al., 2011b). Subsequentsubhorizontal flow led to exhumation of the crust from 30 km depth to about 10km during which horizontal foliation associated with sequential growth of cordierite andandalusite originated. This particular structural evolution resembles ductile thinningmechanism described in subduction wedges (Ring et al 2009, Brandon and Fe,...XXX).In order to image this process I have modelled the vertical exchange of crustal materialas laterally forced gravity overturns followed by ductile thinning of crust bellow thebrittle-ductile transition (Fig. 2.2). The thermal field and P-T-t paths were simulatedfor both exhumation mechanisms in order to correlate petrological and geochronologicaldata with modelling results. We shown that the horizontal spreading is a localised processoperating in several kilometres wide layer bellow orogenic lid which is progressivelydestroyed during ductile thinning process (Skrzypek et al., 2011b, Štípská et al., 2011).However, different type of large-scale horizontal flow of hot lower crustal materialoperates at the eastern margin of Bohemian Massif (Fig. 2.3). Here, the hot orogeniclower crust is thrust over the continental basement which is documented by gravityinversion modelling (Guy et al., 2011). In this work we have shown that the crustalbasement unit is covered by thin layer of low density migmatitic material (Fig. 2.5). Inaddition, new petrological studies of Hasalová et al. (2008a,b) and Štípská et al. (2008)show that the flow of migmatites occurred above the basement at the depth of 25 kmcorresponding to 7 - 8 kbar pressures. We argue, that the migmatitic rocks are flowingabove the basement in form of ductile hot nappe which is progressively fed by lower

2. Mechanisms of lower crustal flow and its thermal and mechanical implications 21Figure 2.2: Schematic evolution of vertical displacement of folded single-layer duringbulk shortening for (a) “no erosion”, (b) “increasing erosion” and (c) “full erosion”models.Figure 2.3: Schematic geological cross-section showing main features of continentalchannel flow along eastern margin of Bohemian Massif.crustal material as the basement promontory is progressively wedging the thickenedorogenic root. We discuss the characteristics of the deformation fabrics in the hot nappeand defined the metamorphic and structural pattern as a result of the crustal scalechannel flow (Schulmann et al., 2008a).I have performed series of numerical models in order to answer the questions of theheat source origin, evolution of transitional rheology of the lithosphere and existence

2. Mechanisms of lower crustal flow and its thermal and mechanical implications 22Figure 2.4: Results of numerical simulations showing distribution of lithologies andtemperature field developed after 20 Ma. First correspond to low radioactive heatproduction, while second, third and fourth columns shows results of simulations withhigh heat production and different solidus for lower crustal layer.of gravitational instabilities controlling vertical extrusion or channel flow mechanismLexa et al. (2011). In this work we show that the thermal structure of the orogenis controlled by a position of felsic lower crust between the MOHO and the formermafic lower crust resulting from lower Paleozoic underplating. This layer, likely richin radioactive elements, is responsible for elevation of geotherm up to solidus of felsiccrust leading to its dramatic softening. Finally, the Rayleigh-Taylor instability developsleading to exhumation of felsic partially molten material into mid-crustal levels (Fig. 2.4).In this study we discuss as whether: 1) the felsic crust had a proportion of radioactiveelements high enough, or 2) as whether the heat source originate from below either from

2. Mechanisms of lower crustal flow and its thermal and mechanical implications 23Figure 2.5: Results of gravity inversion modelling showing the typical distribution ofcrustal layers in the Bohemian Massif.metasomatised upper mantle, or 3) due to up-welling of hot asthenospheric material. It isimportant to note that realistic data of thickness of felsic lower crustal layer, hanging wallmafic lower crust and middle crust derived from geophysical data (Fig. 2.5), includingrealistic values of concentration of the radioactive elements in the felsic lower crust allowsus to “rapidly” generate large temperature anomaly, which is able to trigger the verticaldoming process by gravitational instability. This exhumation mechanism is driven bothby gravity and by the horizontal shortening and we call it “Laterally Forced Overturns”process. Relics of this felsic lower crust are still preserved in the Moldanubian domainin the depth of 40 to 30 km (Guy et al., 2011), while the dense mafic crust occur inmiddle crustal depths (Fig. 2.5).Finally, the remaining problem is the understanding of influx of the felsic crust intothe root. We studied this problem in the Western Carpathians on the example of Vepordome (Jeřábek et al., 2008, 2007) and Orlica-Śnieżnik dome in West Sudetes (Chopinet al., 2011a). The goal of these studies is understanding of formation of burial orogenicfabrics during continental collision processes. Petrological studies of deep gneissesfrom the Vepor basement suggest that the origin of first subhorizontal metamorphicfabric is clearly related to the prograde P-T path (prograde zonation of garnets in Vepororthogneiss and pelites), and therefore to progressive burial. This mechanism is inobvious contrast to the generally accepted model of flat fabric development during theexhumation, namely the origin of “core complexes” due to orogen parallel extension. I

2. Mechanisms of lower crustal flow and its thermal and mechanical implications 24Figure 2.6: Petrological data showing distribution of different prograde P-T pathswithin crustal column. Schematic drawing of thickening from above model showingprogressive evolution of thermal structure in Vepor unit.tried to resolve the paradox, i.e. the development of flat fabric during crustal thickeningby numerical calculations and I offered a model of thermal evolution within thickenedcrust controlled Gemer crystalline overthrust (Jerabek,2009). This process of “thickeningfrom above” plays an essential role for the thermal evolution of a root zone, whichis supported by petrological data. In conclusion, model of thickening of Vepor crustindicates that successful conjunction of petrological, structural and numerical studies isonly possible if they are underpinned by thorough field knowledge.2.1 Perspectives for the numerical modelling of channelflow in the orogens and development of infra- and superstructuretransition zoneExtrusion of the lower crust and its lateral spreading below brittle-ductile transitionrequires a thorough understanding of the rheological and thermal states prevailing duringthe process. First, it is necessary to handle the problem of heat source which may beof various origins. It is likely, that key role is played by process of “relamination” offelsic lower crust above the MOHO resulting from massive material influx as shown in

2. Mechanisms of lower crustal flow and its thermal and mechanical implications 25Figure 2.7: Example of finite element modelling of crustal indentation of orogenicroot. This model was calculated using modified Elmer software (Maierová, 2011) andincludes temperature and strain rate dependent rheology, erosion, isostasy and gravitationalforces to simulate more realistic geodynamic evolution.our study from West Carpathians. To understand this process we need to examine avariety of heat sources such as the mantle delamination, radiocative mantle or radioactivelower crust. We will orient our research in future to understand relative role of variableheat sources, the rheology of crust, kinematics of gravity overturns and more realisticgeometries of rock bodies.Horizontal flow of orogenic lower crust over the Brunia indenter along eastern marginis a problem that also requires the rheological and thermal scaling. For this purpose weset up finite element models (Fig. 2.7), which could solve the problem of material transferefficiency of channel flow, velocity and strain-rate field patterns as well as temperature

2. Mechanisms of lower crustal flow and its thermal and mechanical implications 26relationship between the channel, overburden rigid lid and underlying basement. We testvarious combinations of rheological, thermal properties and boundary conditions whichallow us to understand principal behaviour within parameter space and to identify firstorder factors, which may cause the development of the channel structure, such as thoseobserved along the eastern margin of the Bohemian Massif (FIg ISTRA). It turns outthat the temperature contrast between the continent and the root is not the majorparameter to initiate indentation process, but the rheology contrast of rocks, whichare partially molten within root zone and strong in the indenter represent the mainparameter controlling channel flow extrusion.The last topic that I consider to be very promising is relationship between infraandsuperstructure of orogens in terms of their contrasting structural, rheological andthermal evolution. It is the model of Culshaw et al. (2006) which first introduced theconcept of mechanical coupling during early stages and decoupling during late stagesof orogen evolution. It is during the transition from coupled to decoupled behaviour ofthe crust when the infra and superstructure are formed and when the infrastructure andsuperstructure transition zone (ISTZ) is established. The coupling/decoupling and ISTZformation are clearly associated with extreme localization of deformation accompaniedby major change in mechanical properties of crustal rocks. Therefore in the light ofmechanical reasons for the ISTZ formation, several important questions related to themodel of Culshaw et al. (2006) emerge. (1) Is the origin of ISTZ exclusively controlledby the thermal structure of an orogen or can it also be influenced by other rheologicalparameters such as composition, fluids, differential stress, lithostatic pressure and strainrate? And if yes, to what extent? (2) What is the exact depth of ISTZ break throughin orogens with typically complex internal orogenic architecture? And can we predictits depth in active orogens?In this area, I would like to put efforts of my future master’s students and doctoralcandidates. I believe that areas like contact of Orlica-Śnieżnik dome with Zábřeh crystalline,contact of Teplá crystalline and Barrandian Proterozoic in Bohemian Massif andthe relationship of Vepor dome with Gemer Paleozoic in the Western Carpathians arebest candidates to solve these problems. Relationship between long-lasting structuralrecord of superstructure and juvenile infrastructure (possibly influenced by channel flowprocess) is first-order problem of continental structural geology and in combination withappropriate geochronological and petrological methods promises to bring new resultsnot only for Bohemian Massif and Carpathians but also for other orogenic systems likeCanadian Cordillera, Grenville tract and others. This research should be linked with adetailed microstructural studies combined with thermomechanical modelling allowing acombination of field and laboratory research which may be attractive for students withrather limited mathematical background.

2. Mechanisms of lower crustal flow and its thermal and mechanical implications 272.2 Accompanying publicationsSchulmann, K., Lexa, O., Štípská, P., Racek, M., Tajčmanová, L., Konopásek, J.,Edel, J.-B., Peschler, A., Lehmann, J., 2008a. Vertical extrusion and horizontalchannel flow of orogenic lower crust: key exhumation mechanisms in large hotorogens? Journal of Metamorphic Geology 26 (2), 273–297. 133Schulmann, K., Konopásek, J., Janoušek, V., Lexa, O., Lardeaux, J.-M., Edel, J.-B., Štípská, P., Ulrich, S., 2009b. An Andean type Palaeozoic convergence in theBohemian Massif. Comptes Rendus Geosciences 341 (2-3), 266–286. 159 Lexa, O., Schulmann, K., Janoušek, V., Štípská, P., Guy, A., Racek, M., Jan. 2011.Heat sources and trigger mechanisms of exhumation of HP granulites in Variscanorogenic root. Journal of Metamorphic Geology 29 (1), 79–102. 181Franěk, J., Schulmann, K., Lexa, O., Tomek, v., Edel, J.-B., 2011a. Model ofsyn-convergent extrusion of orogenic lower crust in the core of the Variscan belt:implications for exhumation of high-pressure rocks in large hot orogens. Journalof Metamorphic Geology 29 (1), 53–78. 2052.3 Related publicationsSchulmann, K., Kröner, A., Hegner, E., Wendt, I., Konopásek, J., Lexa, O., Štípská,P., 2005. Chronological constraints on the pre-orogenic history, burial and exhumationof deep-seated rocks along the eastern margin of the Variscan orogen,Bohemian Massif, Czech Republic. American Journal of Science 305 (5), 407–448.Franěk, J., Schulmann, K., Lexa, O., 2006. Kinematic and rheological model ofexhumation of high pressure granulites in the Variscan orogenic root: Exampleof the Blanský les granulite, Bohemian Massif, Czech Republic. Mineralogy andPetrology 86 (3-4), 253–276.Racek, M., Štípská, P., Pitra, P., Schulmann, K., Lexa, O., 2006. Metamorphicrecord of burial and exhumation of orogenic lower and middle crust: a new tectonothermalmodel for the Drosendorf window (Bohemian Massif, Austria). Mineralogy andPetrology 86 (3-4), 221–251.Jeřábek, P., Faryad, W. S., Schulmann, K., Lexa, O., Tajčmanová, L., 2008. Alpineburial and heterogeneous exhumation of Variscan crust in the West Carpathians:insight from thermodynamic and argon diffusion modelling. Journal of the GeologicalSociety 165, 479–498.

2. Mechanisms of lower crustal flow and its thermal and mechanical implications 28Skrzypek, E., Štípská, P., Schulmann, K., Lexa, O., Lexová, M., 2011b. Progradeand retrograde metamorphic fabrics - a key for understanding burial and exhumationin orogens (Bohemian Massif). Journal of Metamorphic Geology 29 (4), 451–472.Skrzypek, E., Schulmann, K., Štípská, P., Chopin, F., Lehmann, J., Lexa, O.,Haloda, J., 2011a. Tectono-metamorphic history recorded in garnet porphyroblasts:insights from thermodynamic modelling and electron backscatter diffractionanalysis of inclusion trails. Journal of Metamorphic Geology 29 (4), 473–496.Chopin, F., Schulmann, K., Štípská, P., Martelat, J., Pitra, P., Lexa, O., Petri,B., 2011b. Microstructural and petrological evolution of a high pressure graniticorthogneiss during continental subduction (Orlica-Śnieżnik dome, NE BohemianMassif). Journal of Metamorphic Geology, submitted.Štípská, P., Chopin, F., Skrzypek, E., Schulmann, K., Lexa, O., Pitra, P., Martelat,J., Bollinger, C., 2011. Juxtaposition of eclogite-facies and mid-crustal rocksduring exhumation: relative role of erosion and crustal scale folding in a continentalwedge (Orlica-Śnieżnik dome, Bohemian Massif). Journal of MetamorphicGeology, submitted.

Chapter 3Quantitative analyses ofmetamorphic microstructures,their visualization and statisticsWithin the framework of microstructural analyses carried out at the Institute of Petrologyand Structural Geology I developed a code to characterise and quantify deformationand metamorphic microstructures. This early version allowed quantification of strainintensity using digitized grain boundaries similar to PAROR and SURFOR methods ofPanozzo (1983, 1984) ans was for the first time used in microstructural styudy of deformedmarbes by Ulrich 92. However, we have quickly recognized the potential of thetechnique and I have implemented a range of statistical methods to quantify mineralmicrostructures of polyphase and monophase recrystallized aggregates. It is in particularthe analysis of grain sizes combined with CSD approach which allows determiningsome kinetic characteristics of grain nucleation and growth. The second approach implementedroutinely in the PolyLX code is the analysis of grain boundary frequencieswhich allow evaluating the relative importance of surface energy of mineral aggregate inquantifying the degree of deviation of grain contact frequencies from random distribution.The textural analysis is a powerful, but underused tool of petrostructural analysis.Except acquirement of common statistical parameters, this technique can significantlyimprove understanding of processes of grain nucleation and grain growth,can bring insights on the role of surface energies or quantify duration of metamorphicand magmatic cooling events as long as appropriate thermodynamical data forstudied minerals exists. This technique also allows systematic evaluation of degree of29

3. Quantitative analyses of metamorphic microstructures 30preferred orientations of grain boundaries in conjunction with their frequencies. Thismay help to better understand the mobility of grain boundaries and precipitationsor removal of different mineral phases.The new open platform, object-oriented MATLAB ® toolbox PolyLX developedby myself, provide several core routines for data exchange, visualization and analysisof microstructural data, which can be run on any platform supported by MATLAB ® .Detailed descriptions of toolbox routines and methods of implementation of newtechniques could be find on dedicated webpage http://petrol.natur.cuni.cz/ ondro.The technique was first used in Lexa et al. (2005) where we examined the problemof microstructural maturity related to the duration of two metamorphic events whichoccurred at similar pressure and temperature conditions. We analysed samples of amphibolitesand tonalitic gneisses that were affected by ductile deformation associatedwith emplacement of thin Carboniferous granodiorite sill. This event was controlledby intense solid state recrystallization due to intense deformation triggered by heatingeffect of adjacent granodiorite sill. On the other hand we studied similar rocks (amphibolitesand tonalitic orthogneiss) that recrystallized during long lasting evolution ofCambro-Ordovician intracontinental rift. This event reveals much longer thermal timescale compared to the Carboniferous event for surprisingly similar PT conditions. The

3. Quantitative analyses of metamorphic microstructures 31Figure 3.1: N 0 -Gt plot showing CSD evolution of plagioclase and hornblende. Thisfigure showing two contrasting positions for metamorphic events with different durations.Long-lived mature microstructures are grouped along so-called fabric attractor.microstructural analysis covered digitization of large number of samples which weresubsequently handled using PolyLX software. As a result of this analysis we discoveredthat the Carboniferous event was associated with grain size distribution whichin CSD space shown dominance of nucleation process with respect to slow growth rate.Increase of ambient temperature correlates with decrease in nucleation density accompaniedwith increase growth (Fig. 3.1). In contrast the Cambro-Ordovician rocks revealedpre-dominance of growth process over the nucleation. Our data plot on continuous curvein the CSD space which clearly indicates absence of Ostwald ripening and dominance oftextural coarsening by DeHoff’s communicating neighbours theory, which governs thechange in shape of CSD curve. The analysis show that the time matters in mineral microstructuresleading to development of grain size distribution which can be consideredas a grain size attractor (Fig. 3.1).This study had shown for the first time the power of grain contact frequency analysis.The rocks deformed by short lived Carboniferous pulse reveal aggregate distributionof both mafic and felsic phases. However, the similar lithologies experiencing long lastingCambro-Ordovician event recorded evolution driving contact frequencies towardsthe field of regular distribution. The process of increase of unlike contacts in rocks is

3. Quantitative analyses of metamorphic microstructures 32interpreted as mechanism of decrease of interfacial surface energy which is not only governedby mineral growth and development of triple point network as suggested by otherauthors. These differences were also underlined by strong shape preferred orientationcoupled with high aspect ratios for short lived deformation event while the protractedheating was characterized by almost random microstructure in terms of shapes of grainand their orientation. Moreover accompanied systematic microstructural analysis ofgabbros deformed at different thermal conditions of 600°C and 700°C (Baratoux et al.,2005b) shown that identical lithologies reveal different CSD patterns, shape preferredorientations and grain contact frequencies which can be attributed to contrasting deformationmechanisms governing rock deformation. It was shown that the temperature isa leading factor controlling processes like mechanical mixing of phases, growth versusnucleation rates and internal strain partitioning within rock. We had shown in this workthat the mineral microstructure has a potential to reflec thermal conditions better thanany other thermometric method.Finally, we examined the microstructural characteristic of amphibolites deformedat different metamorphic grade and subsequently folded by independent deformationevent (Baratoux et al., 2005a). This study showed the power of PolyLX method inevaluating the degree of mechanical anisotropy of rocks in terms of mineral elongation,degree of preferred orientation of grain boundaries and their spatial distribution. Itwas demonstrated that the banded amphibolites with high degree of shape and grainboundary preferred orientation, important aggregate distribution and highly contrastinggrain size reveal high mechanical anisotropy while material that shows low aspect ratio,shape and grain boundary preferred orientation in conjunction with regular grain contactdistribution and similar grain size distributions of phases of contrasting rheologies revealsvery low mechanical anisotropy. These two contrasting mineral architectures than governfurther behaviour of rock when external stress is applied leading to almost ideal bucklingin the first case and passive flow folding in the latter.The PolyLX was successfully applied to determine deformation mechanisms of quartzin study of Jeřábek et al. (2007) where we used the technique for paleostress analysisof recrystallized quartz for determination of flow stress in the Vepor orthogneiss. Themethod revealed high reproducibility of datasets and precision in grain size determinationleading to highly precise stress estimates which are impossible to be achieved bystandard stereometric measurements. This is because the PolyLX involves the solutionof 3D sectional problem which is commonly ignored in similar studies. In combinationwith temperature estimates we were able to produce a stress-temperature deformationmap, which shown positive correlation of stress and temperature for relatively high

3. Quantitative analyses of metamorphic microstructures 33Figure 3.2: Digitized microstructure images of highly anisotropic low temperatureamphibolite fabric and weakly anisotropic high temperature amphibolites.grade samples that was explained in term of progressive burial and temperature increase.Negative correlation of stress and temperature was found for low grade samplesand was explained by stress/strain rate partitioning due to weak matrix effect.During subsequent years we concentrated our efforts on studies of partially moltenrocks especially orthogneisses. The PolyLX method was used by Závada et al. (2007)to study the microstructure and topology of melt seams related to mylonitic flow oforthogneiss deformed at high melt/fluid pressure. We have shown that at high meltpressure and low differential stress conditions the orthogneiss deformed at high temperatureconditions can yield in brittle manner involving so called cavitation process. ThePolyLX method allowed quantifying the topology of intragranular melt filled fracturesthereby providing decisive argument for unusual cavitation dominated grain boundarysliding. The diffusion creep accommodated by cavitation fracturing of feldspar aggregateproduces disproportionally lower ductility of feldspars compared to quartz, whichmodify our view on rheology of the orogenic lower crust.Our melt related study culminated in two papers where the PolyLX software playedthe decisive role. In partially molten orthogneiss from the Kutná Hora crystalline complexSchulmann et al. (2008b) shown that with increasing melt proportion the crystalsize distribution of main phases show protracted growth in comparison to reduced nucleationrate, while the interstitial phases revealed increasing degree of nucleation rate

3. Quantitative analyses of metamorphic microstructures 34Figure 3.3: Rf/phi plot quantifying the orientation and shape of intragranular meltfilledvoids in K-fedspar aggregate and 3d image of melt distribution in monomineralicaggregate.Figure 3.4: Grain contact frequency vs. brain boundary preferred orientation showingthe evolution of microstructure from highly aggregate to random and regular distributionswith increasing deformation.(Fig. 3.4). At the same time, the rock reveals first increase of aggregate distributionrelated with development of banded mylonite structure followed by complete mixing ofall phases related to influx of melt and the rheological collapse of the rock. We correlatedthese stages of orthogneiss deformation with increasing amount of melt pressure whichreduced effective stress at constant (and low) differential stress level conditions. In fact,the amount of very small melt proportion controls not only the topology of melt seamsbut also dramatically weaken rock at melt fraction for about 5%Hasalová et al. (2008a) used the PolyLX method to evaluate the role of melt infiltrationinto orthogneiss microstructure. This work shows surprising CSD evolution from

3. Quantitative analyses of metamorphic microstructures 35Figure 3.5: N 0 -Gt plots showing evolution of CSD from growth dominated to nucleationdominated microstructural stages within crustal channel. This evolution isinterpreted as result of syntectonic melt infiltration.originally high Gt/N 0 values towards low values in conjunction with increasing amountof mineral resorption and mineral overgrowth features. This particular evolution canbe only explained by increasing importance of nucleation rate which is incompatiblewith process of in-situ melting but reflect crystallization of melt in rock pores. TheseCSD plots were therefore used as a major argument for reactive porous flow in felsicrocks, a concept which was for the first time demonstrated in the continental crust. TheCSD results are supported by evolution in grain contact frequency plot which revealssystematically increasing importance of regular grain distributions which is driven tounexpected values. The regular grain distribution simply indicates progressive growthof new phases in the rock aggregate as the melt crystallizes in the rock. Finally, theprogressive melt infiltration is related with loss of grain shape preferred orientation coupledwith loss of aspect ratios, which in turn is connected with increasing developmentof preferred orientation of certain unlike boundaries. All that indicates that the porousflow was a dynamic process related to dynamic dilation of intragranular pores whichmaintained the porosity of rock at high level as long as the reactive porous flow wasactive.Our last studies are focused on quantification of microstructural evolution of granulitetextures (Franěk et al., 2011b). We discovered precursor rock of felsic granuliteswhich is interpreted as coarse grained orthogneiss, that was converted to granulite duringCarboniferous thickening and collision. These studies show that the hypersolvusalkaline feldspar is first decomposed into thick perthites that are rapidly replaced bypolycrystalline, equidimensional aggregate of pure K feldspar and plagioclase (Fig. 3.6).This decomposition is driven by heterogeneous nucleation process resulting from stored

3. Quantitative analyses of metamorphic microstructures 36Figure 3.6: Example of digitized microstructures originated from heterogenous decompositionof alkaline feldspar due to increase of temperature.strain energy in the aggregate. Such a strain energy is driven by decompression associatedwith slight cooling of perthitic aggregate and leads to development of particularmicrostructure which reduces both strain and surface energies of the aggregate. ThePolyLX played decisive role in the characterizing the process and it shows that the microstructurereveals highest nucleation density compared to growth in conjunction withhighest possible regular grain distribution which was recorded in metamorphic rocks sofar. This microstructure showed to be a precursor for granulite mylonitic fabric. ThePolyLX allows tracking the microstructure evolution of the rock and helps to define firstgrain boundary sliding diffusional process replaced subsequently by dislocation creepflow at lower temperatures.3.1 Perspectives for numerical modeling of deformationmicrostructures and their quantificationPolyLX method revealed high applicability in microstructural analysis thanks to its versatileand modular character allowing large amount of applications in paleopiezometry,crystal size distribution, grain contact frequencies and deformation mechanisms studiesaccompanied with texture analyses (Fig. 3.7). The potential is therefore in further useof the method in a variety of microstructural studies which can be combined with kineticgrow/nucleation models if the thermal history is known. For that purpose a coupling

3. Quantitative analyses of metamorphic microstructures 37Figure 3.7: Example of textural data obtained from EBSD analysed by MTEX andsubsequently imported into PolyLX to obtain microstructural quantitative datawith ELLE code is required. In recent collaboration with leaders in ELLE programming(Mark Jessell - Toulouse, Sandra Piazollo - Stockholm and Paul D. Bons - Tubingen)we achieved a full compatibility between both codes so, that now the microstructuralprocesses can be both simulated and quantified. In close future we plan to test simpledeformation - growth simulations for known thermal histories coupled with detailedstatistical analysis of rock microstructure. Our aim is to define parameters allowing toprecisely characterize the microstructure relationships to stress and temperature as it isdone in metallurgy.In nature factors controlling the active deformation mechanisms and hence the rheologyinclude those that are properties of the deforming material (e.g. composition andmineral assemblage, texture, grain size), and those that are imposed on the system fromoutside (confining pressure, differential stress, temperature and fluid pressure). Thereforewe plan coupling of precise identification of PT evolution of rock carried out throughThermocalc and Perple X codes in collaboration with leading petrologists in the field(Roger Powell (Melbourne), Richard White (Mainz) and Pavla Štípská (Strasbourg))in order to attribute to compositional characteristics of mineral microstructures moreor less precise PT conditions and PT trends. This approach will allow us to search formicrostructures that are unique for given PT conditions but also time of rock residencein given thermal conditions. The latter parameter will be quantified by analyses of interfacialangles that are functions of time of annealing, but also phase spatial distributionsand grain size evolutions.

3. Quantitative analyses of metamorphic microstructures 383.2 Accompanying publicationsLexa, O., Štípská, P., Schulmann, K., Baratoux, L., Kröner, A., 2005. Contrastingtextural record of two distinct metamorphic events of similar P-T conditions anddifferent durations. Journal of Metamorphic Geology 23 (8), 649–666. 231Baratoux, L., Schulmann, K., Ulrich, S., Lexa, O., 2005b. Contrasting microstructuresand deformation mechanisms in metagabbro mylonites contemporaneouslydeformed under different temperatures (c. 650°C and c. 750°C).; Deformationmechanisms, rheology and tectonics; from minerals to the lithosphere. Journalof Structural Geology 243 (4), 97–125. 249Závada, P., Schulmann, K., Konopásek, J., Ulrich, S., Lexa, O., 2007. Extremeductility of feldspar aggregates–melt-enhanced grain boundary sliding and creepfailure: rheological implications for felsic lower crust. Journal of Geophysical Research,B, Solid Earth and Planets 112, B10210, doi:10.1029/2006JB004820.279Schulmann, K., Martelat, J.-E., Ulrich, S., Lexa, O., Štípská, P., Becker, J. K., Oct.2008b. Evolution of microstructure and melt topology in partially molten graniticmylonite: Implications for rheology of felsic middle crust. Journal of GeophysicalResearch, B, Solid Earth and Planets 113, B10406, doi:10.1029/2007JB005508.295Hasalová, P., Schulmann, K., Lexa, O., Štípská, P., Hrouda, F., Ulrich, S., Haloda,J., Týcová, P., 2008a. Origin of migmatites by deformation-enhanced melt infiltrationof orthogneiss: a new model based on quantitative microstructural analysis.Journal of Metamorphic Geology 26, 29–53. 315Franěk, J., Schulmann, K., Lexa, O., Ulrich, S., Štípská, P., Haloda, J., Týcová, P.,2011b. Origin of felsic granulite microstructure by heterogeneous decomposition ofalkali feldspar and extreme weakening of orogenic lower crust during the Variscanorogeny. Journal of Metamorphic Geology 29 (1), 103–130. 3413.3 Related publicationsBarraud, J., 2006. The use of watershed segmentation and GIS software for texturalanalysis of thin sections. Journal of Volcanology and Geothermal Research 154 (1-2), 17–33.

3. Quantitative analyses of metamorphic microstructures 39Machek, M., Špaček, P., Ulrich, S., Heidelbach, F., 2007. Origin and orientationof microporosity in eclogites of different microstructure studied by ultrasound andmicrofabric analysis. Engineering Geology 89 (3-4), 266–277.Jeřábek, P., Stunitz, H., Heilbronner, R., Lexa, O., Schulmann, K., 2007. Microstructural- deformation record of an orogen-parallel extension in the Veporunit, West Carpathians. Journal of Structural Geology 29 (11), 1722–1743.Závada, P., Schulmann, K., Lexa, O., Hrouda, F., Haloda, J., Týcová, P., 2009. Themechanism of flow and fabric development in mechanically anisotropic trachytelava. Journal of Structural Geology 31 (11), 1295–1307.Skrzypek, E., Schulmann, K., Štípská, P., Chopin, F., Lehmann, J., Lexa, O.,Haloda, J., 2011a. Tectono-metamorphic history recorded in garnet porphyroblasts:insights from thermodynamic modelling and electron backscatter diffractionanalysis of inclusion trails. Journal of Metamorphic Geology 29 (4), 473–496.Chopin, F., Schulmann, K., Štípská, P., Martelat, J., Pitra, P., Lexa, O., Petri,B., 2011b. Microstructural and petrological evolution of a high pressure graniticorthogneiss during continental subduction (Orlica-Śnieżnik dome, NE BohemianMassif). Journal of Metamorphic Geology, submitted.Oliot, E., Lexa, O., Schulmann, K., Goncalves, P., Marquer, D., 2011. Mid-crustalstrain localization in granitic rocks: constraints from quantitative textural andcrystallographic preferred orientations analyses. Tectonophysics, submitted.Martelat, J.-E., Malamoud, K., Cordier, B., Randrianasolo, B., Schulmann, K.,Lardeaux, J.-M., 2011. Garnet crystal plasticity in the continental crust, new examplefrom south Madagascar. Journal of Metamorphic Geology, submitted.

BibliographyAerden, D., 1998. Tectonic evolution of the Montagne Noire and a possible orogenicmodel for syncollisional exhumation of deep rocks, Variscan belt, France. Tectonics17 (1), 62–79.Baratoux, L., Lexa, O., Cosgrove, J. W., Schulmann, K., 2005a. The quantitative linkbetween fold geometry, mineral fabric and mechanical anisotropy; as exemplified bythe deformation of amphibolites across a regional metamorphic gradient. Journal ofStructural Geology 27 (4), 707–730.Baratoux, L., Schulmann, K., Ulrich, S., Lexa, O., 2005b. Contrasting microstructuresand deformation mechanisms in metagabbro mylonites contemporaneously deformedunder different temperatures (c. 650°C and c. 750°C).; Deformation mechanisms, rheologyand tectonics; from minerals to the lithosphere. Journal of Structural Geology243 (4), 97–125.Barraud, J., 2006. The use of watershed segmentation and GIS software for texturalanalysis of thin sections. Journal of Volcanology and Geothermal Research 154 (1-2),17–33.Behrmann, J., 1988. Crustal scale extension in a convergent orogen: the Sterzing-Steinach mylonite zone in the Eastern Alps. Geodynamica Acta 2, 63–73.Burg, J.-P., 1999. Ductile structures and instabilities: their implication for Variscantectonics in the Ardennes. Tectonophysics 309 (1-4), 1–25.Chopin, F., Schulmann, K., Skrzypek, E., Lehmann, J., Dujardin, J., Martelat, J.-E., Lexa, O., Corsini, M., Edel, J.-B., Štípská, P., Pitra, P., 2011a. Crustal influx,indentation, ductile thinning and gravity redistribution in a continental wedge: keyprocesses in the evolution of a mantle gneiss dome (European Variscan belt). Tectonics,submitted.Chopin, F., Schulmann, K., Štípská, P., Martelat, J., Pitra, P., Lexa, O., Petri, B.,2011b. Microstructural and petrological evolution of a high pressure granitic orthogneissduring continental subduction (Orlica-Śnieżnik dome, NE Bohemian Massif).Journal of Metamorphic Geology, submitted.41

Bibliography 42Culshaw, N., Beaumont, C., Jamieson, R., 2006. The orogenic superstructureinfrastructureconcept: Revisited, quantified, and revived. Geology 34 (9), 733–736.Dewey, J. F., 1975. Finite plate evolution: some implications for the evolution of rockmasses at plate margins. American Journal of Science 275A, 260–284.Dumicz, M., 1979. Tectogenesis of the metamorphosed series of the Klodzko district: atentative explanation. Geologia Sudetica 14 (2), 29–46.England, P., Houseman, G., 1986. Finite strain calculations of continental deformation2. Comparison with the India-Asia collision zone. Journal of Geophysical Research,B, Solid Earth and Planets 91 (3), 3664–3676.Franěk, J., Schulmann, K., Lexa, O., 2006. Kinematic and rheological model of exhumationof high pressure granulites in the Variscan orogenic root: Example of the Blanskýles granulite, Bohemian Massif, Czech Republic. Mineralogy and Petrology 86 (3-4),253–276.Franěk, J., Schulmann, K., Lexa, O., Tomek, v., Edel, J.-B., 2011a. Model of synconvergentextrusion of orogenic lower crust in the core of the Variscan belt: implicationsfor exhumation of high-pressure rocks in large hot orogens. Journal of MetamorphicGeology 29 (1), 53–78.Franěk, J., Schulmann, K., Lexa, O., Ulrich, S., Štípská, P., Haloda, J., Týcová, P.,2011b. Origin of felsic granulite microstructure by heterogeneous decomposition ofalkali feldspar and extreme weakening of orogenic lower crust during the Variscanorogeny. Journal of Metamorphic Geology 29 (1), 103–130.Guy, A., Edel, J.-B., Schulmann, K., Tomek, v., Lexa, O., 2011. A geophysical modelof the Variscan orogenic root (Bohemian Massif): Implications for modern collisionalorogens. Lithos 124 (1-2), 144–157.Harland, W. B., 1971. Tectonic transpression in Caledonian Spitsbergen. GeologicalMagazine 108, 27–41.Hasalová, P., Schulmann, K., Lexa, O., Štípská, P., Hrouda, F., Ulrich, S., Haloda, J.,Týcová, P., 2008a. Origin of migmatites by deformation-enhanced melt infiltration oforthogneiss: a new model based on quantitative microstructural analysis. Journal ofMetamorphic Geology 26, 29–53.Hasalová, P., Štípská, P., Powell, R., Schulmann, K., Janoušek, V., Lexa, O., 2008b.Transforming mylonitic metagranite by open-system interactions during melt flow.Journal of Metamorphic Geology 26, 55–80.Houseman, G., England, P., 1986. Finite strain calculations of continental deformation1. Method and general results for convergent zones. Journal of Geophysical Research,B, Solid Earth and Planets 91 (3), 3651–3663.

Bibliography 43Jeřábek, P., Faryad, W. S., Schulmann, K., Lexa, O., Tajčmanová, L., 2008. Alpineburial and heterogeneous exhumation of Variscan crust in the West Carpathians:insight from thermodynamic and argon diffusion modelling. Journal of the GeologicalSociety 165, 479–498.Jeřábek, P., Stunitz, H., Heilbronner, R., Lexa, O., Schulmann, K., 2007. Microstructural- deformation record of an orogen-parallel extension in the Vepor unit, WestCarpathians. Journal of Structural Geology 29 (11), 1722–1743.Konopásek, J., Schulmann, K., Lexa, O., 2001. Structural evolution of the central part ofthe Krušné Hory (Erzgebirge) Mountains in the Czech Republic; evidence for changingstress regime during Variscan compression. Journal of Structural Geology 23 (9), 1373–1392.Konopásek, J., Schulmann, K., Lexa, O., 2003. Reply to comments by A. Krohe and A.P.Willner on ”Structural evolution of the central part of the Kru né Hory (Erzgebirge)Mountains in the Czech Republic—evidence for changing stress regime during Variscancompression”. Journal of Structural Geology 25 (6), 1005–1007.Koyi, H., Geoffrey Milnes, A., Schmeling, H., Talbot, C., Juhlin, C., Zeyen, H., 1999.Numerical models of ductile rebound of crustal roots beneath mountain belts. GeophysicalJournal International 139 (2), 556–562.Kratinová, Z., Ježek, J., Schulmann, K., Hrouda, F., Shail, R. K., Lexa, O., 2010.Noncoaxial K-feldspar and AMS subfabrics in the Land’s End granite, Cornwall: Evidenceof magmatic fabric decoupling during late deformation and matrix crystallization.Journal of Geophysical Research, B, Solid Earth and Planets 115, B09104,doi:10.1029/2009JB006714.Krohe, A., Willner, A. P., 2003. Discussion on: ”Structural evolution of the centralpart of the Krušné Hory (Erzgebirge) Mountains in the Czech Republic; evidence forchanging stress regime during Variscan compression” by J. Konopásek, K. Schulmannand O. Lexa. Journal of Structural Geology 25 (6), 1001–1004.Lehmann, J., Schulmann, K., Edel, J.-B., Hrouda, F., Lexa, O., Štípská, P., Haloda,J., 2011. AMS and structural record of granitoid sills emplacement during growthof continental gneiss dome: implications for exhumation of deep crust. Journal ofStructural Geology, submitted.Lexa, O., Cosgrove, J., Schulmann, K., 2004. Apparent shear-band geometry resultingfrom oblique fold sections. Journal of Structural Geology 26 (1), 155–161.Lexa, O., Schulmann, K., Hrouda, F., in prep. De-formation of magmatic bodies in hotorogenic channel. Journal of Geophysical Research, B, Solid Earth and Planets.Lexa, O., Schulmann, K., Janoušek, V., Štípská, P., Guy, A., Racek, M., Jan. 2011. Heat

Bibliography 44sources and trigger mechanisms of exhumation of HP granulites in Variscan orogenicroot. Journal of Metamorphic Geology 29 (1), 79–102.Lexa, O., Schulmann, K., Ježek, J., 2003. Cretaceous collision and indentation in theWestern Carpathians: View based on structural analysis and numerical modeling.Tectonics 22 (6), 1066, doi:10.1029/2002TC001472.Lexa, O., Štípská, P., Schulmann, K., Baratoux, L., Kröner, A., 2005. Contrastingtextural record of two distinct metamorphic events of similar P-T conditions anddifferent durations. Journal of Metamorphic Geology 23 (8), 649–666.Machek, M., Špaček, P., Ulrich, S., Heidelbach, F., 2007. Origin and orientation of microporosityin eclogites of different microstructure studied by ultrasound and microfabricanalysis. Engineering Geology 89 (3-4), 266–277.Martelat, J.-E., Malamoud, K., Cordier, B., Randrianasolo, B., Schulmann, K.,Lardeaux, J.-M., 2011. Garnet crystal plasticity in the continental crust, new examplefrom south Madagascar. Journal of Metamorphic Geology, submitted.Matte, P., Lancelot, J., Mattauer, M., 1998. The Montagne Noire Hercynian AxialZone is not an extensional metamorphic core complex but a compresional post-nappeanticline with an anatectic core. Geodinamica Acta 11, 13–22.McKenzie, D. P., Parker, R. L., 1967. The North Pacific: an example of tectonics on asphere. Nature 216 (5122), 1276–1280.Neubauer, F., Genser, J., Kurz, W., Wang, X., 1999. Exhumation of the Tauern window,Eastern Alps. Physics and Chemistry of the Earth, Part A: Solid Earth and Geodesy24 (8), 675–680.Oliot, E., Lexa, O., Schulmann, K., Goncalves, P., Marquer, D., 2011. Mid-crustalstrain localization in granitic rocks: constraints from quantitative textural and crystallographicpreferred orientations analyses. Tectonophysics, submitted.Panozzo, R., 1983. Two dimensional analysis of shape fabric using projections of linesin a plane. Tectonophysics 95, 279–294.Panozzo, R., 1984. Two dimensional strain from the orientation of lines in a plane.Journal of Structural Geology 6, 215–221.Platt, J. P., Vissers, R. L., 1989. Extensional collapse of thickened continental lithosphere:A working hypothesis for the Alboran Sea and Gibraltar arc. Geology 17 (6),540–543.Racek, M., Štípská, P., Pitra, P., Schulmann, K., Lexa, O., 2006. Metamorphic recordof burial and exhumation of orogenic lower and middle crust: a new tectonothermalmodel for the Drosendorf window (Bohemian Massif, Austria). Mineralogy and

Bibliography 45Petrology 86 (3-4), 221–251.Ratschbacher, L., Frisch, W., Neubauer, F., Schmid, S. M., Neugebauer, J., 1989. Extensionin compressional orogenic belts: The eastern Alps. Geology 17 (5), 404–407.Rey, P., Vanderhaeghe, O., Teyssier, C., 2001. Gravitational collapse of the continentalcrust: definition, regimes and modes. Tectonophysics 342 (3-4), 435–449.Rosenberg, C., Brun, J.-P., Gapais, D., November, 2004. Indentation model of the EasternAlps and the origin of the Tauern Window. Geology 32 (11), 997–1000.Sanderson, D. J., Marchini, W., 1984. Transpression. Journal of Structural Geology6 (5), 449–458.Schulmann, K., Edel, J.-B., Hasalová, P., Cosgrove, J., Ježek, J., Lexa, O., 2009a. Influenceof melt induced mechanical anisotropy on the magnetic fabrics and rheology ofdeforming migmatites, Central Vosges, France. Journal of Structural Geology 31 (10),1223–1237.Schulmann, K., Konopásek, J., Janoušek, V., Lexa, O., Lardeaux, J.-M., Edel, J.-B.,Štípská, P., Ulrich, S., 2009b. An Andean type Palaeozoic convergence in the BohemianMassif. Comptes Rendus Geosciences 341 (2-3), 266–286.Schulmann, K., Kröner, A., Hegner, E., Wendt, I., Konopásek, J., Lexa, O., Štípská, P.,2005. Chronological constraints on the pre-orogenic history, burial and exhumation ofdeep-seated rocks along the eastern margin of the Variscan orogen, Bohemian Massif,Czech Republic. American Journal of Science 305 (5), 407–448.Schulmann, K., Lexa, O., Štípská, P., Racek, M., Tajčmanová, L., Konopásek, J., Edel,J.-B., Peschler, A., Lehmann, J., 2008a. Vertical extrusion and horizontal channel flowof orogenic lower crust: key exhumation mechanisms in large hot orogens? Journal ofMetamorphic Geology 26 (2), 273–297.Schulmann, K., Martelat, J.-E., Ulrich, S., Lexa, O., Štípská, P., Becker, J. K., Oct.2008b. Evolution of microstructure and melt topology in partially molten graniticmylonite: Implications for rheology of felsic middle crust. Journal of GeophysicalResearch, B, Solid Earth and Planets 113, B10406, doi:10.1029/2007JB005508.Schulmann, K., Thompson, A. B., Lexa, O., Ježek, J., 2003. Strain distributionand fabric development modeled in active and ancient transpressive zones. Journalof Geophysical Research, B, Solid Earth and Planets 108(B1), 2023, doi:10.1029/2001JB000632.Skrzypek, E., Schulmann, K., Štípská, P., Chopin, F., Lehmann, J., Lexa, O., Haloda, J.,2011a. Tectono-metamorphic history recorded in garnet porphyroblasts: insights fromthermodynamic modelling and electron backscatter diffraction analysis of inclusiontrails. Journal of Metamorphic Geology 29 (4), 473–496.

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Accompanying publications47

Journal of Structural Geology 23 2001) 1373±1392www.elsevier.com/locate/jstrugeoStructural evolution of the central part of the KrusÏne hory Erzgebirge)Mountains in the Czech RepublicÐevidence for changing stress regimeduring Variscan compressionJirÏõ KonopaÂsek a,b, *, Karel Schulmann a , Ondrej Lexa aa Charles University, Institute of Petrology and Structural Geology, Faculty of Science, Albertov 6, 128 43, Praha 2, Czech Republicb Geophysical Institute, Czech Academy of Science, BocÏnõ II/1401, 141 31, Prague 4, Czech RepublicReceived 12 July 2000; revised 17 December 2000; accepted 19 December 2000AbstractIn the central part of the KrusÏne hory Erzgebirge) Mountains, the parautochthonous metasedimentary basement has been overthrust by acrustal nappe of ®ne- and coarse-grained orthogneisses. The thrust boundary is de®ned by the presence of ma®c eclogites with preservedsubduction-related fabric. Westward thrusting of an allochthonous unit is associated with the development of the main metamorphic foliationand lineation in non-eclogitic lithologies. Buttressing of the allochthonous body fromthe west is responsible for the development of late,large-scale folds with N±S trending hinges and vertical axial planes. Subsequent N±S compression leads to large-scale folding of both theparautochthonous and allochthonous units. This deformation produces km-scale antiforms with hinges plunging to the west and is associatedwith the development of the E±W stretching lineation as a result of complete reworking of earlier fabric in the limb zones. N±S shortening isalso associated with the development of small-scale folds and brittle-ductile kink bands suggesting a decrease in temperature, and, thus, upliftof the whole studied area during this event. The last stage of deformation is characterised by the development of kink-band folds and acrenulation cleavage. These structures suggest a sub-vertical direction of principal compression, developed exclusively in those parts of thearea in which the N±S compression produced steep planar fabric. q 2001 Elsevier Science Ltd. All rights reserved.1. IntroductionIn the western part of the Saxothuringian domain of theBohemian Massif, the identi®cation of the major tectonicunits is relatively easy because of considerable metamorphiccontrast between high-grade crystalline rocks andadjacent low-grade metasediments. Here, allochthonoushigh-grade crystalline units represented by the MuÈnchberg,Frankenberg and Wildenfels klippens are thrust over lowgradeCarboniferous sediments Franke et al., 1995).Another important structural pattern is associated withexhumation of the Saxonian granulites. The emplacementof this body into the upper crust is interpreted in terms of theextension associated with the development of a metamorphiccore complex Franke, 1993).In the eastern part of the Saxothuringian domain centraland eastern KrusÏne hory Mountains), the identi®cation ofthe major structural units is more dif®cult because of themedium- to high-grade metamorphism affecting both theallochthonous and autochthonous units. Moreover, at least* Corresponding author. Fax: 1420-2-2195-2238.E-mail address: kony@natur.cuni.cz J. KonopaÂsek).two periods of high-grade metamorphism have affected thepre-Palaeozoic basement in this area, and it is dif®cult, if notimpossible, to separate each other in the ®eld MlcÏoch andSchulmann, 1992; KroÈner et al., 1995). The tectonics of theKrusÏne hory Mountains is commonly interpreted as a resultof a large-scale westward oriented continental collisionMatte et al., 1990). This thrusting event is documentedby the presence of a ¯at foliation, an E±W trending lineationand commonly observed westward directed kinematicsRajlich, 1987; Matte et al., 1990; MlcÏoch and Schulmann,1992). Recently published studies show that this thrustingwas associated with the emplacement of a large-scalecrustal nappe over the Saxothuringian parautochthonKlaÂpova et al., 1998; Krohe, 1998). The main argumentfor the presence of an allochthonous unit is the widespreadoccurrence of ma®c eclogites surrounded by non-eclogiticrock assemblages KonopaÂsek, 1998; RoÈtzler et al., 1998;KlaÂpova et al., 1998; SchmaÈdicke et al., 1992). However,the exact boundary between the parautochthonous andallochthonous units is poorly documented. Closerexamination of mapped geological structures Hoth et al.,1994), as well as a detailed ®eld structural survey, showsthat the E±W thrusting is not the only Variscan tectonic0191-8141/01/$ - see front matter q 2001 Elsevier Science Ltd. All rights reserved.PII: S0191-814101)00003-749

1374J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392Fig. 1. a) Location of the KrusÏne hory Mountains in the European Variscides black square in the northern part of the Bohemian Massif). b) Simpli®edgeological map of the central part of the KrusÏne hory Mountains with the studied area outlined by a black line.event responsible for the building of the ®nal geologicalpattern.In this paper we document that the structural fabric associatedwith emplacement of a crustal nappe is later affectedby an important N±S compression, which is responsible forregional-scale refolding of a previously developed ¯at lyingfoliation. We discuss the succession of deformational eventsin conjunction with the thermal evolution and changingstress regimes through time. We also correlate small-scalestructures and observed ®nite strain ellipsoids with stressesoperating during the crustal-scale folding associated withexhumation of the middle crust.2. Geological settingThe Bohemian Massif represents the easternmost crystallinecomplex of the European Variscides. Its western part iscomposed of the Saxothuringian domain, which has beenalready recognised by Kossmat 1927) as a unit with adistinct lithological and stratigraphic evolution with respectto the more easterly lying unitsÐthe Tepla±Barrandian andthe Moldanubian domains.The area studied is situated in the central part of theKrusÏne hory Mountains Erzgebirge), which representthe easternmost termination of the Saxothuringian domain.50

J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392 1375Fig. 2. Synoptic geological map of the studied area simpli®ed after Sattran 1967) and Hoth et al. 1994)). The inset represents lithotectonic column. Solid black lines show the position of cross-sectionspresented in Fig. 3.51

1376J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392Fig. 3. a)±d) Cross-sections over the studied area with D2±D3 structural data represented by the lower-hemisphere equal area projection. Data are contoured at 1 £ uniformdistribution. See Fig. 2 for theposition of cross-sections.52

J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392 1377Fig. 3. continued)The KrusÏne hory Mountains forman antiform-like structurewith Proterozoic metamorphic rocks in its core and with aPalaeozoic metasedimentary cover. The age of metamorphismwas thought to be dominantly Variscan withdecreasing metamorphic conditions from the east to thewest Kossmat, 1925) Fig. 1).The core of this structure is represented mainly by amonotonous complex of micaschists with subordinate metagreywackesthe Osterzgebirge complex) and bymetapelites, orthogneisses and numerous bodies of eclogitesin the uppermost part the PrÏõÂsecÏnice complex) Hofmann etal., 1988). The metasedimentary cover consists of quartzites,micaschists and phyllites of Lower Paleozoic ageHoth et al., 1979). This subdivision is based exclusivelyon lithological arguments, however, and no structural andmetamorphic features of the area were considered.MlcÏoch and Schulmann 1992) suggested a Cadomianage of metamorphism for anatectic orthogneisses in thecentral part of the KrusÏne hory Mountains, which wereintruded by a large late-Cadomian porphyritic granite. ANeo-proterozoic age for this pluton 554 ^ 10 MaÐU/Pbzircon method), as well as that of surrounding migmatites,was obtained by KroÈner et al. 1995). The Cadomian structurewas subsequently reworked by the Variscan deformationand metamorphism that dominates in the entire areae.g. SchmaÈdicke et al., 1995; Kotkova et al., 1996; Werneret al., 1997; KroÈner and Willner, 1998).PT data fromall major lithologies suggest a high dP/dTgradient during Variscan metamorphism SchmaÈdicke et al.,1992; Holub and SoucÏek, 1994; KlaÂpova et al., 1998;KonopaÂsek, 1998; RoÈtzler et al., 1998). In the Germanpart of the KrusÏne hory Mountains, Willner et al. 1994)have proposed three major high-pressure HP) units, withdecreasing metamorphic grade from the lowermost to theuppermost unit. Krohe 1996) interpreted this feature as aresult of ductile extension during exhumation of thethickened Saxothuringian domain. All these units containma®c eclogites and some of them HP granulites Willner53

1378J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392et al., 1997) with peak PT conditions differing considerablyfrom those of the adjacent metasediments RoÈtzler et al.,1998).Metamorphic conditions in the study area are similar tothose de®ned in Germany. A high dP/dT gradient wasobserved in metasedimentary rocks KonopaÂsek, 1998,2001) which host ma®c eclogites. The peak PT conditionsof the eclogites are not consistent with those of the adjacentparautochthonous metasediments, however, suggesting thatthey have been juxtaposed tectonically KlaÂpova et al.,1998).3. Lithological zonation of the studied areaThe area studied involves four main large-scale structuresexposed in the vicinity of the village of MeÏdeÏnecÐtheMeÏdeÏnec synform, the Oberwiesenthal structure, theMeÏdeÏnec antiformand the KlõÂnovec antiformSÏkvor,1975) Fig. 2). These large structures are composed of thefollowing rocks:1. Plagioclase schists are described by KonopaÂsek 1998)as characteristic metasediments forming the Saxothuringianparautochthon. Typical samples containreversely zoned plagioclase porphyroblasts envelopingnumerous garnet inclusions. As these rocks show a polyphasemetamorphic history, they can be sampled atdifferent stages of their evolution KonopaÂsek, 1998).These rocks are intercalated with metagreywackes'Dichte Gneisse' of Pietzsch 1914)) and metaconglomeratesMehnert, 1939; Sattran, 1963).2. Orthogneisses appear either as ®ne-grained equigranularvarieties or as deformed porphyritic granite 'RoteGneisse' of Scheumann 1935)). Porphyriticorthogneisses are heterogeneously deformed and occurin all deformation stages ranging from protomyloniticmetagranites up to banded ultramylonites.3. Garnetiferous micaschists are characterised by theappearance of numerous garnet porphyroblasts, sometimesup to 2 cm in diameter, surrounded by a whitemica and quartz matrix. In the southern limb of theKlõÂnovec antiform, garnetiferous micaschists are associatedwith a layer of amphibolites Fig. 2).4. Ma®c eclogites bear a typical HP mineral assemblageconsisting of omphacite 1 garnet 1 zoisite 1amphibole 1 rutile 1 quartz and hand specimens arecharacterised by well-developed planar and linear fabricKlaÂpova et al., 1998).The rock-types described above occur in all the largescalestructures, but the lithological zonation within thesestructures is not uniform. In the MeÏdeÏnec synform, plagioclaseschists with metagreywackes represent structurally thedeepest level. These schists pass upward into the layer ofgarnetiferous micaschists associated with the ma®ceclogites. In the hanging wall of garnetiferous micaschist,there occurs a large slab of orthogneisses with a ®ne-grainedvariety at the bottomand a porphyritic type at the top Fig.3aÐnorthern part of the pro®le 1±1 0 ). Locally, fragmentsof the allochthonous orthogneiss slab can be observed alsowithin the parautochthonous micaschists Fig. 3b). Thelithological zonation of the KlõÂnovec antiformis morecomplicated. In its eastern part, the zonation of the southernlimb is the same as that in the MeÏdeÏnec synformFig. 3cÐsouthern part of the pro®le 3±3 0 ). However, farther to thewest, the lithotectonic zonation becomes progressivelyreversed with orthogneisses in the lowermost positionpassing in the garnetiferous micaschists and with plagioclaseschists in the uppermost position Fig. 3dÐsouthernpart of the pro®le 4±4 0 ). Moreover, in the core of theKlõÂnovec antiform, large bodies of eclogites are situated inthe centre of the orthogneiss body Fig. 2).Field observations show that ma®c eclogites, togetherwith a layer of garnetiferous micaschists, are in mostcases exposed along the boundary between the orthogneissesand plagioclase schists. The location of these unitsalong this boundary is critical for understanding the tectonicevolution of the studied area.4. PT evolution of the studied area and the de®nition ofthe allochthonous and parautochthonous domainsThe PT conditions of orthogneiss formation are notknown, but the stability of plagioclase in all the studiedsamples excludes the possibility of them being deformedunder eclogite facies conditions. The peak pressure conditionsof the plagioclase schists were established to be 13±15 kbar at a temperature of 580±6308C KonopaÂsek, 1998).These pressure conditions are in contrast with those reportedby KlaÂpova et al. 1998) fromthe ma®c eclogites wherepeak pressures were estimated to be 26 kbar at 650±7008C. The PT conditions of the associated garnetiferousmicaschists approach those of the eclogites 6408C,22 kbarÐ KonopaÂsek, 2001). The difference between PTestimates from parautochthonous metasediments and eclogiteswas explained in terms of emplacement of previouslysubducted oceanic crust into continental rocks during theearly stages of collision KlaÂpova et al., 1998).As noted earlier, the ma®c eclogite bodies associated withthe garnetiferous micaschists occur systematically at theboundary between plagioclase schists and orthogneisses.This observation suggests that this boundary represents amajor crustal boundary along which ma®c eclogies wereexhumed and incorporated into the middle crust. Accordingto the interpretation of KonopaÂsek 1998) and KlaÂpova et al.1998), the plagioclase schists represent a Saxothuringianparautochthon, which was overthrust by middle-crustalorthogneisses derived froman orogenic root domain,today exposed in the eastern Moldanubian zone. Based onthis interpretation, these allochthonous orthogneisses,54

J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392 1379together with ma®c eclogites and garnetiferous micaschistswill be referred to as the Lower Crystalline Nappe.5. The succession of deformation structures D1±D3)A polyphase tectonic evolution has resulted in four D1±D4) stages of deformation. The structures developed in theeclogites have been described by KlaÂpova et al. 1998) andare, therefore, not examined in detail here, except whennecessary for the interpretation of a particular deformationstage.5.1. D1 structuresThe D1 structures are present exclusively in the eclogitesand consist of syn-metamorphic S 1 foliation and an L 1lineation. As described by KlaÂpova et al. 1998), the S 1metamorphic foliation is mainly the result of a planarorientation of omphacite and a metamorphic layeringcharacterised by an alternation of omphacite-rich layerswith layers rich in garnet. The L 1 stretching lineation ischaracterised by a shape-preferred orientation of omphacitecrystals. The orientation of D1 structures in eclogiticboudins varies according to their position within largescalestructures over the studied area see KlaÂpova et al.1998) for details).5.2. D2 structuresThe D2 deformation in eclogites is marked by thedevelopment of asymmetric internal foliation boudinagewith neck zones ®lled with quartz, rutile, amphibole,paragonite and zoisite. Brittle cracks up to several meterslong have developed; they are either closed or ®lled with thesame assemblage as the neck zones KlaÂpova et al., 1998).The D2 structures in non-eclogitic lithologies can be bestobserved in the northern part of the studied area in theOberwiesenthal and MeÏdeÏnec synforms) where the rocksare only slightly affected by the D3 deformation. Largedomains of D2 fabric only slightly affected by D3 deformationalso appear in the basement metasediments exposedsouth of the KlõÂnovec antiform.In the orthogneisses, the S 2 foliation is characterised bythe alternation of recrystallised quartz and feldspar ribbonswith phyllosilicate rich domains, and in metasedimentsmainly by preferred orientation of micas and by alternationof mica-rich and quartz-rich layers. The orientation of the S 2foliation in the MeÏdeÏnec synformis variable but it generallystrikes N±S and dips gently between 10 and 208. In theOberwiesenthal structure, the S 2 foliation is ¯at lying inits northern part and becomes steeply inclined as onemoves south. The L 2 stretching and mineral lineation isde®ned by an alignment of quartz and feldspar aggregatesand by a stretching of quartz-rich aggregates in theorthogneisses, and as a preferred arrangement of micas inthe metasediments. The orientation of the L 2 lineation isgenerally ESE±WNW Fig. 4).Numerous kinematic indicators such as sigmoidalK-feldspar porphyroclasts and S±C fabrics in porphyriticorthogneiss Fig. 5a) and shear-bands in metasedimentsare consistent with top-to-the-west oriented shearing. Inseveral places, early F 2 isoclinal, recumbent folds wereobserved in basement rocks of semipelitic compositionFig. 5b). These early folds refold the primary compositionalbanding S 0 of the metasediments leading to itscomplete transposition into a penetrative metamorphicfabric. The orientation of early F 2 fold hinges is E±W,being parallel to the L 2 mineral lineation Fig. 4).During late stages of the D2 deformation, the wholeassembled sequence was folded into large-scale periclinalstructures further described as the late F 2 folds) withsubvertical, N±S trending axial planes. These late F 2periclines formthe present day MeÏdeÏnec synformand theOberwiesenthal structure.5.3. D3 structuresIn the southern termination of the MeÏdeÏnec synform, theS 2 fabric dips gently to the NW. As one moves south, theorientation of the metamorphic fabric and the lithologicalbodies rotate so that the foliation dips steeply to the southFigs. 3a and 4). This geometry suggests that the gentlydipping planar fabric of the MeÏdeÏnec synformis refoldedby the large MeÏdeÏnec antiformwith the hinge zone plungingto the west at a shallow angle. This fold is asymmetricalwith a steeply dipping southern limb, a gently dippingnorthern limb and a hinge zone plunging gently to the west.A similar pattern is developed in the Oberwiesenthalstructure and the KlõÂnovec antiform. The Oberwiesenthalstructure has, in its eastern part, a similar geometry to theMeÏdeÏnec synformand shows progressive decrease of theinterlimb angle to the south Fig. 4). The connectionbetween the Oberwiesenthal synformand the KlõÂnovec antiformiseroded and the later structure represents, as in thecase of the MeÏdeÏnec structures, a former N±S trendingsynclinal structure rotated into an E±W direction Fig. 2).In the E±W trending limb, the foliation is subverticalFig. 5c) and the hinge of this large fold is steeply plungingto the SW. This is documented by the linear fabric of ma®ceclogites, which plunges steeply in the same direction. Theappearance of eclogite bodies in the centre of the KlõÂnovecantiform, as well as double thickness of the ®ne-grainedorthogneisses in the northern zone of the KlõÂnovec antiformFig. 2) suggest that the northern part of this E±W trendinglimb was thickened during the the D2 thrusting episode.The mineral lineation in the Oberwiesenthal structure, theMeÏdeÏnec synformand the MeÏdeÏnec antiformshows auniformWNW±ESE orientation Fig. 4), consistent withother parts of the KrusÏne hory Mountains parautochthone.g. MlcÏoch and Schulmann, 1992). In the KlõÂnovecantiform, the mineral and stretching lineation of the55

1380J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392Fig. 4. Structural map of the studied area showing trends of D1, D2 and D3 planar and linear structures.56

J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392 1381Fig. 5. a) S±C fabric in porphyritic orthogneiss of the Lower Crystalline nappe. This fabric originated during D2 and suggests top-to-the-west oriented senseof movement. b) Isoclinal F 2 fold in metagreywackes of the parautochthonous unit. c) Vertical S 2 foliation in the southern limb of the KlõÂnovec antiform. Thisvertical fabric resulted fromrefolding of originally ¯at S 2 foliation during D3. d) F 3 fold with steep axial plane in the autochthonous plagioclase schists southof the KlõÂnovec antiform. e) Conjugate system of D3 kink bands affecting the S 2 foliation in the hinge zone of the MeÏdeÏnec antiform. f) Late F 4 folds withsubhorizontal axial planes in amphibolite of the southern limb of the KlõÂnovec antiform.orthogneisses is systematically plunging gently to the WFig. 4), whereas in the eclogites, the omphacite L 1 lineationis gently plunging to the W±WSW in the central part of theKlõÂnovec antiformand steeply to the SW in the hinge zone.Numerous D3 fold structures up to several metres in sizewith E±W trending subhorizontal hinge zones and steep tointermediate, mostly northward dipping axial planes occurin the parautochthonous plagioclase schists south of theKlõÂnovec antiformFigs. 4 and 5d).The D3 deformation is most intense in the parautochthonousplagioclase schists in the area between the northernlimb of the KlõÂnovec antiformand the southern limb of57

1382J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392Fig. 6. Schematic block-diagram of the studied area. Flat lying S2 foliation can be observed in the northern part of the studied area in the MeÏdeÏnec synform. Going to the south, the early D2 fabric is completelyreworked by D3 deformation. The Oberwiesenthal structure is not shown.58

J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392 1383the MeÏdeÏnec antiform. In this area, the S 2 fabric is completelyreworked into a steep to subvertical E±W trendingS 3 foliation. Here, a strong sub-horizontal stretching ofquartz aggregates can be observed in several places. In thehinge zone of the MeÏdeÏnec antiform, monoclinal or conjugatekink bands are developed affecting the early S 2 metamorphicfoliation of the garnetiferous micaschist andorthogneisses. These kink bands exhibit E±W trendingaxes and kink planes, bimodally distributed with respectto ¯at-lying S 2 foliation Fig. 5e). The geometry of thesekink bands suggests that they are genetically linked to theMeÏdeÏnec antiformal D3 structure but the mechanism and thekinematics of kinking will be discussed in a separate sectionbelow.The general structure of the studied domain is shown inblock-diagram Fig. 6), summarising lithological and structuralobservations discussed above as an effect of the D3folding superimposed on the D2 fabrics.6. Problem of the L 2 and L 3 lineation solved usingaggregate-shape analysisAn important feature is the parallelism of the L 2 minerallineation in the MeÏdeÏnec synforma D2 structure) with thelineation observed in micaschists and orthogneisses of boththe MeÏdeÏnec and the KlõÂnovec antiforms. If the WNW±ESEtrending L 2 lineation in the MeÏdeÏnec synformwere rotatedby late F 2 folding, the originally EW trending, gentlyplunging L 2 lineation would become subvertical Fig. 7).This can be observed in N±S trending, steep S 2 foliationin the southern part of the Oberwiesenthal structure. Subsequentrotation of the S 2 foliation by F 3 folds with a steephinge zone will keep the orientation of the L 2 in a subverticalposition Fig. 7) . The lineation in the steep limbsof the KlõÂnovec and the MeÏdeÏnec antiforms is always foundto plunge gently to the west, however, indicating that it wasformed during the D3 deformation L 3 lineation). Consequently,the change of the L 2 into an L 3 lineation must beassociated with complete modi®cation of the D2 fabricduring the D3 deformation. In order to investigate theproposed changes to the D2 fabric ellipsoid during the D3deformation, we have carried out a shape analysis of mineralaggregates in the orthogneisses fromthe MeÏdeÏnec synformand the MeÏdeÏnec and the KlõÂnovec antiforms.In order to determine the role of the F 3 folding inpossible fabric modi®cations, an aggregate-shape analysisof K-feldspar and quartz was carried out on the porphyriticorthogneisses. Selected porphyritic augen orthogneisseswere sampled in the hinge zones and limbs of both theMeÏdeÏnec and the KlõÂnovec antiforms, as well as in theMeÏdeÏnec synformand the Oberwiesenthal structure. Twosections in each sample were studied: a) perpendicular tometamorphic foliation and parallel to mineral lineation theXZ section of the ®nite strain ellipsoid), and b) perpendicularto both foliation and lineation the YZ sectionof the ®nite strain ellipsoid). The shapes of the K-feldsparand quartz aggregates were traced on transparent sheet,digitised and then analysed. Because of the large size ofmeasured aggregates, only a limited number of measurementscould be carried out on each sample Table 1). Finitestrain ratios were calculated fromthe two principal strainellipsoid planes using a harmonic mean method Lisle,1977; Ramsay and Huber, 1983, p. 80). These ratios wereFig. 7. Succession of the lower-hemisphere equal area projections shows schematically expected evolution of the orientation of the L 2 lineation during D2 andD3 folding. Early L 2 lineation is trending E±W being subhorizontal. During late F 2 folding, the L 2 lineation does not change the orientation, but becomessteeper. The successive F 3 folding changes the orientation of steep S 2 foliation; the L 2 lineation should rotate into N±S direction maintaining its steep plunge.In the D3 structures the MeÏdeÏnec and the KlõÂnovec antiforms), however, the lineation is always E±W oriented and mostly subhorizontal.59

1384J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392Table 1Harmonic means of measured X/Z and Y/Z ratios and corresponding calculated X/Y ratios and k and d values for orthogneiss samples from the D2 domains theMeÏdeÏnec synformand the Oberwiesenthal structure) and the D3 domains the KlõÂnovec and MeÏdeÏnec antiforms). See Fig. 8 for corresponding Flinn's graphand the d/k plotSample Locality Object X/Z Y/Z X/Y k d Number of objects measured:in XZin YZD2MR10 MeÏdeÏnec synformQtz 5.907 2.26 2.614 1.28 2.047 18 25Kf 3.438 1.372 2.507 4.055 1.552 4 7MR7 MeÏdeÏnec synformQtz 4.914 1.865 2.634 1.888 1.849 24 35Kf 2.525 1.831 1.379 0.456 0.914 8 16MR8 MeÏdeÏnec synformQtz 4.594 2.34 1.963 0.719 1.65 22 21Kf 1.752 1.285 1.364 1.276 0.462 7 4OW407 Oberwiesenthal str. Qtz 3.614 3.033 1.192 0.094 2.042 11 9Kf 3.794 2.234 1.699 0.566 1.418 12 14OW417 Oberwiesenthal str. Kf 3.891 2.372 1.64 0.466 1.514 13 13D3MR115 KlõÂnovec antiformQtz 5.271 2.955 1.783 0.401 2.106 10 9Kf 3.171 1.913 1.657 0.72 1.125 10 8MR11 MeÏdeÏnec antiformKf 4.209 2.488 1.692 0.465 1.641 21 49MR12 MeÏdeÏnec antiformKf 10.45 9.001 1.161 0.02 8.002 20 23MR3 KlõÂnovec antiformQtz 10.26 6.301 1.628 0.118 5.338 23 13Kf 5.487 3.194 1.718 0.327 2.309 15 11MR401 KlõÂnovec antiformQtz 12.33 8.385 1.47 0.064 7.399 10 14Kf 6.961 5.168 1.347 0.083 4.182 12 13used to construct the k-parameter of Flinn 1962) andd-parameter e.g. Ramsay and Huber, 1983, p. 202) toestimate the shape of the strain ellipsoid and strain intensity,respectively.The ®nite strain analysis shows that the augenorthogneisses in the MeÏdeÏnec synform, as well as those inthe hinge zone of the MeÏdeÏnec antiform, exhibit ellipsoidsranging fromconstrictional to plane-strain or slightly oblatesymmetry, k-parameters vary from 0.72 to 1.89 for quartzand 0.46 to 4.06 for feldsparsÐFig. 8 and Table 1). Exceptionallyoblate fabric was observed for quartz aggregates inthe sample OW407 k ˆ 0.094). Quartz generally showshigher strain intensity d-parameters from 1.65 to 2) thanfeldspar d-parameters from 0.9 to 1.55). In studied samples,the orthogneisses have a character of augen orthogneiss withwell preserved K-feldspar porphyroclasts. In the southern,steep limb of the MeÏdeÏnec antiform, and in the orthogneissbodies of the KlõÂnovec antiform, the shapes of strainellipsoids are oblate k ˆ 0.4±0.06 for quartz andk ˆ 0.72±0.02 for K-feldsparÐFig. 8 and Table 1). As inthe previous case, quartz generally shows higher strainintensity d-parameter is from 2.1 to 7.4) than feldspard-parameter from 1.1 to 8). The increase of deformationintensity is connected with the disappearance of augenstructure and development of banded orthogneiss. It isnoted that samples with weak oblate symmetry show welldevelopedL 3 lineation characterised by stretching of feldsparand quartz aggregates. In contrast, samples with strongoblate symmetry do not show any stretching lineation andthe L 3 fabric is characterised only by alignment of whitemica in the foliation plane.We suggest in agreement with the mesoscopic structuraldata) that the symmetry of the D2 deformation is representedby a plane strain to constrictional or slightly oblatestrain ellipsoid resulting fromwestward directed noncoaxialshearing. This symmetry is preserved in both theMeÏdeÏnec synformand the Oberwiesenthal structure andprobably also in the hinge zone of the MeÏdeÏnec antiform.In contrast, the shapes of K-feldspar and quartz aggregatesin the limbs of the KlõÂnovec antiformshow oblate geometry,high strain intensity and X-axes oriented E±W. Asmentioned above, this type of geometry cannot be producedby passive rotation of the D2 fabric during the F 3 folding andwe propose that the large-scale F 3 folding is responsible forthe modi®cation of the shapes of the D2 strain ellipsoid. Theaggregate shape analysis also shows that the large scale F 3folding occurred under still elevated thermal conditions,which allowed ductile deformation of feldspar clasts.7. Brittle-ductile structures late D3 and D4)As reported by KlaÂpova et al. 1998), the D4 deformationin the eclogites resulted in the development of two main setsof structures. Most frequently, the D4 deformation produceslate S 4 retrograde shear zones, which crosscut the S 1eclogitic foliation at high angles. In the thick eclogiteboudins, the D4 deformation has resulted in the refoldingof the early S 1 foliation. There, D4 structures are open toclosed recumbent, asymmetrical F 4 folds. The D4 deformationis best recorded in amphibolites, garnetiferousmicaschists and basement plagioclase schists, and is60

J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392 13853X/Y2D2K=1D2D2D3Quartz - D2Quartz - D3K-feldspar - D2K-feldspar - D3D3D3D3dD211 2 3 4 5 6 7 8 910864Y/ZQuartz - D2Quartz - D3K-feldspar - D2K-feldspar - D3200 1 2 3 4 5kFig. 8. Flinn's graph and d/k plot obtained from measurements of K-feldspar and quartz aggregates in coarse-grained orthogneisses of the MeÏdeÏnec nappe. Inthe MeÏdeÏnec synformand northern part of the MeÏdeÏnec antiform D2 structures), both K-feldspar and quartz aggregates show plane strain to prolate symmetryat relatively low strain intensities. On the other hand, the same aggregates in the southern limb of the MeÏdeÏnec antiformand in the KlõÂnovec antiformD3structures) show oblate symmetry and higher strain intensities see text for methods and details). Grey areas represent starting and ending positions of the D2ellipsoids used for modelling of the D3 fabric. Results of this modelling are presented in Fig. 10.represented by late F 4 folds and kink bands Fig. 5f). Thelatest deformation in the metasediments is represented bysemi-brittle normal faults dipping to the SW. These occur onthe southern limb of the KlõÂnovec antiform. Evidence of theD4 deformation in orthogneisses is rare. The D4 structuresappear mostly in the southern limb of the KlõÂnovec antiformas discrete sets of brittle-ductile S 4 cleavage with homogeneousspacing obliquely oriented to the main S 2 foliationFig. 9).7.1. The D3 and D4kink-band folds in metasediments andorthogneissesIn several places, conjugate arrays of kink-band foldswere observed, but in most cases only one single set isdeveloped monoclinal kink-band folds after Ramsay andHuber 1987), p. 427). All these kink-band folds are ofcontractional type Ramsay and Huber, 1987, p. 427) andtheir common feature is an E±W orientation of their hinges.In the northern limb of the MeÏdeÏnec antiformclose to thehinge zone, Fig. 9) several localities with well-developedkink-band folds occur. These appear in both the metasedimentsand the orthogneisses in zones where the S 2 fabricdips gently Loc. 130, 130.V, 6, 138ÐFig. 9). The sametype of kink-band folds can be locally observed in zones of¯at lying S 2 metamorphic fabric south of the KlõÂnovec antiformLoc.307). These kink-band folds, which are eitherconjugate or forma single set with a uniformorientation ofthe kink-plane, are related to the D3 deformation.The D4 kink-band folds are developed exclusively inzones of steep S 2 ±S 3 fabric in both the plagioclase schistsand the garnetiferous micaschists, and are mostfrequently observed in close proximity to the allochthonousorthogneiss body. Three groups of D4 kink-band folds arerecognised and these are shown in Fig. 9. A group a) ismade up from a set of monoclinal kink-band folds, whichaffect foliations dipping steeply to the north. They indicatetop-to-the-north normal movement. Group b) consists of61

1386J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392Fig. 9. Structural map of the studied area shows D3 and D4 monoclinal and conjugate systems of kink bands. Orientation of the kink bands is presented in thelower-hemisphere equal area projection together with the orientation of S 2 ±S 3 foliation see lower left part of the ®gure for principal fabric elements). For eachlocality numbered inside or above each projection), all measured kink bands and foliations were averaged and presented as a mean value. Estimated s 1 and s 3directions are shown at localities with developed conjugate kink bands. The equal area projection on the right hand side of the ®gure represents poles to brittleductile,late D4 shear bands in orthogneisses and micaschists. Three groups of F 4 compressional kink bands observed in the studied area are shown in the lowerpart of the ®gure. a) Kink bands showing northward normal kinematics affecting northward dipping steep foliation. b) Conjugate set of kink bands affectingsubvertical foliation. c) Kink bands showing southward normal kinematics affecting southward dipping steep foliation.symmetrically developed conjugate kink-band folds andthese are developed in a vertical foliation. Group c)consists of monoclinal kink-band folds which deform asteep, south-dipping foliations and indicates top-to-thesouthnormal kinematics.In the southern limb of the KlõÂnovec antiform, most of thekink-band folds belong to group 3, with groups 2 and 1observed less frequently. On the southern limb of theMeÏdeÏnec antiform, the three groups of kink-band folds aremore equally developed.62

J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392 13877.2. Analysis of kink-band foldsThe geometry of symmetrical kink-band folds was usedto determine the orientation of stress axes according to themethod of Ramsay 1962a). The orientation of s 1 is perpendicularto the kink-band fold axis and lies in the plane thatbisects the obtuse angle between the conjugate kink-bandfolds. If only one set of the conjugate kink-band foldsdevelops, it implies that the maximum principal compressivestress s 1 was oblique to the foliation. The approximateorientation can be deduced but, as the angle betweens 1 and the kink-band fold is determined by the value of themechanical anisotropy Cobbold et al., 1971) and because ofthe phenomenon of stress de¯ection, it is not possible todetermine the exact orientation of s 1 .The sub-horizontal orientation of s 1 , deduced fromthegeometry and orientation of the D3 kink-band folds, isconsistent with earlier deduced N±S compression associatedwith the D3 deformation. Some of them show signsof later rotation, which probably occurred during theprogressive development of the D3 folding e.g. Locality6 in Fig. 9).The uniformorientation of the D4 fold axes and theessentially bimodal distribution of kink-band folds in thesouthern limb of the KlõÂnovec antiformFig. 9) indicatethat they probably represent a conjugate set groups a)and c), Fig. 9). This conjugate set is developed throughoutthe whole study area. Locally, only one set of the kink-bandfolds is found to have developed and, as discussed above,this can be interpreted as a result of variations of the S 2 ±S 3orientation with respect to principal compression. The stressanalysis of both the monoclinic and conjugate kink-bandfolds shows that the principal compression s 1 during theD4 deformation was subvertical.8. Discussion8.1. Kinematic interpretation of the D1±D2 structuralsuccessionÐbuilding of the nappe pileThe D1 fabric in eclogites has developed during a pressureand temperature peak in spatially different setting thanthe present-day country rocks. After their emplacement inthe crust, the eclogites were passively transported at the baseof the Lower Crystalline Nappe during the D2 thrusting andbehaved as rigid inclusions in a weak matrix of micaschists.This kind of behaviour resulted in the non-uniformorientationof the D1 structures within the isolated eclogiticboudins which were developed mainly during the largescaleD3 folding KlaÂpova et al., 1998).The D2 structures in the parautochthonous plagioclaseschists, and in the allochthonous orthogneisses, are associatedwith the main metamorphic event. They representsyn-metamorphic structures developed during thewestward thrusting of the Lower Crystalline Nappe overthe metasedimentary Saxothuringian parautochthon, asinferred from numerous kinematic indicators. KlaÂpova etal. 1998) suggested that the D2 structures in the eclogitesdeveloped in the same kinematic regime.During the ®nal stages of D2, the principal compressionthat was acting E±W caused large-scale buckling of thewhole nappe sequence and of the parautochthonous domain.These buckle folds have vertical axial planes and foldhinges perpendicular to the L 2 stretching lineation. Weinterpret these structures as a result of the buckling of thestacked nappe sequence at the end of a D2 thrusting episodecaused by the buttressing effect of a rigid autochthon. TheMeÏdeÏnec synformand the Oberwiesenthal structureoriginated in the same manner.8.2. F3 refolding and regional interference patternTwo large-scale F 3 antiforms later refolded the wholenappe system, together with the parautochthonoussequence. These antiforms have E±W striking steep axialplanes and westward plunging hinge zones. In the KlõÂnovecantiform, the D3 deformation is associated with the rotationof the omphacite L 1 lineation in the eclogites as a result ofthe active rotation of eclogitic boudins within weakerorthogneisses and metasediments.The increase in the intensity of the D3 folding in the southof the study area may be the result of the original geometryof the D2 synforms. If the late F 2 folds have a large wavelengthand small amplitude as, for example, the MeÏdeÏnecsynform), later compression parallel to their axes will resultin the development of dome and basin structures interferencepattern type I of Ramsay 1962b)). The observedstructural pattern indicates, however, that late F 2 folds areoverprinted by F 3 folds with moderate to steeply plunginghinges. The development of this kind of F 3 fold suggests thatlate F 2 folds were non-cylindrical domes and basinselongated in a N±S direction with increasing amplitude tothe south. Such periclinal basins and domes are commonlydescribed fromthe Zagros Mountains Iran) or the JuraMountains Price and Cosgrove, 1990, pp. 262±263). TheN±S compression of such a structure may produce theobserved large folds with steeply plunging hinges andE±W vertical axial plane. In this manner, an interferencepattern of type II originates Ramsay, 1962b). The D3 foldingintensity is thus controlled by D2 fold shape with basementdepth increasing towards the south and interlimb angleincreasing towards the north.8.3. Strain pattern in the F 3 fold limbs and F 3 fold mechanicsThe strain ellipsoid in the limb zones of both the KlõÂnovecand MeÏdeÏnec antiforms is characterised by oblate shapessuggesting ¯attening perpendicular to the F 3 fold axialplane Fig. 8). Moreover, the X-axis of ®nite strain inthese zones is sub-horizontal and E±W in direction as aresult of the D3 deformation. We suppose that in E±W63

1388J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392Fig. 10. Calculated evolution of the D3 fabric as a result of homogeneous isovolumic pure-shear deformation applied to the D2 strain ellipsoid with N±Strending Z-axis and vertical X-axis. Black and grey lines show the evolution of the strain ellipsoid with 5% step of strain black dots) being labeled by percentof strain accommodated by the X-axis of the D2 strain ellipsoid negative values represent shortening). a) Results of calculation starting fromthe D2 ellipsoidwith the plane strain symmetry and low strain intensity grey ®eld labeled D2). Provided that almost the entire D3 deformation is accommodated by the E±Wtrending Y-axis of the D2 ellipsoid, after 40±50% of shortening by pure-shear, the resulting fabric will show E±W trending long axis, oblate symmetry, andstill relatively low intensity of deformation grey ®eld labeled D3). b) Results of calculation starting from the D2 ellipsoid with prolate symmetry and mediumintensity of deformation grey ®eld labelled D2). Provided that almost the entire D3 deformation is accommodated by the E±W trending Y-axis of the D2ellipsoid, after 40±60% of shortening by pure-shear, the resulting fabric will show weak vertical elongation, oblate symmetry and high strain intensity grey®eld labelled D3).oriented limbs, the ductile shortening was superimposed onthe original D2 plane strain fabric.We have simulated superposition of homogeneouscoaxial deformation on a plane-strain ellipsoid with steepX-axis and vertical E±W trending XY plane D2 fabric afterF 3 passive foldingÐFig. 7) and the results are shown inFig. 10. Fromthe above presented geological arguments,it is assumed that the D2 and D3 Z-axes were parallelafter rotation of S 2 foliation during large scale F 3 folding.Thus, considered isovolumic changes of the shape offabricellipsoidareinducedbyshorteningoftheformerZ-axis. This isovolumetric strain is accommodated bydifferential elongation of former Y- and X-axes of theD2 ellipsoid. Paths in Fig. 10 represent differentamounts of elongation of the former X-axis negativevalues represent shortening). Points on the strain paths64

J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392 1389aNS 2L 2L 2bD 3 kink-bandsNL 1L 2L 3D 4 kink-bandsFig. 11. Schematic sketches showing the evolution of large-scale structures in time. a) Situation after the late D2 deformation is interpreted as a result ofbuttressing of the allochthonous unit from the west. Continuous E±W compression leads to the development of large-scale periclinal structures the MeÏdeÏnecsynform and Oberwiesenthal structure). b) Subsequent N±S compression leads to refolding of the late D2 structures by km-scale antiforms the MeÏdeÏnec andKlõÂnovec antiforms). Brittle-ductile structures in the ¯at lying S 2 foliation appear during the late D3 deformation. Disappearance of the N±S oriented D3compressive stress leads to strengthening of the role of overburden. Vertical D4 compression produces D4 kink-band folds to kink-bands.show 5% increments of strain imposed on a former D2fabric ellipsoid.In those areas that were not affected by the D3 deformation,we have identi®ed three types of the D2 strains symmetriesFig. 8) and these approximately represent startingpoints of trajectories in the Fig. 10. An attempt was made to®nd those ®nite strains observed in natural samples from themegafold limbs. Moreover, different ®nite D3 fabrics mustbe produced after an equal number of increments of the D3deformation. Our analysis demonstrates Fig. 10) that thiscase is reached when most of the imposed shortening isaccommodated by elongation of former Y-axis and smallchange of former X-axis. For plane-strain to constrictionalvertical D2 fabric with higher strain intensity Fig. 10a),65

1390J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392>>1σ 1 (N-S)σ 1 (vertical)temperature (oC)400 500 6002a1132b011depth (km)30 50Fig. 12. Diagramof the orientation of s 1 versus depth and temperature during D3 and D4 shows two possible scenarios of such an evolution. Path 1±2a±3shows N±S compression during exhumation of the studied area, which is demonstrated by progressive development of brittle-ductile D3 structures. After thedecreasing role of the N±S compression at shallow crustal levels and low temperature, the prevailing vertical compression will produce the D4 kink-bands.This scenario is consistent with ®eld observations. The path 1±2b shows the diminishing role of the N±S compressive stress and increase in the importance ofvertical compression caused by the overburden at high temperatures in depth. This evolution would result in refolding of the F 3 folds by ductile F 4 folds withsubhorizontal axial planes. The PT path adopted fromKonopaÂsek 2001).superimposed D3 deformation will produce strong oblatefabrics without macroscopically visible aggregate lineation.In the case of low D2 strain intensities Fig. 10b), the resultingfabric is marked by L±S symmetry with intermediateintensities and visible horizontal L 3 aggregate stretchinglineation former Y-axis).The ®eld observations suggest that the F 3 folds developedby mechanisms of buckling of a single layer of orthogneisssurrounded by micaschist. Strong ¯attening in the limbareas of the KlõÂnovec antiformalso indicates conditions ofmoderate viscosity contrast between strong orthogneissesand weak micaschists.8.4. The origin of the D3±D4 kink bands and other D4brittle-ductile structuresA question arises: were the above-described D4 kinkbandsdeveloped as post-D3 structures, or were they createdduring the early D3 deformation and then re-oriented duringthe D3 folding? As described above, the steep metamorphicfabric in metasediments between the northern limb of theKlõÂnovec antiformand the MeÏdeÏnec antiform, as well as thezone of metasediments in the southern limb of the KlõÂnovecantiform, are believed to represent the S 3 cleavage. Therefore,the kink-bands affecting the S 3 cleavage must be post-D3. This is the case of those kink-folds developed in thesteep southern limb of the MeÏdeÏnec antiformand in thesteep zones south of the KlõÂnovec antiform. On the otherhand, the kink-bands developed in the hinge zone of theMeÏdeÏnec antiformactually correspond to the axial cleavageof this structure and are, therefore, developed in the earlystages of D3 folding.If we accept a hypothesis that the kink-bands in metasedimentsclose to the KlõÂnovec and the MeÏdeÏnec antiformswere not generated in the same stress ®eld as the D3 largescale folds, another local source of stress must have existedto generate these structures. A possible source of stressduring the D4 is the weight of the overburden after therelaxation of the D3 stresses. This will produce subverticals 1 , which has been documented in the southern part of theKlõÂnovec antiform. Moreover, the expected deformationassociated with the relaxation of the D3 stresses is ratherlow. This is consistent with the development of kink bandsand brittle-ductile clevage in orthogneisses and metasediments.The amount of strain achieved during theirdevelopment is probably more than two orders of magnitudeless than the deformation associated with the D1±D3 deformationand, thus, rather insigni®cant at the scale of thestudied area.8.5. Variations of principal compression direction throughtimeThe analysis of the D2, D3 and D4 structures shows66

J. KonopaÂsek et al. / Journal of Structural Geology 23 2001) 1373±1392 1391complex stress±temperature±time evolution Fig. 11). Asdiscussed above, the D2 structures are associated with westwardthrusting of the Lower Crystalline nappe over theSaxothuringian basement suggesting non-coaxial deformationwith E±W oriented s 1 direction.Strain analysis within D3 fold limbs indicates a N±Soriented horizontal compression s 1 and important verticalstress s 2 inhibiting elongation in the vertical direction.Then, the only elongation accommodating N±S horizontalcompression s 1 is possible in the horizontal E±W direction.This indicates an important role of rigid overburden forstresses acting in a low viscosity layer at depth. SubhorizontalN±S compression leads to lateral ¯ow of materialwithout the vertical component in the early stages ofD3 deformation. The same stress regime, however, isresponsible for the formation of kink band structures inthe hinge of the MeÏdeÏnec antiform. Kink band structuresrepresent a brittle-ductile regime Dewey, 1965, 1969)and, thus, indicate a decrease in temperature under thesame orientation of the D3 stress ®eld. It is, therefore,suggested that the horizontal N±S compression operatedduring an uplift of the whole region fromthe deep to thesupracrustal level.The presence of D4 kink bands must be associated withthe disappearance of horizontal stresses and increase insubvertical s 1 compression. If this change of stress regimewould occur at high temperature conditions, then the F 3folds would be refolded by large-scale ductile F 4 foldswith subhorizontal axial planes. Fig. 12 documents thatthe F 3 folding started at peak temperature conditions ca.6008C) enabling viscous buckling of the nappe sequencecaused by shortening in the N±S direction. This deformationphase continued during a temperature decreaseassociated with uplift, resulting in the development of D 3kink-bands and a crenulation cleavage in the hinge zones oflarge-scale anticlines. The latest F 4 phase occurred underrelatively low temperature conditions and indicates thedisappearance of horizontal stress, allowing vertical shorteningof steep D 3 fabric. These changes in stress can beinterpreted as an increase of the role of overburden aftertermination of the D3 compressive stress.8.6. Exhumation of eclogites and the importance ofextension in the Czech part of the KrusÏne hory Erzgebirge)MountainsIn contrast with the western Saxothuringian domain,where the boundary between eclogites-bearing nappes andsupracrustal autochthon can be easily identi®ed, a similarlimit is dif®cult to establish in the eastern KrusÏne horyMountains. Using structural and petrological criteria, wehave de®ned the boundary between the parautochthonousSaxothuringian metasediments and allochthonous eclogitesbearingnappe. Thrusting-associated structures developed inboth the parautocthonous and allochthonous units documentthat the nappe emplacement occurred in the middle crust at adepth corresponding to 13±15 kbar. Field observations,however, are not able to provide any information aboutthe mechanism of emplacement of eclogites from a depthcorresponding to 26 kbar to the base of non-eclogiticorthogneiss nappe. This work shows mechanical behaviourof the crust during and after the nappe emplacement witheclogites as a part of lithological assemblage. The exhumationof assembled parautochthonous and allochthonous unitsto supracrustal levels is associated with complex structuralreworking of originally simple fabric during the subsequentN±S shortening, which is responsible for the ®nal pattern ofthe central part of the KrusÏne hory Mountains in the CzechRepublic.There is a range of publications emphasising the role ofthe late Variscan extensional deformation for the ®naltectonic and metamorphic pattern of the German part ofthe eastern Saxothuringian domain Willner et al., 1994;Krohe, 1996, 1998; RoÈtzler et al., 1998). These interpretationsare based on regional distribution of metamorphicunits and orientation of shear structures in different areasof the Erzgebirge. In this study, we have examined a particularlysuitable area with steeply developed anisotropyaffected by vertical shortening and demonstrated that itachieves no more than ®rst percents of the bulk strain duringthis event. If the extensional tectonics would be a keyregime responsible for ®nal geometry of studied crystallinecomplexes then signi®cantly more important vertical shorteningshould be expected ®rst in areas with subverticallydeveloped anisotropy. Therefore, we suggest that the structuralpattern of the whole Saxothuringian domain should bere-evaluated in terms of detailed structural analysis todemonstrate the real signi®cance of late orogenic extensionin this part of the Bohemian Massif.AcknowledgementsWe are very grateful to John Cosgrove for helpfulcomments on an early version of the manuscript. Thepaper has also bene®ted from the comments of G. Oliverand W. Franke. We also thank Jaroslav Synek for drawingFig. 6. This work was funded by the Grant Agency of theCzech Republic, grant no. 205/96/0279.ReferencesCobbold, P.R., Cosgrove, J.W., Summers, J.M., 1971. Development ofinternal structures in deformed anisotropic rocks. Tectonophysics 121), 23±53.Dewey, J.F., 1965. 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JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 108, NO. B1, 2023, doi:10.1029/2001JB000632, 2003Strain distribution and fabric development modeledin active and ancient transpressive zonesKarel Schulmann, 1 Alan Bruce Thompson, 2 Ondrej Lexa, 1 and Josef Ježek 3Received 23 May 2001; revised 26 April 2002; accepted 20 May 2002; published 16 January 2003.[1] A model based on kinematics of transpression allows the measurable internalparameters (simultaneous strain, fabric, and vertical elevation) to be simulated in relationto the macroscopically determined external parameters of transpressive zones (i.e., zonewidth, velocity, angle of convergence, and zone depth). We present a diagram ofconvergence angle against time for homogeneous transpression, where isolines of strainintensity D, strain symmetry K, and vertical elevation of rock samples are superposed.However, the distribution of internal strain parameters is sensitive to three types of strainpartitioning: (1) Discrete partitioning results in general decrease of finite strainaccumulations and in increase of pure shear component, (2) ductile partitioning splits thetranspressional domain into a pure shear zone where strain accumulations decreases and ina wrench-dominated zone where strain accumulations increases, and (3) viscositypartitioning is marked by different strain rates in zones of different viscosity and therefore bydifferent strain parameters. INDEX TERMS: 8025 Structural Geology: Mesoscopic fabrics; 8110Tectonophysics: Continental tectonics—general (0905); 0905 Exploration Geophysics: Continental structures(8109, 8110); KEYWORDS: transpression, strain partitioning, exhumation, oblique convergenceCitation: Schulmann, K., A. B. Thompson, O. Lexa, and J. Ježek, Strain distribution and fabric development modeled in active andancient transpressive zones, J. Geophys. Res., 108(B1), 2023, doi:10.1029/2001JB000632, 2003.1. Introduction[2] Modern tectonic studies face the problem of understandingthe relationships between small-scale structuresand large-scale geometry in orogenic zones. The linksbetween the external and internal parameters governingthe problem are particularly important. These factors maybe viewed as far-field causes related to local effects. Weexamine here these relationships in ancient and activetranspressive zones. External structural parameters are representedby the geometry of orogenic zones (i.e., zone width(distance between colliding plates), obliquity (angle ofconvergence) and depth of the transpressive zone. Irregularitiesof plate boundaries (shape of indenting block), andthe velocity and duration of plate convergence) may alsoplay an important role. Internal (local) structural parametersthat can be examined are fabric and strain intensity, symmetryof fabrics, orientations of strain axes, pressure (depth)memory, and metamorphic facies of the rocks. Theseinternal structural parameters are determined by the interactionsof the external forces with local lithological heterogeneitiesand rheologies in the transpressive zone.[3] Relative motion of lithospheric plates on a sphericalsurface is such that the plate convergence vectors are often1 Institute of Petrology and Structural Geology, Faculty of Science,Charles University, Prague, Czech Republic.2 Department Erdwissenschaften, ETH Zurich, Switzerland.3 Institute of Applied Mathematics and Computer Science, Faculty ofScience, Charles University, Prague, Czech Republic.Copyright 2003 by the American Geophysical Union.0148-0227/03/2001JB000632$09.00not orthogonal to plate boundaries [McKenzie and Parker,1967; Dewey, 1975]. These plate boundaries experiencecombined transcurrent and convergent displacements associatedwith development of deformation zones of differentsize. Within continental blocks, the deformation is not onlyrestricted to active plate boundaries but occurs within zonesof weakness inside rigid continental domains [e.g., Tommasiand Vauchez, 1997] and can be approximately described asa deformation of a weak zone bounded by rigid blocks withsteep parallel walls. All the mentioned types of deformationzones can be more or less described by a model called‘‘transpression,’’ introduced first by Harland [1971], developedby Sanderson and Marchini [1984], and elaborated bymany others. Despite its simplicity, it seems that the modelstill has much to contribute toward our understanding of thenature of convergent orogeny. In this work we develop themodel to quantify the effects of external (macroscopic)parameters on temporal development of internal strainparameters in transpressive (obliquely convergent) weakzones of finite width.2. Structural Definition of Transpression andProblems to Be Solved[4] The classical zone of transpressive deformation is atabular weak region subjected between its steep walls to asimultaneous pure shear and simple shear. In the model ofSanderson and Marchini [1984] the material is able to slipfreely upward (vertically extruded) along the walls of thetranspression zone. The transpressive deformation zonedefined by Sanderson and Marchini [1984] was also limiteddownward by a rigid horizontal plate (like a rigid floor),ETG 6 - 169

ETG 6 - 2 SCHULMANN ET AL.: STRAIN DISTRIBUTIONthus allowing extrusion only in the vertical direction.Fossen and Tikoff [1998] suggested boundary conditionswhere zones of transpressive deformation are able to growthor shrink vertically as well as horizontally along strike.[5] Fossen and Tikoff [1993] defined two types of transpressionon the basis of the orientation of the instantaneousstretching axes: (1) pure-shear-dominated transpression inwhich v 1 and v 3 (eigenvectors of instantaneous straincorresponding to maximum and minimum elongation)define a vertical plane and the v 1 (lineation) is vertical,and (2) wrench-dominated transpression in which v 1 and v 3are horizontal, while the v 2 (intermediate eigenvector) isvertical. Pure-shear-dominated transpression operates whenthe angle between relative plate convergence direction andplate boundary is greater than 20° [Pinet and Cobbold,1992], and the main axis of finite strain is always vertical. Inwrench-dominated transpression acting at an angle of convergence(a) of

SCHULMANN ET AL.: STRAIN DISTRIBUTION ETG 6 - 3Figure 1. Block diagrams and principal strain axes orientations for transpressive systems with externalparameters: (a) plate velocity (v = 1 to 10 cm yr 1 ) and initial width (d = 50 to 300 km), which give theratio R vd from 2 to 0.03, respectively, and rigid floor depth (RFD = 70, 40 km). As internal parameters,the instantaneous orientation of foliation plane (XY ) and lineation (X ) are related to angle of convergence(a) of(c)a =60° and (d) a

ETG 6 - 4 SCHULMANN ET AL.: STRAIN DISTRIBUTIONFigure 2. (a) Diagram showing relation between strain intensity (D) and time (expressed by timeparameter, k t , where t = k t /R vd ). The dot-dashed curves show time evolution of D for different angle ofconvergence (a). The distribution curve at the left shows the envelope of natural strains summarized byPfiffner and Ramsay [1982], Hrouda [1993], and others. The horizontal line through the maximum (at D =1.2) shows the strain observable by field measurements. (b) The relationship between strain intensity (D)and angle of convergence (a) through three time sections (5, 10, and 15 Myr). The contours show variationsof R vd values, which distinguish narrow and fast converging transpressive orogens (R vd > 0.5) from wideand slowly converging ones (R vd < 0.5). Vertical lines at convergence angles 30° and 50° show strainaccumulations for R vd values of active transpressive zones in New Zealand and Sumatra, respectively.suggests that the base of lithospheric transpressional zonesis not always present as a rigid floor. In such a case, verticalmovements within the system are controlled by isostaticalresponse and can thus lead also to downward motion ofmaterial early in the shortening history.4. Modeling of Internal Parameters ofTranspressional Systems[16] We calculate the internal strain parameters of transpressionalsystems (strain rate, finite strain intensity, andsymmetry and orientation of finite strain axes) in terms ofexternal parameters defined above. The description ofmethod of calculations and derivation of equations are givenin Appendices A–D.4.1. Strain Rates and Finite Strain Intensities[17] The first result of our model is that in the range ofassumed R vd , a strain rate interval from 10 14 s 1 to 10 16s 1 is obtained. This is within the range of generallyassumed strain rates extrapolated from experimental laboratorydata [Carter and Tsenn, 1987] and corresponds wellto that deduced for natural orogens by Pfiffner and Ramsay[1982]. Theoretical curves of finite strain accumulation fordifferent obliquities are presented in Figure 2b. In Figure 2bthe strain intensity parameter D (see Appendix A) is plottedon the vertical axis against the time parameter, k t , whichrelates the time of deformation with R vd , the ratio ofconvergence velocity and initial zone widthk t ¼ tR vd :The introduction of such a time parameter follows from thedefinition of the transpression model and allows us to graphthe temporal developments corresponding to zones ofdifferent width and convergence velocities (Figure 2a). Ifthe ratio R vd = 1, the scale of the k t axis represents timedirectly in million years. For other R vd values we obtain thecorresponding time from equation (1).[18] In order to compare Figure 2a with those on Figure 7of Pfiffner and Ramsay [1982] each of the curves calculatedfor convergence angles varying from a =0° to 90° is labeledwith its strain rate value. Because of the triaxial character ofthe deformation in transpressive zones we use a strain ratecalculated as rate of change of the square root of theminimum eigenvalue corresponding to short axis of instantaneousstrain tensor, and instead of R = X/Z, which is a goodcharacteristic for plane strain we use the D parameter.ð1Þ72

SCHULMANN ET AL.: STRAIN DISTRIBUTION ETG 6 - 5[19] The strain rates are expressed in terms of R vd . Thefinite strain development is more rapid for pure frontalconvergence than for a transpressional zone of any obliquity.The convergence with a =90° and a =0° end-membercurves represent pure shear and simple shear strain ratepaths in Figure 7 of Pfiffner and Ramsay [1982]. Thedensity distribution curve on the left side of Figure 2aindicates the distribution of finite strains plotted fromPfiffner and Ramsay [1982] and Hrouda [1993] and completedwith finite strain data of Rajlich et al. [1988],Schultz-Ela and Hudleston [1991], Schulmann et al.[1994], and Kirkwood [1995].[20] The strain intensity D is strongly dependent on boththe angle of convergence a, and on the ratio R vd .Itistherefore highest for frontal collision at high R vd . Withdecreasing R vd and decreasing convergence angle, the strainintensity decreases rapidly (Figure 2b). Starting from zonesmarked by the value R vd = 0.1 the variations in strainintensity are negligible with time. However, zones withR vd values higher than 0.2 show rapid increase of strainintensity after only a short duration of convergence. Thehighest gradient of strain intensity increase occurs for lowconvergence angles (wrench-dominated transpression). Forhigher convergence angles, the strain intensity increasessmoothly with increasing convergence angle. The strainintensity value also increases steadily with time, which isconsistent with progressive accumulation of finite strainduring shortening of a collisional belt.4.2. Temporal Evolution of Foliation and Lineation[21] In our model, in pure-shear-dominated transpression(a >20°) the mineral growth lineation is vertical for anywidth of the belt and any angle of convergence. In wrenchdominatedtranspression (a < 20°), the orientation oflineation varies for different angle of convergence a andfor different belt width d (Figure 3a). The orientation of thelongest strain axis, s 1 , starts at a maximum angle of 45° (fora simple shear zone) with respect to zone margin and tendstoward parallelism with simple shear direction (a = 0)represented by zone boundaries with increasing amount ofdeformation. However, the rate of lineation reorientation isvery rapid in the case of high R vd . For instance for R vd =2and convergence at an angle of 10°, the lineation rotatesfrom an initial angle of 40°–5° in 3 Myr. For R vd = 0.03 itwill take 200 Myr. An important conclusion is that in widebelts and slow highly oblique convergence (e.g., R vd = 0.03)the stretching lineation should be oriented for long times ata high angle (20°–35°) with respect to plate boundariesafter a long period of convergence (20–50 Myr, Figure 3a).In contrast, narrow oblique belts of rapid convergence (e.g.,R vd = 2) should have stretching lineations which are subparallel(

ETG 6 - 6 SCHULMANN ET AL.: STRAIN DISTRIBUTIONFigure 3. (a) Diagram showing orientation of mineral lineation (x axis) through time (parameter k t )forwrench-dominated transpression (a

SCHULMANN ET AL.: STRAIN DISTRIBUTION ETG 6 - 7Figure 4. (a) Diagram showing the development of strain symmetry (K ) in time (parameter k t )forvarious convergence angles (a). (b) Diagram showing the relationship between strain symmetry (K ) andangle of convergence (a) through three time sections (5, 10, and 15 Myr). The curves show only slightdependence of K on the R vd values and on time.[31] However, the RFD is not easily defined in ancienttranspressional zones, and we can make only rough estimatesbased on petrological data and observed depth ofMoho. Such analysis shows that in the case of shear zonesof southern Madagascar the depth in which the rocks wereoriginally located was around 60–70 km [Martelat et al.,1997; Pili et al., 1997]. Martelat et al. and Pili et al. alsosuggested that the lithospheric shear zones are coupled withunderlying mantle so that the RFD may be located evendeeper in the mantle lithosphere [Teyssier and Tikoff, 1997;Vauchez et al., 1998].6. Superposition of Strain Parameters inTranspressive Belts[32] The preceding considerations and results enable us tocreate a type of map of strain parameters registered in rockselevated to the surface. Figure 6 shows such a map whereisolines of strain intensities D and isolines of strain symmetryK are superposed on a diagram of convergence angleagainst time parameter (a - k t space). We have addedcontours of RFD/z 0 to this (dashed curves in Figure 6).These shows the times when the samples are elevated closeto the surface. The ratio RFD/z 0 also expresses the acrosswidthshortening of the zone. The shaded area in a -] k tspace shows the range of naturally observed strain intensities.For these D values (strain intensities) the strainsymmetry parameter K corresponds well with high or lowa (angle of convergence), respectively, for vertical andhorizontal orientation of lineation. On the other hand theintensity of strain (D) is not very sensitive to the angle ofconvergence greater than 20° (pure-shear-dominated transpression)but is strongly dependent on time. For the case ofvery obliquely convergent zones (low a) the time needed toaccumulate observed strain intensities are almost twice aslong as for high convergence angles. The maximum strainsat the right side of the shaded area in Figure 6 may beattributed to rock samples elevated from thickened midcrustaldepths (40 km) of transpressional zones. However,this is only valid for zones with convergence angles >50°.For more oblique zones, only very shallow samples (uppermost25% of the zone) can be exhumed. This also meansthat horizontal stretching lineations may be exhumed fromvery shallow depths in soft transpressional zones.7. Effects of Strain Partitioning on TemporalStrain Parameter Development[33] Next we examine the effects of three different typesof strain partitioning on the temporal development of finitestrain parameters.[34] Discrete displacement partitioning is modeled usingan approach of Teyssier et al. [1995]. In their model,variable amounts of total lateral displacement can be consideredto be consumed by discrete faulting. This isexpressed by a ratio p 1 between fault-accommodated lateraldisplacement and total lateral displacement. The evolutionof strain parameters in the viscous domain of homogeneous75

ETG 6 - 8 SCHULMANN ET AL.: STRAIN DISTRIBUTIONFigure 5. Vertical elevation rate for transpression expressed in terms of angle of convergence (a) andtime parameter (k t ) for different base depths, RFD = 40, 70 km, and different original sample depths z 0 =30, 60 km. The curves show elevation achieved by these samples in a given time. For example, in Figure5a (RFD = 40 km, z 0 = 30 km) for convergence angle a =80° sample is elevated to depth 10 km after k t =1.2. For the case of R vd = 0.1 the time of elevation corresponds to 12 Myr.deformation, where the rest of lateral displacement and thewhole across strike shortening are accommodated, can bededuced using our hypothetical strain map (Figure 6). Toevaluate the influence of partitioning on strain parameters,recalculation of values R vd and a are made using theequations (B1) and (B2) (in Appendix B) or they can betaken from Figure 7. The effect of discrete partitioning onfinite strain parameters is expressed by a virtual increase ofconvergence angle and decrease of R vd (decrease of plateconvergence velocity or increase of weak zone width). Thusfor a transpression zone with an angle of convergence of45°, R vd equal 0.1, and with 50% of lateral displacementconsumed by discrete faults, we use values of a 0 = 63.43°0and R vd = 0.079 (Figure 7). For 10 Myr of convergencewithout discrete partitioning, finite strain parameters are D =1.64 and K = 0.58. When discrete partitioning accommodates50% of the lateral displacement, the finite strainparameters are D = 0.9 and K = 0.8.[35] Ductile partitioning splits the deformed domain intoa pure shear zone (PSZ) and a wrench-dominated zone(WDZ). We assume that the pure shear-across-strike shorteningand elevation are homogeneously distributed acrossthe whole system, while simple shear-lateral displacement isaccommodated only in the WDZ. We examine developmentof strain parameters for different widths of the WDZ,expressed as ratio p 2 of the width of WDZ and the widthof the whole transpressional zone. Such a partitioning ofpure shear and simple shear within transpressional zones isresponsible for decomposition of the velocity gradienttensor into two separate tensors according to equations(C1) and (C2) (in Appendix C).[36] While the evolution of strain parameters in the PSZmay be obvious, the evolution of strain parameters in theWDZ zones is less so. The results of calculations for WDZzones of different width are shown in Figure 8. Fifty percentof ductile partitioning (Figure 8a) exhibits a similar patternof strain parameters as a nonpartitioned system. Nevertheless,the domain with horizontal lineation and the domainof oblate symmetry are both slightly enlarged. This isrelated to the shift of the lineation switch, which occurs76

SCHULMANN ET AL.: STRAIN DISTRIBUTION ETG 6 - 9Figure 6. Superposition of strain intensity (D), strain symmetry (K ), and sample elevation for variousvalues of RFD and initial sample depth z 0 , plotted in terms of angle of convergence (a) and time parameter(k t ). The values of D show a strong dependence on R vd and time (but not much on a)fora =20° to 90°;fora

ETG 6 - 10 SCHULMANN ET AL.: STRAIN DISTRIBUTIONFigure 7. Relation between amount of discrete partitioning (amount of lateral displacement0accommodated by fault) and change of angle of convergence a and R vd ratio in transpression zone.Knowing the angle of convergence and the amount of discrete partitioning, new angle of convergence a 00and new R vd values can be depicted and used for estimation of strain parameters in diagram Figure 6.Thick black lines indicate amount of discrete partitioning for different active transpressive zones.mulate strain during the entire exhumation path. A rock’scapacity to accumulate strain is determined by its thermal,microstructural and rheological development. The bestexamples are syntectonically emplaced granites in transpressionalshear zones [e.g., Melka et al., 1992; Parry et al.,1997; Brown and Solar, 1998]. There is only a very shorttime of granite intrusion when the magma has enoughcrystals to record the deformation. Further cooling of themagma is responsible for hardening of granites and freezingof the fabrics at certain depth levels.[39] Numerous regional studies of transpressional zonesshow that rock complexes have experienced ductile deformationfollowed by late folding and brittle fracturing in thesame tectonic regime. This means that the accumulation ofductile strain ends at some depth below the surface dependingupon the ambient thermal regime. It is also common thatsamples originally located deep in lithospheric transpressionalzones have had their mineralogy changed due toretrograde metamorphism during exhumation. This meansthat deformation registered in transpressional zones wasdeveloping only during a small time interval during theconvergent activity.[40] To be able to correlate the external and internalparameters using the transpressional model, we need tospecify the time span (or the duration of the verticalelevation path) in which the accumulated strain is attributed.This means that in Figure 6 the time axis together withelevation contours (ratios of R vd /z 0 ) needs to be rescaled sothat times corresponding to the finite strain parameters, Dand K, are higher.9. Discussion[41] Transpressional models in general assume that aweak and deformable zone bounded by the rigid walls ofadjacent lithosphere is progressively shortened in the courseof convergence. We can put forward a question as towhether the measured internal (microscopic) parameters ofancient transpressional zones (lineation, foliation, K and Dvalues) may be used to estimate the initial external (macroscopic)parameters (R vd , RFD, and a)?9.1. Ancient Zones[42] The strain symmetry parameter K is sensitive to theangle of convergence (a), whereas the strain intensityparameter D is well correlated with across-width shorteningof a transpressional zone. On the basis of Figure 6 and usingmeasured natural finite strains and orientation of lineation in78

SCHULMANN ET AL.: STRAIN DISTRIBUTION ETG 6 - 11Figure 8. Superposition of strain intensity (D) and strain symmetry (K ) in terms of angle ofconvergence (a) for different widths of wrench-dominated zone: (a) 50%, (b) 20%, (c) 10%, and (d) 5%.For 100%, see Figure 6. These diagrams may correspond to different depth levels through thetranspressional zone with increasing width of wrench-dominated zone with depth in agreement with themodel of Pinet and Cobbold [1992].the rock (vertical for a >20°) we can conclude that in theframework of our model the maximum initial width ofancient transpressional zones should not exceed doubletheir actual width. Transpressive zones with horizontallineation corresponding to highly oblique convergence(low a) would have been 1.4 times wider originally. Thereforethe maximum initial width of ancient transpressionalzones could vary from 20 to 100 km. This width isapparently less than those of recently active transpressionalsystems.[43] Provided the average range of plate velocities ofcontinental convergence did not change through geologicalhistory, from upper Proterozoic to Recent, we can estimatethe range of lifetimes of ancient transpressional zones. Asexamples of highly obliquely convergent transpressionalzones with large amounts of finite strain data, we canconsider the Central Bohemian Shear Zone (horizontallineation, K =0.8 1, vertical foliation, D =0.5 1.5[Rajlich et al., 1988]), and the transpressional zone developedin the eastern Variscan Culm Basin (horizontal lineation,K = 0.6–1.2, vertical foliation, D = 0.4 1.5[Rajlich, 1990]). These transpressional zones are associatedwith homogeneous ductile deformation over a width of upto 15 km. Because these transpressional zones are associatedwith low- to very low grade metamorphism, weconsider also an example of finite strain analysis from ahighly oblique transpressional zone developed at high-grademetamorphic conditions (kyanite-sillimanite zone) in theEastern Bohemian Massif which show horizontal lineationwith strain symmetry ranging from to oblate to plane strainand D from 0.6 to 1.2 [Schulmann, 1990]. As an example oftranspression with a supposed high angle of convergenceand with a large amount of data on finite strain, we considera transpressional zone in the Archean greenstone belt ofMinnesota (vertical lineation and steep foliation, K variesaround plane strain, D ranges from 1.5 to 2.5 [Schultz-Elaand Hudleston, 1991]), where the deformed zone reaches anapproximate width of 30 km. Assuming plate velocities in79

ETG 6 - 12 SCHULMANN ET AL.: STRAIN DISTRIBUTIONthe range 1–5 cm yr 1 , the lifetimes of these shear zonesmight be in the range 0.5–4 Myr. These times are too shortfor an assumed duration of orogenic events. If we considerthat the duration of orogenic events is long (over an intervalof several tens of million years), then the computed strainintensities for such given external parameters are unrealisticallyhigh (up to D = 100).9.2. Active Zones[44] Our calculations show that for a transpressional zonewith R vd < 0.2, the strain rate does not exceed a value of 6 10 15 s 1 for any angle of convergence (a). This R vd valuecorresponds to the active zones of Sumatra (d = 300 km, v =70 mm yr 1 , R vd = 0.23) and San Andreas (d = 200 km, v =50 mm yr 1 , R vd = 0.25). For a zone with value R vd = 0.48,corresponding to the Alpine Fault zone in New Zealand,there will be maximum strain rate of 1.58 10 14 s 1distributed over a width of 100 km for a plate velocity of 48mm yr 1 . Zones with R vd >2 (small d, large v) show highstrain rate values of 3.2 10 14 s 1 and will apply fornarrow continental zones (10–20 km) with a mean platevelocity of 5 km yr 1 . Note that all of these values areconsistent with the strain rate estimates suggested by Carterand Tsenn [1987].[45] Assuming that the above mentioned strain valuesdevelop in large zones of homogeneous deformation, inFigure 6b we have plotted recent macroscopic convergenceparameters of well-known active transpressive zones. Wecan take the above listed macroscopic parameters of thesetranspressive zones and ask at what time are the averagestrain intensities developed? Figure 6b shows that theaverage strain of D = 1.2 would be produced in the Sumatrazone after 3.75 Myr, in San Andreas after 4.8 Myr, and inAlpine Fault Zone, New Zealand, after 1.7 Myr.[46] The average strains, characteristic for most of thezones with dispersed deformation are achieved in 5 Myr forzones with R vd = 0.2 and after 10 Myr for zones with R vd =0.1. We note that the New Zealand zone of distributedfaulting would have a strain intensity corresponding to D =13 after 5 Myr, while the San Andreas and Sumatra zoneswould remain within a realistic range of strain intensities (Dof x to y). We note that for the entire lifetime of the SanAndreas system (20 Myr) only the last 5 Myr are attributedto dextral transpression [Walcott, 1993] which is caused byPacific plate rotation [Luyendyk et al., 1985]. Similarfeatures are reported from paleomagnetic investigations ofthe New Zealand system [Walcott, 1987], and therefore anydirect application of the transpressional model to strainaccumulations in time is problematic.9.3. What Effects Could Explain These Discrepancies?[47] We are unable to correlate succinctly the external andinternal parameters of homogeneous transpression. Thisinconsistency, apart from plate rotation, is well explainedby the three concepts of strain partitioning discussed above.Discrete partitioning results in general decrease of finitestrain accumulations and in increase of pure shear componentin deformed zone. The effects of ductile partitioningbecomes important for narrow wrench-dominated zone (lessthan 30% of the width of the whole transpressional zone)and is responsible for decrease of strain accumulation in thepure-shear-dominated zone and increase in the wrenchdominatedzone. We note that the pure-shear-dominatedzone exhibits all structural characteristics of frontal shortening.Viscosity partitioning is marked by different strainrates in domains of different viscosity leading to differentstrain accumulations. In strongly oblique zones (a

SCHULMANN ET AL.: STRAIN DISTRIBUTION ETG 6 - 13logical data allowing estimate of depth changes throughtime exist but so far are related to intrusions of magmas intranspressional regimes [e.g., Melka et al., 1992; Parry etal., 1997]. Future work should certainly concentrate onacquisition of fabric data together with detailed petrologicalinvestigations, to delimit the depth-temperature regime, andwith geochronological investigations, to understand theduration of transpressional orogeny.9.4. Importance of the Switch of Lineation[52] As an additional result of our modeling, we haveobtained information about the switch of lineation thatdeserves to be mentioned in conclusion. Analysis of thelineation produced by the classical transpression modelshows that the switch of lineation during progressive shorteningof the zone is theoretically possible. However, thecorresponding strain intensities are very high and the degreeof oblateness does not permit meaningful measurement of thechange in the linear fabric. An interesting observation is thatoblate fabrics are measured only exceptionally, more oftenwe have to deal with lineation (vertical or horizontal andclose to plain strain symmetries). In contrast to the model, innature horizontal and vertical lineations are often observedsimultaneously in transpressional zones [Hudleston et al.,1988; Schultz-Ela and Hudleston, 1991; Melka et al., 1992;Tikoff and Greene, 1997]. This fact underlines the ideas ofpartitioning of the strain in transpressional zones [Tikoff andGreene, 1997] or perhaps to later superposition of simpleshear zones on already existing fabrics [Jones and Tanner,1995].Appendix A: Strain Parameters[53] In the transpressional model, computation of thethree-dimensional motion path of rock samples and strainparameters in the transpressive belt are based on thevelocity gradient tensor:0 10 _g 0L ¼ @ 0 _e 0 A: ðA1Þ0 0 _eThe pure shear strain rate component _e (corresponding tovertical extrusion and horizontal shortening) and simpleshear rate component _g (corresponding to lateral horizontaldisplacement) are calculated using the respective equations:_e ¼ n d sin a_g ¼ n cos a;d ðA3ÞðA2Þwhere v is the relative plate velocity, d is zone width, and ais the angle of convergence (Figure 1). In our model we willconsider an example of constant strain rate during temporaldevelopment of a transpressional zone, so for given periodfinite strain parameters can be obtained from followingfinite strain tensor:0_g1_ecos a ½ 1 exp ð _et ÞŠ01F ¼ @ 0 exp ð _et Þ 0 A: ðA4Þ0 0 exp ð_etÞDuring deformation we trace rock samples moving in thetranspressive zone (Figure 1b). At each position of a rocksample we calculate finite strain geometry, i.e., the principaldirections and the principal stretches by taking theeigenvectors and square roots of eigenvalues, respectively,of ‘‘Cauchy-Green tensor’’ FF T [Truesdell and Toupin,1960], which are further used to calculate the parameters Kand D [Ramsay and Huber, 1983, pp. 201–202].[54] Strain intensity is expressed by the D value:qffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiD ¼ R xy 1 2þ Ryz 1 2: ðA5ÞThe strain symmetry K value is calculated asK ¼ R xy 1R yz 1 ; ðA6Þwhere R xy = S 1 /S 2 and R yz = S 2 /S 3 represent elongationaspect ratios measured on orthogonal axes. Furthermore,when K = 0 a (S) planar fabric is indicated, and for K =1(LS) a plane strain fabric in which foliation is equally strongas lineation, and K = 1 for a linear fabric without foliation.[55] Vertical displacement produced by the pure shearcomponent can be expressed by a valuezt ðÞ¼z 0 exp ð_etÞ;ðA7Þwhere z 0 = z(t 0 ) is the initial vertical distance of a rocksample above a reference level (rigid floor depth) of zeroelevation.Appendix B: Discrete Partitioning[56] To evaluate strain parameters in a domain wherediscrete partitioning operate, we can use above definedapproach using following recalculated values of angle ofconvergence a and R vdtan a 0 ¼tan að1 p 1 ÞðB1ÞqffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiR 0 vd ¼ R vd ð1 p 1 Þ 2 cos 2 a þ sin 2 a; ðB2Þwhere p 1 is ratio between fault-accommodated lateraldisplacement and total lateral displacement.Appendix C: Ductile Partitioning[57] Ductile partitioning in transpression zone results indecomposition of velocity gradient tensor into two componentsfor pure shear zone (L PSD ) and wrench-dominatedzone (L WDZ ) as follows:010 0 0L PSD ¼ @ 0 _e ð1 p 2 Þ 0 A0 0 _e ð1 p 2 ÞðC1Þ010 _g 0L WDZ ¼ @ 0 _ep 2 0 A; ðC2Þ0 0 _ep 281

ETG 6 - 14 SCHULMANN ET AL.: STRAIN DISTRIBUTIONwhere p 2 is ratio of the width of WDZ and the width of thewhole transpressional zone.Appendix D: Viscosity Partitioning[58] Viscosity partitioning assumes different strain ratesor different R vd , respectively, in two adjacent domains whilethe angle of convergence remains constant. The values ofR vd areR m 2for low-viscosity domain andr mvd ¼ R vd1 þ p 3 r m 1 ðD1Þvd ¼ R 1vd1 þ p 3 r m 1 ðD2ÞR m 1for high-viscosity domain, whereis viscosity contrast.r m ¼ m 1; ðm m 1 > m 2 Þ2[59] Acknowledgments. The Czech National Foundation grant 205/98/K004, the Grant Agency of Charles University grant 296/1997/B GEO,the Schweizerische National Fonds and ETH Research credits are gratefullyacknowledged for financial support. We thank Basil Tikoff and an anonymousreviewer for very helpful reviews.ReferencesBrown, M., and G. S. Solar, Shear-zone systems and melts: Feedbackrelations and self- organization in orogenic belts, J. Struct. 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Thompson, Modelling of strain partitioningin transpression zones, in Proceedings of the fourth annual conferenceof the International Association for Mathematical Geology, editedby A. Buccianti, G. Nardi, and R. Potenza, pp. 827 – 832, De Frede,Naples, Italy, 1998.Jones, R. R., and P. W. G. Tanner, Strain partitioning in transpression zones,J. Struct. Geol., 17, 793–802, 1995.Jones, R. R., R. E. Holdsworth, and W. Bailey, Lateral extrusion in transpressionzones: The importance of boundary conditions, J. Struct. Geol.,19, 1201–1217, 1997.Kirkwood, D., Strain partitioning and progressive deformation history in atranspressive belt, northern Appalachians, Tectonophysics, 241, 15–34,1995.Lisle, R., and J. Savage, Factors influencing Rock Competence - Data Froma Swedish Deformed Conglomerate, Geol. Foren. Stockholm Forh., 104,219–224, 1982.Luyendyk, B. P., M. J. Kamerling, R. R. Terres, and J. S. Hornafius, Simpleshear of southern California during Neogene time suggested by paleomagneticdeclinations, J. Geophys. Res., 90, 2454–2466, 1985.Martelat, J. E., C. Nicollet, J. M. Lardeaux, G. Vidal, and R. Rakotondrazafy,Lithospheric tectonic structures developed under high-grade metamorphismin the southern part of Madagascar, Geodin. Acta, 10, 94–114, 1997.Martelat, J. E., J. M. Lardeaux, C. Nicollet, and R. Rakotondrazafy, Strainpattern and late Precambrian deformation history in southern Madagascar,Precambrian Res., 102, 1–20, 2000.McKenzie, D., and R. L. Parker, The North Pacific: An example of tectonicson a sphere, Nature, 216, 1276–1280, 1967.Melka, R., K. Schulmann, B. Schulmannova, F. Hrouda, and M. Lobkowicz,The evolution of perpendicular linear fabrics in synkinematicallyemplaced tourmaline granite (central Moravia-Bohemian Massif ),J. Struct. Geol., 14, 605–620, 1992.Newman, J., W. M. Lamb, M. R. Drury, and R. L. M. Vissers, Deformationprocesses in a peridotite shear zone: Reaction-softening by an H 2 O-deficient,continuous net transfer reaction, Tectonophysics, 303, 193–222,1999.Nyman, M. W., R. D. Law, and S. S. Morgan, Conditions of contactmetamorphism,Papoose Flat Pluton, eastern California, USA-Implicationsfor cooling and strain histories, J. Metamorph. Geol., 13, 627–643, 1995.Parry, M., P. Stipska, K. Schulmann, F. Hrouda, J. Jezek, and A. Kroener,Tonalite sill emplacement at an oblique plate boundary; northeasternmargin of the Bohemian Massif, Tectonophysics, 280, 61–81, 1997.Pfiffner, O. A., and J. G. Ramsay, Constraints on geological strain rates:Arguments from finite strain states of naturally deformed rocks, J. Geophys.Res., 87, 311 – 321, 1982.Pili, E., Y. Ricard, J. M. Lardeaux, and S. M. F. Sheppard, Lithospheric shearzones and mantle-crust connections, Tectonophysics, 280, 15–29, 1997.Pinet, N., and P. R. Cobbold, Experimental insights into the partitioning ofmotion within zones of oblique subduction, Tectonophysics, 206, 371–388, 1992.Rajlich, P., Strain and tectonic styles related to Variscan transpression andtranstension in the Moravo-Silesian Culmian basin Bohemian Massif,Czechoslovakia, Tectonophysics, 174, 351–367, 1990.Rajlich, P., K. Schulmann, and J. Synek, Strain analysis on conglomeratesfrom the central Bohemian shear zone, Krystalinikum, 19, 119–134,1988.Ramsay, J. G., and M. I. Huber, The Techniques of Modern StructuralGeology, vol. 1, Strain Analysis, Academic, San Diego, Calif., 1983.Ranalli, G., and D. C. Murphy, Rheological stratification of the lithosphere,Tectonophysics, 132, 281–295, 1987.Robin, P. Y., and A. R. Cruden, Strain and vorticity patterns in ideallyductile transpression zones, J. Struct. Geol., 16, 447–466, 1994.Royden, L., Coupling and decoupling of crust and mantle in convergentorogens: Implications for strain partitioning in the crust, J. Geophys. Res.,101, 17,679–17,705, 1996.Sanderson, D. J., and W. R. D. Marchini, Transpression, J. Struct. Geol, 6,449–458, 1984.Schulmann, K., Fabric and kinematic study of the bites orthogneiss (southwesternMoravia)—Result of large-scale northeastward shearing parallelto the Moldanubian Moravian boundary, Tectonophysics, 177, 229–244,1990.Schulmann, K., R. Melka, M. Lobkowicz, P. Ledru, J. M. Lardeaux, andA. Autran, Contrasting styles of deformation during progressive nappestacking at the southeastern margin of the Bohemian Massif (ThayaDome), J. Struct. Geol., 16, 355–370, 1994.Schultz-Ela, D. D., and P. J. Hudleston, Strain in an Archean greenstonebeltof Minnesota, Tectonophysics, 190, 233–268, 1991.Stuwe, K., and T. D. Barr, On uplift and exhumation during convergence,Tectonics, 17, 80–88, 1998.Teyssier, C., and B. Tikoff, Study of strike-slip partitioning in Californiaresolves San Andreas discrepancy and constrains lithospheric structure,Geol. Soc. Am. Abstr. Programs, 29, 347, 1997.Teyssier, C., B. Tikoff, and M. Markley, Oblique plate motion and continentaltectonics, Geology, 23, 447–450, 1995.Thompson, A. B., K. Schulmann, J. Jezek, and V. Tolar, Thermally softenedcontinental extensional zones: Arcs and rifts as precursors to thickenedorogenic zones, Tectonophysics, 332, 115–141, 2001.82

SCHULMANN ET AL.: STRAIN DISTRIBUTION ETG 6 - 15Tikoff, B., and D. Greene, Stretching lineations in transpressional shearzones: An example from the Sierra Nevada Batholith California, J. Struct.Geol., 19, 29–39, 1997.Tikoff, B., and C. Teyssier, Strain modeling of displacement-field partitioningin transpression orogens, J. Struct. Geol., 16, 1575–1588, 1994.Tommasi, A., and A. Vauchez, Continental-scale rheological heterogeneitiesand complex intraplate tectono-metamorphic patterns: Insights froma case study and numerical models, Tectonophysics, 279, 327 – 350,1997.Truesdell, C. A., and R. A. Toupin, The classic field theory, in Encyclopediaof Physics, vol.III,Principles of Classical Mechanics and FieldTheory, edited by S. Flügge, pp. 226–793, Springer-Verlag, New York,1960.Vauchez, A., and A. Nicolas, Mountain building-Strike-parallel motion andmantle anisotropy, Tectonophysics, 185, 183–201, 1991.Vauchez, A., A. Tommasi, and G. Barruol, Rheological heterogeneity, mechanicalanisotropy and deformation of the continental lithosphere, Tectonophysics,296, 61–86, 1998.Walcott, D., Neogene tectonics and kinematics of western North America,Tectonics, 12, 326–333, 1993.Walcott, R. I., Geodetic strain and the deformational history of the NorthIsland of New Zealand during the Late Cainozoic, Philos. Trans. R. Soc.London, Ser. A, 321, 163–181, 1987.White, J. C., and C. K. Mawer, Deep-crustal deformation textures alongmegathrusts from Newfoundland and Ontario—Implications for microstructuralpreservation, strain rates, and strength of the lithosphere, Can.J. Earth Sci., 29, 328–337, 1992.Wood, D. S., Current views of the development of slaty cleavage, Annu.Rev. Earth Planet. Sci., 2, 369–401, 1974.J. Ježek, Institute of Applied Mathematics and Computer Science,Faculty of Science, Charles University, Albertov 6, 128 43 Prague, CzechRepublic. ( jezek@natur.cuni.cz)O. Lexa and K. Schulmann, Institute of Petrology and Structural Geology,Faculty of Science, Charles University, Albertov 6, 128 43 Prague, CzechRepublic. (lexa@natur.cuni.cz; schulman@natur.cuni.cz)A. B. Thompson, Institut für Mineralogie und Petrographie, ETHZentrum, NO E 64, CH-8092 Zürich, Switzerland. (thompson@erdw.ethz.ch)83

TECTONICS, VOL. 22, NO. 6, 1066, doi:10.1029/2002TC001472, 2003Cretaceous collision and indentation in the West Carpathians:View based on structural analysis and numerical modelingOndrej Lexa and Karel SchulmannInstitute of Petrology and Structural Geology, Charles University, Prague, Czech RepublicJosef JežekInstitute of Applied Mathematics and Computer Science, Charles University, Prague, Czech RepublicReceived 29 October 2002; revised 13 March 2003; accepted 30 July 2003; published 21 November 2003.[1] A model of indentation of a rigid promontory intoweak metasedimentary rocks during Cretaceousconvergence is suggested for tectonic evolution ofsouthern part of West Carpathians. An early arcuatecleavage fan has developed in front of northwardmoving southern rigid basement block. The interactionof the moving indenter with western stationarybasement promontory is responsible for developmentof the boundary-parallel shear zone along which themain southern indenter is shifted to the east. This resultsin development of a new steep transpressional cleavageoverprinting the early fabric. Eastward displacement ofthe southern indenter causes the development of a thrustzone parallel to the margin of the eastern stationarypromontory. A proposed numerical model of thedeformation of a thin viscous sheet in front of ovalrigid indenter reliably simulates the development of theobserved deformation pattern. Modeled discretepartitioning between the western promontory and theindenting block fully agrees with the observedsecondary cleavage associated with the transpressionalshear zone. Our numerical model interconnects thiscomplex kinematic frame with finite strain pattern,which was to date possible only for simple boundaryconditions. In addition, the model explains thepolyphase cleavage patterns in terms of complexshapes of promontories and changes in movements ofindenting blocks. INDEX TERMS: 5475 Planetology: SolidSurface Planets: Tectonics (8149); 8005 Structural Geology: Foldsand folding; 8020 Structural Geology: Mechanics; 8110Tectonophysics: Continental tectonics—general (0905);KEYWORDS: Cretaceous collision, indentation, west Carpathians,numerical model. Citation: Lexa, O., K. Schulmann, and J. Ježek,Cretaceous collision and indentation in the West Carpathians: Viewbased on structural analysis and numerical modeling, Tectonics,22(6), 1066, doi:10.1029/2002TC001472, 2003.1. Introduction[2] Existing interpretations of Mesozoic tectonic evolutionof Alpine and Carpathian chains are based on broadCopyright 2003 by the American Geophysical Union.0278-7407/03/2002TC001472$12.005 - 1knowledge of metamorphic and structural data obtainedfrom studies of internal parts of these orogenic belts[Genser et al., 1996; Schmid et al., 1996; Trumpy, 1973].These data underline subduction-related tectonics associatedwith lithospheric-scale subduction of oceanic and Europeanlithosphere below the African indenter [Allemand andLardeaux, 1997]. The Mesozoic evolution of the WestCarpathian belt is an integral part of the closing of theTethyan ocean, being characterized by Jurassic subductionand Late Jurassic exhumation of high pressure rocks[Faryad, 1995; Faryad and Henjes-Kunst, 1997; Maluskiet al., 1993; Mock and Reichwalder, 1992].[3] The West Carpathian structure is traditionally interpretedas a result of Cretaceous crustal shortening ofVariscan basement crustal segments associated withdécollement of Mesozoic sedimentary sequences. TheMesozoic strata originally formed basins separating individualcrustal segments, which were inverted and passivelytransported toward the north in the form of twolarge-scale nappes (the lower Krížna and upper Chočnappes) [Andrusov, 1936, 1958; Biely et al., 1968;Plašienka, 1991; Plašienka et al., 1997]. The Cretaceoustectonic evolution in West Carpathians was mostly studiedin Mesozoic rocks with well-known stratigraphy[Biely and Bystrický, 1967; Bystrický, 1967; Vozár,1978]. Consequently, the Cretaceous stacking of Mesozoicsequences, the kinematics and displacement ofMesozoic nappes and the degree of their metamorphicreequilibration are well defined and understood [cf.Plašienka, 1995]. However, the Alpine structural andmetamorphic evolution of pre-Mesozoic crystalline rocksis poorly described due to difficulties in distinguishingVariscan and Alpine fabrics.[4] This paper presents new information on structural andmechanical behavior of the southern part of West Carpathianpre-Mesozoic crystalline segments during Cretaceousconvergence. We have investigated the polyphase cleavagepatterns developed in low-grade metasedimentary EarlyPaleozoic sequences. It is generally accepted that suchcleavage patterns reflect the geometry and directionof movement of adjacent rigid blocks [Sintubin, 1999;Woodcock et al., 1988]. The detailed knowledge of shapesof rigid Variscan promontories and the geometry andsuperposition of cleavage patterns allow modeling of thecollisional process in time. We use the finite elementnumerical method introduced by England et al. [1985] to85

5 - 2 LEXA ET AL.: COLLISION IN WEST CARPATHIANSmodel the finite strain pattern in time and space for complexboundary conditions.2. Geological Setting[5] The West Carpathians have been traditionally dividedinto the outer and inner structural zones. However, takinginto account the recent studies of Mesozoic evolution, thetriple division into Inner, Central and Outer West Carpathiansis commonly accepted (Figure 1) [see Plašienka etal., 1997]. The Outer West Carpathians (OWC) are representedby flysch dominated Paleogene formations andMesozoic rocks lacking pre-Mesozoic basement. The Innerand Central West Carpathians (IWC and CWC), separatedby a Late Jurassic oceanic suture, are composed of a pre-Late Cretaceous imbricated nappe system comprisingcrystalline basement units with characteristic Late Paleozoicand Mesozoic sequences [Matějka and Andrusov, 1931] andpostnappe Late Cretaceous to Neogene sedimentary andvolcanic formations.2.1. Pre-Cretaceous Geology of the Studied Area[6] The studied area is located in the southern part of theCWC and comprises three major lithological and tectonometamorphicunits (Figure 1). From the north to the southand from the bottom to the top they are (1) Variscancrystalline basement (Vepor Unit) with Late Paleozoic andMesozoic cover sequences, (2) Early to Late Paleozoic,basinal, mostly low grade turbiditic sequences (GemerUnit), and (3) Mesozoic accretionary wedge containingblueschist facies relics overlain by (4) flat, nonmetamorphosedSilica nappe.2.1.1. Variscan Crystalline Basement: The Vepor Unit[7] The crustal rocks of the Vepor basement are composedof two contrasting Variscan metamorphic domainsexhibiting pre-Alpine thrust tectonics. The structurallylower domain, generally dipping to the north, is composedof medium-grade schists exhibiting peak PT conditions inKy-St micaschists corresponding to maximum of 10 kbarand 550–600°C [Korikovskij et al., 1989; Méres andHovorka, 1991]. The classic Barrow type metamorphiczonation is difficult to establish due to the Alpine greenschistfacies overprint but in the eastern part of the Veporbasement (Čierna Hora Mountains), a metamorphic zonationranging from the biotite zone in the south to thestaurolite zone in the north is documented [Jacko et al.,1990; Korikovskij et al., 1990]. The structurally highercrystalline unit is represented by a domain of heterogeneouspara- and ortho-derived migmatites intruded byporphyritic to medium-grained peraluminous granites.The Variscan age of metamorphism is documented byTh-U-Pb dating of monazite (370–350 Ma [Janák et al.,2001a]) and by sporadically preserved 40 Ar/ 39 Ar coolingages ranging from 358 to 312 Ma [Dallmeyer et al.,1996]. High-grade fabrics represented by compositionallayering and stromatitic banding in migmatites are dippingeither to the NW or to the N under steep to mediumangles. Peak temperature conditions yield 680–730°C atpressures of 4–6 kbar (Figure 2) [Siman et al., 1996].These PT conditions most likely correspond to lateexhumation and decompression melting. This is deducedfrom other parts of West Carpathians where relics ofhigh-pressure assemblages preserved in similar types ofmigmatites were found [Hovorka and Méres, 1989;Janák et al., 1996].[8] The southern Vepor Variscan crystalline basement isunconformably covered by Late Carboniferous (Stephanian)sandstones and shales [Planderová and Vozárová, 1978].These metasediments were intruded by granitoids that wereresponsible for contact metamorphism ranging from biotiteto cordierite zones. Contact metamorphic conditions of500°C and 2 kbar were established by Vozárová [1990].Vozárová also suggests that contact metamorphism overprintedregional greenschist facies assemblages of Carboniferousmetasediments. The Vepor crystalline complexes aswell as the Late Carboniferous sequences are overlain byPermian and Triassic clastics and locally of Middle and LateTriassic carbonates.2.1.2. Early to Late Paleozoic Gemer Unit[9] The Gemer Unit consists of three principal groups thatdiffer in lithology and metamorphic grade. The amphibolitefacies metamorphic conditions are reported from the KlátovGroup [Faryad, 1990], which tectonically overlies thegreenschist facies metabasites and phyllites of the RakovecGroup. The Variscan metamorphic conditions correspond to440–480°C at 6–8 kbar for the northern part and 350–430°C at 4–5 kbar (Figure 2) for the southern part of theRakovec Group [Faryad and Bernhardt, 1996; Vozárová,1993]. The metamorphic foliation is represented by compositionallayering in metabasites dipping to the NNW undershallow to intermediate angles. The tectonically lowermostGelnica Group builds the major part of the Gemer Unit. Thelower part of this unit consists of a thick turbiditic sequenceof Ordovician age [Soták et al., 1999; Vozárová etal., 1999]gradually passing into volcanosedimentary sequencesaccompanied by products of massive rhyolite-dacite volcanismof probably Silurian-Devonian age [Cambel et al.,1990]. The top of the Gelnica Group is composed ofphyllites and black shales containing sporadic carbonatelenses. The metamorphic conditions indicate temperaturesof 350–400°C and pressures of 2.5–3.5 kbar [Faryad, 1992,1994] for regional metamorphism, which is overprintedby contact metamorphism (450–550°C and 1.5–2 kbar[Faryad, 1992]) associated with intrusion of granites incentral part of the Gelnica Group. Strong greenschist faciesFigure 1. (opposite) Geological and structural map of the studied area with trajectories of cleavage based on this work andthat of Snopko and Reichwalder [1970]. The inset in the upper left corner shows the Carpathian arc with location of thestudied area.86

LEXA ET AL.: COLLISION IN WEST CARPATHIANS 5 - 387

5 - 4 LEXA ET AL.: COLLISION IN WEST CARPATHIANSFigure 2. Idealized E-W cross section showing pre-Cretaceous structural and metamorphic pattern ofthe studied area. Insets show Variscan (Gelnica, Rakovec, and Klátov Formations) and Jurassic (Bôrkanappe) metamorphic PT conditions. The eastern ‘‘unknown’’ block represents hypothetical active marginduring closure of Meliata ocean. The legend is same as for Figure 1.metamorphic foliation and relics of sedimentary bedding aredipping conformably to the NNW. The metamorphic gradeand degree of structural transposition are gradually vanishingto the south.[10] Already consolidated Rakovec and Gelnica Groupsare discordantly covered by Early Carboniferous flysch andPermian clastics of ‘‘red bed’’ type. Sedimentation wasterminated by the development of Late Permian-EarlyTriassic evaporite formations. The metamorphism of Permianclastics yielded temperatures of 200–250°C [Šucha andEberl, 1992].2.1.3. Mesozoic Meliata Accretionary Wedge andSuperficial Triassic Nappe[11] Mesozoic sequences are located mainly along thesouthern margin of the Gemer Unit, locally rimming itsnorthern margin or forming klippens on the Gemer Unit.The bottom part of the Meliata accretionary wedge isformed by the sequence of thrust sheets (initially thesouthern slope of the European continent) consisting ofthinned continental margin, Permian clastics, Triassic limestonesand blueschist facies metabasalts (Figure 2) (380–460°C at 9–13 kbar [Faryad, 1995]).[12] These rocks are overlain by subduction-relatedmelange composed of deep-water oceanic sediments of theMeliata ocean and relics of oceanic mantle rocks. Theuppermost part of the accretionary wedge is composed ofTriassic (Turňa) carbonates, shales and pelagic sediments ofnorthern slope of the Apulian continent. The high pressurerocks of the Meliata accretionary wedge exhibit strongductile and polyphase deformation of all lithologies. Themylonitic foliation is generally dipping to the east, bears anintense southeast plunging stretching lineation and kinematicindicators, such as sigmoidal pebbles of metabasalts, whichsuggest top to the northwest transport. The Turňa shalesshow very low grade metamorphic overprint and locallydeveloped slaty cleavage subparallel to the bedding.[13] The accretionary wedge and the underlying GemerUnit are tectonically overlain by extensive horizontallylying Silica nappe derived from the Apulian shelf. The baseof the nappe is composed of Late Permian-Early Triassicevaporites and shales followed by Middle and Late Triassiccarbonates, which were affected by brittle faulting [Hók etal., 1995].2.2. Summary of Pre-Cretaceous TectonometamorphicEvolution of the Studied Area[14] The metamorphic and structural patterns describedabove are consistent with north-south polarity of the Variscanorogenic evolution marked by southward thrusting ofdeep-seated lower crustal and middle crustal complexesover low-grade foreland. The Vepor Unit is regarded as arelic of the Variscan internal domain in which the anatecticlower crust is thrust over the middle crustal Barroviancomplex. The original Paleozoic relationship between Veporand Gemer Units is largely obscured by later Mesozoicevolution. North dipping pre-Mesozoic fabrics and increasein metamorphic grade from the south to the north in both theGemer and Vepor Units suggest an existence of N-S polarityof Paleozoic orogeny [Faryad, 1990]. On the basis of thesedata, the Klátov and Rakovec groups may be interpreted asrelics of an Early Paleozoic basin underthrust beneath thehigh-grade rocks of the internal domain represented by theVepor Unit exposed in the north. In our model the wholenappe stack was further thrust over Early Paleozoic basinrepresented by the Gelnica Group, which is probablyunderlain by a Neo-Proterozoic basement (Figure 2). Thelack of thrust-related structures in Middle Carboniferousmetasediments indicates that the nappe stacking occurred inpre-Westfalian times.[15] The southward thrusting of high-grade gneisses overthe low-grade mostly metasedimentary foreland formed twopromontories separated by embayment of weakly metamorphosedsediments. This irregular geometric distribution ofgneissic complexes and soft sediments played a critical rolein the subsequent Mesozoic tectonic evolution.[16] The last major pre-Cretaceous tectonic event responsiblefor the final structural pattern was the Jurassic southeastwardsubduction (in recent coordinates) of the Meliataocean and the southern passive margin of the Europeanplatform. This process resulted in formation of an accretionarywedge and its northwestward obduction over both88

LEXA ET AL.: COLLISION IN WEST CARPATHIANS 5 - 5the Gemer and Vepor Units during the Late Jurassic (150–160 Ma [Dallmeyer et al., 1996; Faryad and Henjes-Kunst,1997; Maluski et al., 1993]) (Figure 2). In the southern partof the Gemer Unit, sedimentary bedding is well preservedand folded by large-scale open folds with N-S trendinghinges (Figure 3). This folding is connected with developmentof spaced cleavage steeply dipping to the east suggestingthat the thinned continental margin was intensivelyreworked during Jurassic subduction processes.3. Cretaceous Polyphase Structural Evolution:Collisional Stage[17] The Cretaceous collisional evolution of the Gemerand Vepor Units is marked by four major distinct tectonicevents: (1) formation of the Gemer Cleavage Fan (GCF)structure affecting the central part of the Gemer Unit,(2) extensional deformation of the western Vepor promontory,(3) transpressional shearing affecting the western Veporpromontory and development of the Trans-Gemer ShearZone (TGSZ), and (4) extrusion of the Gemer Unit over theeastern Vepor promontory along the Eastern Gemer Thrust(EGT).[18] In order to evaluate strain variations in conglomerateswe used published data [Németh et al., 1997], while forslates the degree of deformation associated with developmentof crenulation cleavage was evaluated using criteriaintroduced by Bell and Rubenach [1983]. In our work weused five stages of crenulation cleavage development[Passchier and Trouw, 1996, Figure 4.17ab] to semiquantitativelycharacterize the strain gradient in studied area.3.1. Early Cretaceous Deformation:Gemer Cleavage Fan[19] The Gemer Cleavage Fan represents the most spectacularstructure overprinting the pre-Mesozoic metamorphicfabric of the Gelnica and Rakovec Groups as well asJurassic fabric to the south. This cleavage forms an asymmetricpositive fan structure developed across the entirelength of the Gemer Unit (Figure 3). The intensity andmetamorphic grade of the cleavage are highest in a 5 kmwide, E-W oriented axial zone of the fan structure. Here, theonly relics of pre-Mesozoic fabric in the form of rootlessfolds are preserved in intensively developed steep slatycleavage corresponding to stages 4 and 5 of Bell andRubenach [1983] classification (Figure 3). Finite strainmeasurements of Németh et al. [1997] show oblate finitestrain symmetry and rather moderate strain intensities. Thiszone is characterized by the presence of numerous bodies ofgranites. These intrusions have been originally consideredCretaceous in age [Kantor, 1957; Máška, 1957; Vozár et al.,1996] and later, based on monazite (273 Ma [Finger andBroska, 1999]) and zircon U-Pb (275–245 Ma [Poller et al.,2002]) dating as Permian to Early Triassic. The maincleavage of the GCF is in the axial zone affected by kinkbands with kink planes perpendicular or oblique to stronglydeveloped vertical anisotropy (Figure 5a). These structuresare interpreted as a result of vertical shortening of steepcleavage associated with the evolution of GCF.[20] The intensity and degree of metamorphism of thenorth dipping, spaced cleavage of southern flank of GCFrapidly decrease to the south. The cleavage intensity correspondsto stages 1 and 2 of Bell and Rubenach’s [1983]classification. Here, the Cretaceous deformation is so weakthat the structures associated with Jurassic obduction ofMeliata accretionary wedge are well preserved. In thenorthern flank of GCF, the cleavage is dipping to the south.The dip angle gradually decreases to the north in conjunctionwith decreasing intensity of cleavage developmentcorresponding to the stages 1 and 2 of Bell and Rubenach[1983] (Figure 5b). In northern parts of the Gelnica Group,the north vergent kink bands developed mainly withinincompetent shales while competent lithologies like quartzitesand volcanics were gently folded by open folds withwavelengths of a few hundreds of meters (Figure 3a). Thedecrease in cleavage intensity is even more pronounced inamphibolite and greenschist facies metabasites of the RakovecGroup where a spaced to fracture cleavage is developedindicating very low finite strains [Passchier and Trouw,1996] (Figures 3a and 5c).[21] The Late Carboniferous and Permian cover of theGemer Unit, directly overlying the Rakovec Group, is alsoaffected by GCF. In this domain, the deformation is stronglyheterogeneous leading to the development of small-scale(up to several hundreds of meters) positive cleavage fanstructures. Carboniferous conglomerates contain pebbles offoliated metabasites of the Rakovec Group. However, thesecondary cleavage known from the Rakovec Group hasnever been discovered in the pebbles. This means that thesame post-Permian cleavage affects the Gemer Unit andLate Paleozoic sediments.[22] The features of GCF described above are developedprominently along the central part of the Gemer Unit(Figure 4a). Lateral extension of GCF toward western andeastern Vepor promontories (Figures 3a and 4b) is markedby change in cleavage trend, so that it becomes parallel totheir boundaries. In addition, the positive fan structuredisappears and the cleavage is dominantly vertical and veryintense. Triassic and Jurassic sediments of the Meliataaccretionary wedge in the front of the western Veporpromontory exhibit a strong cleavage development andreworking of Jurassic fabrics in continuation of GCF.Structural succession as well as 40 Ar/ 39 Ar cooling agesranging from 106 to 82 Ma from Gemer Unit [Dallmeyeret al., 1996] are our major arguments for Cretaceousshortening of the Gemer Unit associated with the developmentof GCF.3.2. The First Cretaceous Deformation:Vepor Extensional Mylonitization[23] In contrast to compressional deformation of theGemer Unit, the earliest Cretaceous structures developedin the Vepor basement are associated with extensionaltectonics (Figure 4a). This is manifested by mylonitizationof basement rocks and Permian-Triassic cover, characterizedby the development of anastomosed network ofgreenschist facies shear zones. The intensity of mylonitizationis highest in the cover sequences and gradually89

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LEXA ET AL.: COLLISION IN WEST CARPATHIANS 5 - 7decreases toward the deeper part of the Vepor basement,where the Variscan high-grade fabrics are preserved [Simanet al., 1996]. The Cretaceous age of this deformation isconfirmed by several Ar/Ar cooling ages of micas rangingfrom 90 to 85 Ma [Dallmeyer et al., 1996; Janák et al.,2001b; Kováčiketal., 1996; Maluski et al., 1993].[24] The style of deformation changes toward the NEedge of the Vepor basement, where narrow normal shearzone marked by strongly noncoaxial deformation associatedwith top-to-the-NE normal shearing is developed. Thisdeformation mainly affects the Late Paleozoic and TriassicVepor cover. Overlying Gemer slates and relics of theMeliata accretionary wedge contain a cleavage interpretedas manifestation of the GCF, but they are not affected byextensional reworking (Figure 4a).3.3. Cretaceous Transpressional Deformation:Trans-Gemer Shear Zone (TGSZ)[25] The transpressional deformation is marked by developmentof a several kilometer wide zone of NE trendingsteep cleavage along the southern boundary of the westernVepor promontory (Figure 4b). Here, in the Gemer Unit, theCarboniferous and Permian cover of the Vepor basementand Mesozoic rocks of the Meliata accretionary wedge, allpreviously developed structures, are intensively reworkedunder lower greenschist facies conditions (Figure 5d). In thestrongly attenuated Gemer Unit, relics of E-W trendingGCF fabric and the new NE trending cleavage form amap-scale sigmoidal domain surrounded by highly shearedLate Paleozoic rocks (Figure 4b). Locally, the early developedfoliation is refolded by synschistose noncylindricalfolds with steeply to subhorizontally plunging hinges,which become subparallel to horizontal stretching lineationswith increasing finite strain. These features are consistentwith progressive folding in transpressional shear zones[Fossen and Tikoff, 1998; Treagus and Treagus, 1992].[26] The boundary of the Vepor basement and LatePaleozoic cover is intruded by sheets of peraluminousgranites, which show magmatic fabric parallel to the transpressionalshear zone. In some places, a heterogeneous,steep NE trending mylonitic S-C fabric indicates sinistralshearing. Granite apophyses and dikes are folded andsheared or they crosscut the foliation of host rocks withoutinternal deformation. All these features are consistent withsyntectonic character of granite emplacement. Shallow levelof magma emplacement is documented by contact metamorphismof host Late Paleozoic schists (500°C and 2 kbar[Vozárová, 1990]).[27] Toward the NE, this 5 km wide zone of steep cleavagecontinues into central part of the Gemer Unit (Figures 3a and4b). This NE trending zone of shear deformation, the Trans-Gemer Shear Zone (TGSZ), overprints all previouslydeveloped metamorphic fabrics (Figure 5e) exhibiting a20 to 25 km sinistral offset of lithological stripes andaxial zone of GCF. The displacement and intensity ofdeformation gradually vanishes toward the NE edge of theGemer Unit.[28] Two several kilometers wide, NE trending zones ofgreenschist facies transpressional deformation affectedinternal parts of the Western Vepor promontory (Figure 6b).Inside these shear zones, the Variscan as well as Mesozoicmylonitic extensional fabric are strongly refolded andtransposed. The character of folding as well as numeroussense-of-shear criteria underlines the sinistral sense ofmovement. These zones are preferentially developed inweak lithologies such as garnetiferous micaschists orhighly anisotropic extensional mylonites. Outside of theseshear zones, a weak E-W trending crenulation cleavageaffects the flat Cretaceous extensional mylonitic foliation.The structural mapping shows anticlockwise rotation ofthis crenulation toward parallelism with shear zones,which confirms the sinistral displacement along shearzones.3.4. Eastern Gemer Thrust (EGT) andAssociated Deformations[29] The southern part of the Gemer Unit is displacedalong sinistral TGSZ toward NE, and consequently is thrustover the eastern Vepor promontory along large-scale compressiveshear zone, the Eastern Gemer Thrust (EGT)(Figure 3b). This zone is marked by imbrication of basementand cover (both Paleozoic and Mesozoic), intenselower greenschist facies grade mylonitization of all lithologiesacross a few kilometers wide belt. An importantfeature is the incorporation of the Gemer Permian-Triassiccover as well as the Triassic-Jurassic Vepor cover intoimbricated thrust system in the form of large-scale isoclinalsynclines (Figure 3b). The foreland duplexes and myloniticFigure 3. (opposite) Idealized cross sections of central and eastern parts of the Gemer Unit. (a) Section characteristic of30 km wide and 50 km long almost E-W trending central part of the Gemer Unit. In the north, Late Carboniferous, Permian,and Triassic (Silica nappe) sequences unconformably overlie the metabasites of the Rakovec Group of the Gemer Unit.Gelnica Group builds the major part of the Gemer Unit farther to the south, where it is unconformably overlain by Permianmetasandstones. The southernmost part of the section is formed by most completely preserved sequences of Meliataaccretionary wedge and Silica nappe at the top. (b) Eastern part showing complete section through the Meliata accretionarywedge and the Silica nappe in the south and the Gemer Unit and the eastern Vepor promontory with its Mesozoic cover tothe NE. All tectonic units are trending NW-SE. The Vepor basement rocks are imbricated with their Mesozoic cover. In thehanging wall, there occurs Late Paleozoic cover of the Gemer Unit represented by Early Carboniferous slates and Permianmetaclastics and Triassic limestones. This sequence is tectonically overlain by the Rakovec and the Gelnica Groups whichare locally covered by klippens of Silica nappe. To the southwest, the Gelnica Group covered by Permian clastics and byMeliata accretionary wedge is located. The legend is same as for Figure 1.91

5 - 8 LEXA ET AL.: COLLISION IN WEST CARPATHIANSFigure 4. Idealized cross sections across the southern and eastern margin of the Vepor promontory.(a) Southwestern section of the studied area characteristic of 40 km wide and 70 km long outcrop of thewestern Vepor promontory. Southeastern margin of this crystalline complex is rimmed by low-grade LateCarboniferous metasedimentary formation (Slatviná Formation) intruded by peraluminous granites ofuncertain age. Farther to the southeast, a belt of Permian greenschist facies metasediments occursfollowed by belt of Early Carboniferous greenschist facies slates (Ochtiná Formation) of variable width.The Gemer Unit consists exclusively of extremely reduced Gelnica Group. The southernmost part of thesection consists mostly of carbonates of Meliata accretionary wedge and Silica nappes. (b) NW-SEtrending contact of the western Vepor promontory with the Gemer Unit. The NE edge of the Veporbasement is rimmed by the Late Carboniferous formation (Slatviná Formation) and by Permian andTriassic greenschist facies grade metasandstones of the Vepor cover sequences. Along this border theLate Carboniferous formation is pinching out to the north so that the Permian rocks of the coversequences directly overlie the Vepor basement. Farther to the east, the Permian metasediments areoverlain by slates of the Gelnica Group. Large klippen of Permian rocks and metacarbonates of the Bôrkanappe overlie the Gelnica Formation. The legend is same as for Figure 1.Figure 5. (opposite) Field photographs of main structural features. (a) Steep cleavage (160/70) of the GCF affected by latesubhorizontal kink bands (350/25) in axial part of GCF (Štós Saddle). (b) Weakly developed south dipping axial planecleavage (204/45) affecting early fabric (90/60) in the northern part of the Gelnica Group (road cut north of GemerskáPoloma). (c) Weak spaced cleavage (150/30) reworking Variscan metamorphic fabrics (335/60) of the Rakovec Group(Nálepkovo). (d) Steep cleavage (150/80) and vertical noncylindrical folds of TGSZ affecting early composite S 0–1 fabricof Triassic limestones (Blh valley near Ploské village). (e) Late steep cleavage of TGSZ (338/85) reworking the earlycleavage of the Gemer Cleavage Fan (178/70) in the Gelnica Group (Smolnícka Huta).92

LEXA ET AL.: COLLISION IN WEST CARPATHIANS 5 - 993

5 - 10 LEXA ET AL.: COLLISION IN WEST CARPATHIANSFigure 6. Schematic block diagrams showing tectonic evolution of the studied area. (a) Development ofthe Gemer Cleavage Fan due to indentation of sub-Gemer basement. (b) Development of the Trans-Gemer Shear Zone and Eastern Gemer Thrust Zone resulting from interaction between the Veporpromontories with sub-Gemer basement.foliation of the EGT system are dipping to the SW, bear anintense stretching lineation plunging to the SW and show atop-to-the-NE sense of shearing.4. Tectonic Model of Cretaceous Collision[30] The structural pattern and succession of deformationstructures described above allow modeling of the tectonicevolution of SW Carpathians and development of polyphasecleavage patterns. Our interpretation is based on mutualstrength relationships between individual units at the onsetof Cretaceous convergence. Cretaceous tectonics of theWest Carpathians is characterized by north vergent collisionof a southern continent with the northerly lying WestCarpathian domain (European plate). Fragments of thesouthern continental domain, now located in northern Hungary(Bükk mountains), are considered to be Neo-Proterozoicin age [Pantó et al., 1967]. Here, the absence ofVariscan overprint is manifested by continuous sedimentationfrom the Early to Late Paleozoic. We suggest that thissouthern continent behaved as a rigid indenter controllingthe deformation of all northerly foreland crustal units, and inour coordinate system was actively moving toward thenorth. The mechanical contrast between the Vepor basementpromontories and the Gemer slates results from their contrastinglithologies and pre-Cretaceous evolution. Theinterval of thermal relaxation of the Vepor quartzofeldspathiccrust between the last (Late Carboniferous) importantthermal perturbation and Cretaceous collisioncorresponds to about 180 Ma. This indicates, that thegeotherm of the Vepor crust was equilibrated at the onsetof Cretaceous orogeny [Cloetingh and Burov, 1996; Morganand Ramberg, 1987]. Therefore we suggest that the VariscanVepor basement, composed of gneisses and granites, representeda mechanically strong promontory of irregular shape.In contrast, the Gemer Unit is represented mainly by lowgradeslates composed of fine-grained hydrous mineralswith rheology controlled by diffusion type of deformationmechanisms as pressure solution and diffusive mass transfer[Knipe, 1979, 1989]. For the same geotherm, as comparedwith laterally adjacent quartzofeldspathic rocks, the strengthof Gemer slates was incomparably lower. Taking intoaccount these rheological assumptions, the Gemer Unitduring the Cretaceous event is considered to be the weakestdomain accommodating most of the viscous deformation.[31] In agreement with Woodcock et al. [1988] andSintubin [1999], the cleavage patterns in deformable weakrocks reflect the geometry and direction of movement ofrigid blocks. In order to model the development of superposedcleavage pattern described above, it is important todefine boundary conditions.4.1. Definition of Kinematic Frame[32] The asymmetry of the GCF can be interpreted as aresult of movement of rigid indenting block to the north andback thrusting of metasediments over its northern margin.The rigid basement does not crop out, but it can be traced indeep seismic lines 2T and G1 [Tomek, 1993; Vozár andŠantavý, 1999; Vozár et al., 1996]. The seismic profilingshows that the Gemer Unit is about 5 km thick sheet-likebody resting on a basement of unknown age. This majorlithological boundary is represented by series of stronghorizontal reflectors. The most spectacular structure inseismic line G1 [Vozár and Šantavý, 1999; Vozár et al.,1996] is a highly reflective south dipping zone along whichthe horizontal base of the Gemer Unit is displaced to thenorth. This zone is interpreted as the major Sub-Gemerthrust fault responsible for northward thrusting of rigidbasement over weak sediments resulting in the developmentof GCF (Figure 6).[33] Important question in our model is the displacementof Vepor basement rocks with respect to the Gemer Unit. Itis well known that the whole central and southern Carpathiandomain was actively moving to the north (in recentcoordinates) as documented by Cretaceous progressiveclosure of Mesozoic Fatric basinal domain north of theVepor basement [Plašienka, 1997]. In this kinematic frameall units are shifted to the north but only differential move-94

LEXA ET AL.: COLLISION IN WEST CARPATHIANS 5 - 11Figure 7. (a) Initial geometry and boundary conditions of the numerical model. The arrow indicates thevelocity and trajectory of the indenter northern margin, with v = 0; zero Dirichlet boundary condition forvelocity; outflow, zero Neumann boundary condition for velocity. (b) Finite strain pattern developed inweak zone after 1 Myr of shortening. Distribution of strain intensity expressed in D value and orientationpattern of XY plane of finite strain ellipsoid. The foliations trajectories are shown by lines and theorientation of lineation is expressed by the color of foliation trace. White line corresponds to horizontallineation and black line to vertical one. (c) Finite strain pattern developed in weak zone after 3 Myr ofshortening. Distribution of strain intensity expressed in D value and orientation pattern of XY plane offinite strain ellipsoid. (d) Distribution of finite strain symmetry expressed in K value.ments within the Vepor-Gemer system are responsible for itsinternal deformation. Important observation is that thedeformation intensity in central part of the Gemer Unitvanishes to the north. In this area, the Vepor basement ispresent (covered by Tertiary sediments) but no increase inCretaceous deformation intensity has been observed. Thismeans that this Vepor segment does not creating importantdeformation resulting from possible movement to the south.Therefore we suggest that the Vepor promontories did notmove actively to the south, and extreme deformation alongwestern and eastern Vepor promontories was imposed bygenerally northward flow of weak material. In conclusion,the only differentially moving rigid block is the northwardthrusted part of sub-Gemer basement. All other basementunits can be further considered kinematically fixed.[34] Our field studies showed that apart from GCF, anintense deformation was concentrated along southeasternedge of the western Vepor promontory producing TGSZ andalso along southwestern edge of the eastern Vepor promontoryresponsible for the origin of EGT (Figure 6b). Thedevelopment of TGSZ probably results from a major changein mutual translation direction of southern sub-Gemer blockand western Vepor promontory due to their oblique collisionat deeper crustal levels. Localized transpressional deformationin upper crustal levels is a typical expression of obliqueconvergence in many active regions, e.g., San Andreas faultzone [Teyssier and Tikoff, 1997], Sumatra [Tikoff andTeyssier, 1994], or Alpine fault in New Zealand [Teyssieret al., 1995].4.2. Numerical Modeling of ProgressiveDeformation of the Gemer Unit[35] The presented numerical approach enables to modelthe deformation in a weak zone surrounded by rigid blocksor free boundaries. The approach is based on the thinviscous sheet approximation being similar to that one usedby England et al. [1985] for modeling the deformation ofthe whole lithosphere. We assume a horizontal weak tabulardomain to have been subjected to flow with no tractions attop and bottom surface. We consider vertical gradients ofthe horizontal velocity to be negligible, which allows us tointegrate the equations of motion over the vertical dimensionand to work with vertical averages of stress and strainrates. When linear constitutive relation between stress andstrain rates is considered, the procedure leads to a system ofelliptic partial differential equations for two horizontalvelocity components (see Appendix A). The system issolved by the finite element method, with the Dirichletand Neumann boundary conditions applied to segments ofthe domain boundaries corresponding to the describedgeological settings (rigid indenter, free inflow or outflow95

5 - 12 LEXA ET AL.: COLLISION IN WEST CARPATHIANSFigure 8. Finite strain pattern developed in weak zone after 7 Myr of shortening. (a) Distribution ofstrain intensity expressed in D value and orientation pattern of XY planes and X axes of finite strainellipsoid. (b) Distribution of angles between instantaneous and finite XY planes expressing the degree ofpossible simple shear reactivation of existing finite anisotropy. (c) Distribution of finite strain symmetryexpressed in K value. (d) Distribution of finite topography in meters after 7 Myr of shortening in front ofthe indenter.of material). The vertical strain rate and velocity are relatedto the horizontal velocity field by the incompressibilityequation. Ježek et al. [2002] have described the thin sheetmodel as sensitive to the angle of collision and maybeproducing a zone dominated by lateral simple shear close tothe indenter and a zone of dominant pure shear farther awayfrom the indenting boundary. In addition, we show thatthese general features can strongly interfere with finitedimension of the modeled area and imposed boundaryconditions.4.2.1. Definition of Domain Geometryand Boundary Conditions[36] As the Vepor promontories were considered to havebeen kinematically fixed throughout the whole deformationalhistory, the zero velocity (Dirichlet) boundary conditionwas applied to their boundaries (Figure 7a), so that theirgeometry corresponds to the Vepor basement recent shape.The northern boundary of our model connecting eastern andwestern promontories is characterized by a zero velocitygradient (Neumann) boundary condition allowing free outflowof material to the north. This is in agreement with theabsence of a deformation gradient in this area. Similarboundary conditions are applied to the rest of examineddomain boundaries since the extension of all geologicalunits in these particular areas is unknown, being covered byTertiary sediments. The southern edge of the model representsthe actively moving rigid body of oval shape followingprescribed trajectory with constant velocity of 1 cm/yr(Figure 7a). This velocity of plate movement is deliberatelychosen to demonstrate the principal tendencies in finitestrain patterns.4.2.2. Results of Numerical Modeling[37] The main results of our modeling are presented in aseries of three time steps equal to 1, 3 and 7 Myr. Wepresent maps of strain intensities, orientations of XY planesof finite strain and orientations of X strain axes (Figures 7b,7c, and 8a). Already after 1 Myr we can observe thedevelopment of arcuate pattern of XY trajectories aroundthe rigid indenter (Figure 7b). Another feature is thedecrease in strain intensity from the south (D =0.5)tothenorth, where the strain intensity is negligible. The strainintensity increases in the western part of the model close tothe western promontory. The strain symmetry is of planestrain type in the deformed area irrespective to the strainintensity. The X axis of finite strain is vertical in the entiredomain indicating predominant pure shear deformationregime.[38] After 3 Myr we can observe several domains withcontrasting strain parameters (Figures 7c and 7d). The centraldomain in front of the indenter shows exponential decrease instrain intensity from indenter margin (D = 1) to the north(D = 0.1). The strain gradient is poorly defined toward easternpromontory and the XY planes tend to be parallel with thepromontory margin. In the west, the strain intensity is highestacross the whole width of the shortened domain and thetrajectories of XY planes are fully parallel with the westernVepor promontory margin. The strain symmetry shows dominantflattening, K close to zero, along wide zones parallel to96

LEXA ET AL.: COLLISION IN WEST CARPATHIANS 5 - 13the Vepor promontories, whereas in the central domain theincrease of K parameter up to 0.5 occurs close to the indentermargin (Figure 7d). Pure shear is dominating along the wholedeformed area except small part southeast of the indenter,where subhorizontal stretching starts to develop. In this area,we can examine angle f between XY planes of instantaneousand finite strain ellipsoid. Here, the angle f exceeds 25°, butwith progressive northward movement of the indenter, thearea of high angle f enlarges. According to Tikoff and Teyssire[1994] and Ježek et al. [2002], high angle f is critical for thedevelopment of discrete partitioning and subsequent slipalong highly inclined surfaces. In this concept, the instantaneousstretching tensor has similar significance as does astress tensor for the development of faulting and reactivationof preexisting anisotropies [Tommasi and Vauchez, 2001].Therefore the sufficiently high angle between finite stainanisotropy and instantaneous maximum stretching axis maygenerate high resolved shear stress on preexisting surfacesand reactivate these planes as strike-slip faults. In our model,the discrete partitioning starts to operate after 3 Myr leading tocontinuous change in the indenter movement direction. Weshow the pattern developed after 7 Myr, where the observedtendencies are fully developed (Figure 8b). The central part ofweak domain shows exponential strain intensity decreasefrom D = 2.5 to approximately D = 0.5 in the most remotearea in the north (Figure 8a). Our highest calculated strainintensities correspond to maximum shortening of 70%,which is in agreement with values estimated for stronglydeveloped slaty cleavage [Siddans, 1972]. Therefore wesuggest that the zones of highest modeled strain intensitiescorrespond to the domains in the field with most intenselydeveloped cleavage mapped as stages 4 and 5 of Bell andRubenach [1983]. Similarly, modeled low strain intensities ofD = 0.5 (25% shortening) may be compared to heterogeneouslydeveloped cleavage stages 1 and 2 of Bell andRubenach [1983] in the field. Modeled strain symmetriesreach values of K = 0.5, which roughly correspond tomeasured values of Németh et al. [1997]. The area betweenthe indenter and western promontory forms now a narrowchannel in which the strain intensity is significantly higherand strain symmetry more oblate. Moreover, the X axis offinite strain becomes horizontal and starts to be controlled bysimple shear deformation. At that time, following the prescribedtrajectory (Figure 7a), the indenter bulk translationvector becomes parallel to the southern margin of the westernVepor promontory (Figure 8a). This change in indentermovement strongly affects the style of deformation in theeastern part of deformable domain. In this area, an intensestrain develops up to D = 2, the foliation trajectories beingparallel to the eastern promontory margin. The strainsymmetry remains oblate with K parameter close to zero(Figure 8c).5. Discussion5.1. Validation of Numerical ModelWith Respect to Field Data[39] The numerical model of thin viscous sheet deformationgenerated by an indenter of oval shape simulates thedevelopment of deformation pattern characteristic of GCF.We note that strain intensities decrease exponentially fromthe margin of indenting block, which is in agreement withcleavage pattern observed in the field.[40] The development of discrete partitioning betweenwestern Vepor promontory and indenting block agrees wellwith the observed secondary cleavage associated withTGSZ. We suggest that the TGSZ accommodates thechange in bulk indenter translation from northward tonortheastward during the deformation. The transcurrentmovement along the western promontory is responsiblefor propagation of TGSZ into interior of the Gemer Unit.This leads to separation of GCF into northern domain withpreserved northward movement related structures, whilethe southern domain becomes frontally convergent withthe eastern promontory. This process is responsible for thedevelopment of EGT.[41] However, the presented model has serious limitations.We are not able to simulate the deformation of thoseparts of the viscous sheet, which were thrust over rigidpromontories. This particularly concerns extensional strippingof the Gemer Unit from northeastern part of thewestern Vepor promontory. The noncoaxial extensionaldeformation in this area is most likely related to the activityof TGSZ and corresponds to pulling the allochthonousGemer Unit associated with sinistral shearing along thismajor shear zone. The model is unable to demonstrate theeffects of strain localization associated with discrete partitioning.In fact, the TGSZ is passively translating southernpart of the Gemer Unit without significant internal deformation.Similarly, the development of EGT appears to be amore localized feature than is shown in our model, andleads also to passive thrusting of the Gemer Unit over theeastern Vepor promontory.[42] Despite of these limitations, the presented modelallows to predict the strain pattern in front of indenting platein an area with complex boundary conditions. Our model isintended to quantify the cleavage patterns developed due tothe movement of rigid blocks as suggested by Woodcock et al.[1988], Sintubin [1999], and others. The basis of our modelingis the assumption that the cleavage represents the XYplaneof finite strain ellipsoid [Cloos, 1947; Sorby, 1853; Wood,1974]. Our model works with deformation of originallyisotropic medium and does not take into account problemsof existing internal anisotropy [Cobbold et al., 1971]. However,the major advantage of our approach is the interconnectionof complex kinematic frame with finite strain pattern,which was so far possible only for extremely simple boundarycondition models, e.g., simple shear, transpression, etc. Inaddition, the model explains the polyphase cleavagepatterns in terms of the complex shape of promontories andchanges in movements of indenting blocks. Moreover, usingthe regional mapping of cleavage patterns, we are now able todistinguish actively moving blocks from stationary rigidpromontories.[43] We are aware that infinite numbers of boundaryconditions exist, which may generate different strain distributionand superposition of structures. Therefore we deliberatelyselected the set of boundary conditions, which satisfy97

5 - 14 LEXA ET AL.: COLLISION IN WEST CARPATHIANSthe structural evolution in the weak domain represented by theGemer Unit. Such a type of modeling could be used tovalidate chosen boundary conditions, i.e., the role of rigidpromontories for complex structural evolutions in terrainswith polyphase deformation.5.2. Timescales of the Proposed Model[44] The timescale of the model is controlled by velocity ofthe indenting block. We have chosen the arbitrary velocity of1 cm/yr for the sake of simplicity. However, in the caseof Vepor and Gemer Units, we can define the velocity ofmovement of our kinematically fixed system. On the basis ofthe knowledge of approximate initial width and stratigraphicrecord of the Mesozoic (Fatric) basin in front of theVepor basement [Plašienka, 1997], the rate of shortening isestimated to be about 1 cm/yr, and the duration of theshortening process is estimated at about 20 Myr. Plašienka[1997] also demonstrated that the original frontal closure ofthe Fatric Basin passed to transpressive movements after20 Myr. This means that the differential movement of rigidindenter, which moves together with the whole kinematicsystem, has to generate a defined finite strain at the sameperiod of time. Moreover, the initiation of TGSZ activity maycorrespond to a transition from frontal to transpressionalmovements recorded in the northern edge of the wholekinematic system. Once this rough timescale is established,then the absolute velocity of our indenter should be four timesslower than suggested in the model to generate the observedstrain pattern.5.3. Development of Topography, Exhumation, andAsymmetry of GCF[45] The model allows estimation of average verticalstrains and, because of a fixed lower boundary condition,also the vertical elevation. We can expect that the surfaceelevation represent local topography generated by shorteningof the viscous sheet. The lateral distribution of topographyfollows the exponential distribution of finite strain inareas of pure shear-dominated deformation. Figure 8dshows the distribution of topography in front of an indentingblock after 7 Myr of shortening. It is to be noted that thedomain of highest topography follows the axial zone of theGCF, where the degree of metamorphism associated withthe development of cleavage is most important.[46] Although our model predicts vertical cleavage in theentire domain, we observe that the cleavage forms a positivefan-like structure. We interpret this pattern as a result ofdifferent amount of vertical shortening due to differentgravitational potential across the GCF. This mechanism ismanifested by the development of late kink bands with kinkplanes perpendicular or oblique to strongly developed verticalcleavage.Appendix A[47] The equations governing the deformation of a thinviscous sheet were published by England and McKenzie[1982]. We provide here the derivation of equations for thesimplest case of Newtonian rheology as they have beenused for our modeling. The model assumes a relatively thinviscous plate with no tractions at top and bottom surfaceand negligible vertical gradients of horizontal velocitycomponents. Creep equation reads@t ij@x j¼ @p@x iðA1Þwhere T is the deviatoric stress tensor and p is the pressure,and repeated index means summation. Assuming a linearconstitutive relationt ij ¼ 2h_e ijequation (1) can be written as2h @_e ij@x j¼ @p@x i[48] The strain rate tensor of the form01_e 11 _e 12 0_E ¼_e 21 _e 22 0BC@A0 0 _e 33ðA2ÞðA3ÞðA4Þis assumed. Then the equation containing the vertical strainrate _e 33 reduces to2h @_e 33@x 3¼ @p@x 3ðA5Þ[49] Integrating the equation over the vertical dimensionx 3 , we obtain2h_e 33 ¼ p þ fðx 1 ; x 2 Þ ðA6Þwhere the upper strike means vertical average. Because ofthe model assumptions we can put f (x 1 ,x 2 ) = 0 everywhere.We substitute from equation (A6) in the first equation (A3)integrated over the vertical dimension and obtain2h @ _e 1j@x j¼ @p@x 1¼ 2h @ _e 33@x 1and similarly for the second equation. The resulting equationscan be written as@_e ij@x j@_e 33@x i¼ 0 i ¼ 1; 2; j ¼ 1; 2 ðA7Þwhere we omit the signs of vertical averaging. Massconservation for incompressible flow requires that_e 33 ¼ ð_e 11 þ _e 22 Þ ðA8Þ98

LEXA ET AL.: COLLISION IN WEST CARPATHIANS 5 - 15Thus2 @_e 11þ @_e 12þ @_e 22¼ 0@x 1 @x 2 @x 1@_e 11þ @_e 21þ 2 @_e 22¼ 0@x 2 @x 1 @x 2ðA9Þ[52] Although we compute the horizontal velocity field,the vertical strain rate can be estimated by equation (8) thatallows us to assess at every step the tensors of instantaneousand finite deformation and related parameters of deformation,such as intensity[50] We replace the strain rate tensor by horizontalcomponents of velocity u 1 , u 2_e ij ¼ 1 @u iþ @u jðA10Þ2 @x j @x iand finally obtain a system of elliptic partial differentialequationssymmetryqffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiD ¼ R xy 1 2þ Ryz 1 2ðA12ÞK ¼ R xy 1 R yz 1 ðA13Þ4 @2 u 1@x 2 1@ 2 u 2@x 2 1þ @2 u 1@x 2 þ 3 @2 u 2¼ 02@x 1 @x 2þ 4 @2 u 2@x 2 þ 3 @2 u 1¼ 02@x 1 @x 2ðA11Þand orientations of lineations and foliations. Because of theassumed form of the strain rate tensor, equation (A4), themodel can produce only horizontal or vertical lineations andfoliations and can be regarded as generalized (locallyvariable) transpression.[51] The system can be solved by the finite elementmethod. Dirichlet and Neumann boundary conditions maybe applied to segments of the domain boundaries so thatthey correspond to geological settings (rigid indenter, freeinflow or outflow of material).[53] Acknowledgments. We are grateful to the Geological Survey ofthe Slovak Republic for significant financial support during the initialstages of our research. This work has been also supported by the CharlesUniversity Agency grant 216/1999/B-GEO. The salaries of K. Schulmannand O. Lexa were covered by Ministry of Education grant 24313005.ReferencesAllemand, P., and J. M. Lardeaux, Strain partitioningand metamorphism in a deformable orogenicwedge: Application to the Alpine belt, Tectonophysics,280, 157 – 169, 1997.Andrusov, D., Subtatranské príkrovy v Západných Karpatoch,Carpatica, 1, 3 – 50, 1936.Andrusov, D., Geológia Československých Karpát Zv.1, Vydavatel’stvo SAV, Bratislava, Slovakia, 1958.Bell, T. H., and M. J. Rubenach, Sequential porphyroblastgrowth and crenulation cleavage developmentduring progressive deformation, Tectonophysics,92, 171 – 194, 1983.Biely, A., and J. Bystrický, Mesozoic of the inner westCarpathians and the klippen belt, in Guide toExcursion, 44 pp., Int. Geol. Congr., Prague, 1967.Biely, A., J. Bystrický, and O. Fusán, De l’appartenancedes nappes des Karpates occidentales internesTranslated Title: Relations of the nappes of theinner western Carpathians, paper presented at Int.Geol. 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Journal of Structural Geology 26 (2004) 155–161www.elsevier.com/locate/jsgApparent shear-band geometry resulting fromoblique fold sectionsOndrej Lexa a, *, John Cosgrove b , Karel Schulmann aa Institute of Petrology and Structural Geology, Charles University, Prague, Czech Republicb Department of Earth Sciences and Engineering, Royal School of Mines, Imperial College, London SW7 2BP, UKReceived 1 May 2002; received in revised form 1 February 2003; accepted 4 April 2003AbstractSmall-scale shear zones inclined at intermediate angles to an earlier anisotropy are often observed in deformed rocks. They aretraditionally described as shear-bands, C-bands, extensional crenulation cleavage or normal kink-bands formed as a result of extension alongthe anisotropy. Their asymmetries are widely used to describe the large-scale kinematics of deformation and the deformational history of agiven area. We demonstrate that when various three-dimensional fold structures are observed on two-dimensional outcrop surfaces or in thinsection, they can appear geometrically identical. We have developed a simple technique that allows the geometrical evaluation of any sectionacross a cylindrical fold of arbitrary geometry. The ranges of planar sections on which a fold exhibits shear-band like geometry are presentedon a stereographic projection in order to simplify the determination of critical orientations. We demonstrate that for any fold geometry, thereare two distinct groups of sections showing shear-band like geometry with opposite ‘senses of shear’ criteria systematically arranged aroundthe axial plane and which are inclined at a high angle to the major anisotropy. We provide a field example from Western Carpathians, wherekinematic analysis, mainly based on apparent extensional shear-bands, led to overemphasis of the role of post-orogenic extension on the finalstructural pattern of the belt.q 2003 Elsevier Ltd. All rights reserved.Keywords: Shear-band geometry; Oblique fold sections; Small-scale shear zones1. IntroductionSmall-scale shear zones inclined at intermediate angles to aprevious anisotropy are commonly observed in deformedrocks. They are traditionally presented as shear-bands (White,1979), C 0 -bands (Ponce and Choukroune, 1980), extensionalcrenulation cleavage (Platt, 1979, 1984; Platt and Vissers,1980), asymmetric boudinage, asymmetric folds or normalkink-bands (Dewey, 1965; Cobbold et al., 1971; Cosgrove,1976) formed as a result of extension along the olderanisotropy or shortening normal to the anisotropy. Their‘sense of shear’ and geometrical relations are widely used todescribe the large-scale kinematics of deformation (Berthéet al., 1979; Simpson and Schmid, 1983; Lister and Snoke,1984) or the tectonic settings of the deformational history(Platt and Vissers, 1980; Behrmann, 1987).Shear bands may resemble the compressional crenulation* Corresponding author. Tel.: þ420-22195-1531; fax: þ420-22195-1524.E-mail address: lexa@natur.cuni.cz (O. Lexa).0191-8141/03/$ - see front matter q 2003 Elsevier Ltd. All rights reserved.doi:10.1016/S0191-8141(03)00072-5cleavage but develop by extension of the older foliationrather than by shortening (Passchier and Trouw, 1996). Thisled some authors to use the terms compressional (CCC) andextensional (SBC) crenulation cleavages (Platt and Vissers,1980). Passchier and Trouw (1996) presented a summary ofdifferences between these two contrasting structures. Theirmain argument for distinction between both kinds ofstructures is the angle of CCC with the older foliation,which generally ranges between 45 and 908, while for SBCthe angle to earlier foliation is less than 458. However, theangular distinction between CCC and SBC is not alwaysvalid. The compressional crenulation cleavage changes thegeometry in the profile section towards the hinge directionof the folded domain, so that the internal rotation becomesless than 458 and may be easily misinterpreted as an SBC(Price and Cosgrove, 1994, p. 263, Fig. 10.50).From a kinematic point of view, CCC develops at a highangle to bulk shortening while SBC represents a single shearplane at small angle to the foliation (Passchier and Trouw,1996). In order to interpret the kinematic significance ofboth kinds of structures, they have to be observed in plane,101

156O. Lexa et al. / Journal of Structural Geology 26 (2004) 155–161Fig. 1. Shear band geometry and definition of angles a, b and g according toPlatt and Vissers (1980).perpendicular to the intersection of CCC and SBC with theolder foliation.With the advent of modern kinematic analysis in structuralgeology in the early 1980’s, the XZ plane of finite strainellipsoid became extremely important. This plane is traditionallydefined as a plane parallel to the stretching lineation andperpendicular to the foliation. However, the stretchinglineation in phyllites or phyllonites can be difficult to define,and it can be easily confused with corrugations or intersectionlineation. Moreover, the presence of well-defined shear bandson a rock surface in the field is in many cases considered asatisfactory indicator to consider this surface as an XZ plane.Because, shear bands are such noticeable structures and theirkinematic significance is straightforward, these structureshave been widely used as first order kinematic indicators inmany orogenic belts (Behrmann, 1987).This paper aims to demonstrate that: (1) distinguishingbetween CCC and SBC is not always an easy task, (2) unlessthis is fully appreciated there is a great danger ofmisinterpretation when the shear bands are used as kinematicindicators, and (3) the CCC and SBC can appear identicalwhen seen on flat outcrop surfaces or in thin sections.2. Geometrical characteristics of oblique sections of foldsFolding or flow partitioning are commonly presented intwo dimensions while geological structures are threedimensional(3D) features. In order to justify the use ofthe two-dimensional (2D) analyses to investigate a threedimensionalproblem, certain assumptions are made.Because the displacement fields predicted by the 2D-theoryare limited to a plane (usually one of the principal planes ofthe strain ellipsoid), it is assumed that displacement isidentical in any parallel plane and therefore that theresulting fold structures are cylindrical. As a result, theaxes of the folds are perpendicular to the plane ofthe displacement field. Based on these assumptions wewill consider 2D sections normal to the fold axes to be ‘true’sections and all other sections ‘apparent’.Geometrically, shear-bands can be characterized by threeangles a, b and g (Fig. 1), corresponding to the relativeangles of the limbs, hinge trace and enveloping surface(Platt and Vissers, 1980). In order for the fold or flexureprofiles to have shear-band geometry, a should have a valuebetween 10 and 508, while b and g should range between 10and 258, so that the interlimb angle exceeds 1308. Platt andVissers (1980) show that these angles are mean values fromobservations in the field, for example from the Beticmovement zone (Behrmann, 1987). We will show that for agiven fold profile in the ‘true’ section (i.e. a section normalto the fold axis; Fig. 2) there is a specific range of ‘apparent’sections (i.e. a section oblique to the fold axis; Fig. 2) inwhich the fold profiles satisfy the geometrical criteria ofshear-bands.Apparent sections through three-dimensional cylindrical,coaxial folds can be treated as fold axis parallel projectionsof ‘true’ fold section onto the section plane. When ax þby þ cz ¼ 0 is the equation of a section plane and ( p, q, r)isthe vector parallel to the fold axis in an arbitrarily chosenreference frame, then the coordinates ½x yzŠ of points in a‘true’ section and the coordinates ½x 0 y 0 z 0 Š of points on an‘apparent’ section are related to each other by matrixequation:23bq þ cr 2ap 2arD D D2bp ap þ cr 2br½x yzŠ6 D D D¼½x 0 y 0 z 0 Š7642cpD2cqDap þ bqDwhere D ¼ðap þ bq þ crÞ.The 2D coordinates on the chosen section are obtained ascoefficients of linear combinations of any two perpendicularunit vectors coplanar with the section plane. We developeda simple MATLABw script, which visualises any sectionprofile across a given fold geometry and determine the threeangles a, b and g (Fig. 1) for the oblique section profile.Using this script we can examine any section across a fixedfold geometry and determine which of them satisfies thegeometric requirements for shear-bands. The results arepresented in a stereographic projection (Figs. 3 and 4)sharing the area containing poles to sections that satisfy theshear-band criteria. The area of sections exhibiting shearbandlike geometry is shaded on the basis of the interlimbangle a (Fig. 1).Sections through symmetrical folds generally have foldprofiles with an asymmetric geometry (Fig. 3) and theplanes on which the apparent fold geometry satisfies theshear-band criteria fall into two distinct groups separated bythe fold axial plane. These two groups contain poles tosections on which shear-band geometry exhibit opposingasymmetry, i.e. opposing ‘shear-sense’ criteria. The area ofthese domains on a stereographic projection, i.e. the range oforientations of section planes in space that display shearbandgeometries is related to the interlimb angle of thesymmetrical folds, so as the interlimb angle increasestowards 1808, the range of suitable sections displayingshear-band geometry increases.Sections across asymmetric folds that display shear-band75102

O. Lexa et al. / Journal of Structural Geology 26 (2004) 155–161 157Fig. 2. (a) Block diagrams of folded anisotropy and ‘apparent’ fold sections showing shear-band geometries. (b) Orientation of apparent fold sectionsexhibiting opposite sense of shear criteria. (c), (d) Principle of presentation of ‘apparent’ fold sections in stereographic projection.geometries are shown in Fig. 4. It can be seen that the range ofsections suggesting a sinistral and dextral ‘sense of shear’ isunequal in this example and a preferred enlargement of onegroup occurs. This asymmetry is controlled by the anglebetween the fold axial plane and the main anisotropy (Fig. 4).3. Geological exampleIn the Vepor basement of the Central WesternCarpathians, a flat-lying mylonitic foliation containing awell-developed stretching lineation is affected by late103

158O. Lexa et al. / Journal of Structural Geology 26 (2004) 155–161Fig. 3. Stereographic projection showing influence of interlimb angle. The range of section planes showing shear-band geometry increases with interlimb angleof folds affecting main anisotropy.Fig. 4. Stereographic projections showing influence of angle between fold axial plane and main anisotropy. The probability of occurrence of opposite shearcriteria decreases with decreasing angle of axial plane and plane of main anisotropy.104

O. Lexa et al. / Journal of Structural Geology 26 (2004) 155–161 159Fig. 5. Structural elements from studied area. (a) Poles to main Alpine metamorphic anisotropy showing girdle distribution around the axis sub-parallel tohinges of late Alpine folds—116 measurements; (b) orientation of stretching lineation—68 measurements; (c) distribution of poles to main fracture systemsdeveloped in studied area (Hók et al., 2001) overlaid by shaded areas indicating ‘apparent’ fold sections in which shear-band geometry can be observed; (d)rose diagram of main fracture directions in studied area (Hók et al., 2001). Note that a main maximum of fracture orientations coincides with orientations ofapparent fold sections. All data are plotted on Schmidt net and projected from lower hemisphere. In (a) and (b), the contour levels are even multiples of standarddeviation.folding, which generated a crenulation cleavage. The foldaxes are generally sub-parallel to the stretching lineation(Fig. 5). Most of the studies on the kinematic andtectonic evolution of this region are based on the use ofshear-bands and asymmetric porphyroclasts as kinematicindicators. These studies have resulted in the generallyaccepted model of post-orogenic, orogen-parallel extension(Plašienka et al., 1999; Lupták et al., 2000; Janáket al., 2001). In contrast to the interpretation given byprevious workers in which the major extensionaldeformation post-dates the episode of compression, astudy by these authors shows that extension was the firstAlpine deformation in the area and that this pre-dates thecompressional stage of tectonic evolution (Lexa et al.,2003). The extensional phase is highly asymmetrical andlocally non-coaxial. The majority of the studied outcropsshow that we are dealing with oblique sections acrosssmall-scale folds, which are likely to be misinterpreted asshear-bands (Fig. 6).In order to evaluate the probability of encounteringoblique sections with shear-band like geometry in aparticular field area, we need to identify the dominantgeometries of the small-scale folds and, moreover, wehave to understand the distribution of surfaces on whichthe structures are observed. It should be pointed out thatthe majority of observations are from natural rather thanman-made outcrops, where the orientation of the exposedsurfaces is controlled mainly by fractures (typicallyjoints). In addition, sections that are sub-parallel to thelineations are specifically selected as being appropriatefor the study of kinematic indicators. We plotted therange of sections with shear-band like geometry, therange of naturally occurring fractures and the distributionof lineations on a stereographic projection (Fig. 5c),which shows that there is a high probability of systematicallyobserving oblique sections across the small-scalefolds showing opposite shear-band geometries. Wepropose that such field observations led some authors105

160O. Lexa et al. / Journal of Structural Geology 26 (2004) 155–161Fig. 6. Field photos of late folding and crenulation cleavage affecting early Alpine fabric in the Vepor micaschists of the West Carpathians. True fold sections(a) micaschists in the Mútnik area; (b) micaschists in the Katarínska huta area. Apparent fold sections showing shear-bands geometry (c), (d) Katarínska huta.Axial plane: 326/75; fold axis: 52/14; and general outcrop surface: 318/65.(e.g. Siman et al., 1996) to interpret the extensionaltectonics in Vepor basement to be symmetrical.4. Discussion and conclusionsShear-bands and folds can generate identical geometrieswhen seen on a flat exposure surfaces in some orientations.These geometries can lead to misinterpretation of folds asshear bands and to erroneous structural and kinematicinterpretation.The ambiguity arises since oblique sections throughsmall-scale folds in anisotropic materials have identicalgeometries to XZ finite strain sections through shear-bandbearing rocks. Misinterpretations are most likely in areas ofmultiple deformations when earlier fabrics are folded.Determination of the XZ section in phyllonites and phyllitesis often complicated by the presence of microscopiccorrugations on foliation surfaces, which may be easilyconfused with stretching lineation. These corrugations arecommonly associated with gentle folds of a larger scale.Particularly in this case, geologists would tend to look forkinematic criteria in sections in which the larger folds wouldgenerate profiles similar to shear band geometry.In order to be confident that the geometry displayed on a2D exposure surface is related to true shear-bands and cantherefore be used as a valid kinematic indicator, it isessential to know the 3D geometry of the structure and theorientation of the exposure surface with respect to fold axesand lineations. For this reason, systematic studies of fractureand joint systems in folded areas should accompanykinematic analyses.106

O. Lexa et al. / Journal of Structural Geology 26 (2004) 155–161 161AcknowledgementsWe are grateful to the Geological Survey of the SlovakRepublic for significant financial support during the initialstages of our research. This work has been also supported bythe Grant of Charles University Agency No. 216/1999/B-GEO. The salaries of K. Schulmann and O. Lexa werecovered from the grant of the Ministry of Education No.24313005.ReferencesBehrmann, J.H., 1987. A precautionary note on shear bands as kinematicindicators. In: Cobbold, P.R., Gapais, D., Means, W.D., Treagus, S.H.(Eds.), Shear Criteria in Rocks. Journal of Structural Geology 9(5–6).Pergamon, Oxford–New York, International, pp. 659–666.Berthé, D., Choukroune, P., Jegouzo, P., 1979. Orthogneiss, mylonite andnon coaxial deformation of granites; the example of South Armoricanshear zone. Journal of Structural Geology 1 (1), 31–42.Cobbold, P.R., Cosgrove, J.W., Summers, J.M., 1971. Development ofinternal structures in deformed anisotropic rocks. Tectonophysics 12,23–53.Cosgrove, J.W., 1976. The formation of crenulation cleavage. Journal of theGeological Society of London 132 (Part 2), 155–178.Dewey, J.F., 1965. Nature and origin of kink-bands. Tectonophysics 1 (6),459–494.Hók, J., Lacika, J., Madarás, J., Kohút, M., Nagy, A., Ivanička, J., Siman,P., Král’, J., Töröková, I., 2001. Neotektonický a geomorfologickývývoj študijných lokalít – 1. čast’. Projekt: Vývoj hlbinného úložiskavyhoreného jadrového paliva a vysokoaktívnych Ra – odpadov vpodmienkach Slovenskej republiky pre obdobie 1998–2000. Úloha:Výber lokality. Číslo etapy: VYL-01-00. Štátny geologický ústavDionýza Štúra, Bratislava. Manuskript–archív ŠGÚDŠ, Bratislava,pp. 1–171.Janák, M., Plašienka, D., Frey, M., Cosca, M., Schmidt, S.T., Lupták, B.,Méres, S., 2001. Cretaceous evolution of a metamorphic core complex,the Veporic unit, Western Carpathians (Slovakia): P–T conditions andin situ Ar-40/Ar-39 UV laser probe dating of metapelites. Journal ofMetamorphic Geology 19 (2), 197–216.Lexa, O., Schulmann, K., Ježek, J., 2003. Cretaceous collision andindentation in SW part of West Carpathians:view based on structuralanalysis and numerical modeling. Tectonics, accepted.Lister, G.S., Snoke, A.W., 1984. S–C mylonites. Journal of StructuralGeology 6 (6), 617–638.Lupták, B., Janák, M., Plašienka, D., Schimdt, S.T., Frey, M., 2000.Chloritoid-kyanite schists from the Veporic unit, Western Carpathians,Slovakia: implications for Alpine (Cretaceous) metamorphism. SchweizerischeMineralogische Und Petrographische Mitteilungen 80 (2),213–223.Passchier, C.W., Trouw, R.A.J., 1996. Microtectonics, Springer-Verlag,Berlin, Heidelberg.Plašienka, D.a., Janák, M., Lupták, B., Milovskú, R., Frey, M., 1999.Kinematics and metamorphism of a Cretaceous core complex: theVeporiuc Unit of the Western Carpathians. Physics and Chemistry ofthe Earth (A) 24 (8), 651–658.Platt, J.P., 1979. Extensional crenulation cleavage. In: Cobbold, P.R.,Ferguson, C.C. (Eds.), Description and Origin of Spatial Periodicity inTectonic Structures; Report on a Tectonic Studies Group Conference.Journal of Structural Geology 1. Pergamon, Oxford–New York,International, pp. 95–96.Platt, J.P., 1984. Secondary cleavages in ductile shear zones. Journal ofStructural Geology 6 (4), 439–442.Platt, J.P., Vissers, R.L.M., 1980. Extensional structures in anisotropicrocks. Journal of Structural Geology 2 (4), 397–410.Ponce, d.L.M.I., Choukroune, P., 1980. Shear zones in the Iberian Arc. In:Carreras, J., Cobbold, P.R., Ramsay, J.G., White, S.H. (Eds.), ShearZones in Rocks. Journal of Structural Geology 2(1/2). Pergamon,Oxford–New York, International, pp. 63–68.Price, N.J., Cosgrove, J.W., 1994. Analysis of Geological Structures,Cambridge University Press, Cambridge.Siman, P., Johan, V., Ledru, P., Bezák, V., Madarás, J., 1996. Deformationand P–T conditions estimated in “layered migmatites” from southernpart of Veporicum crystalline basement (Western Carpathians,Slovakia). Slovak Geological Magazine 3–4/96, 209–213.Simpson, C., Schmid, S.M., 1983. An evaluation of criteria to deduce thesense of movement in sheared rocks. Geological Society of AmericaBulletin 94 (11), 1281–1288.White, S., 1979. Large strain deformation; report on a Tectonic StudiesGroup discussion meeting held at Imperial College, London on 14November 1979. Journal of Structural Geology 1 (4), 333–339.107

DTD 5ARTICLE IN PRESSJournal of Structural Geology xx (xxxx) 1–24www.elsevier.com/locate/jsgThe quantitative link between fold geometry, mineral fabric andmechanical anisotropy: as exemplified by the deformation of amphibolitesacross a regional metamorphic gradientLenka Baratoux a, *, Ondrej Lexa a,b , John W. Cosgrove c , Karel Schulmann ba Institute of Petrology and Structural Geology, Charles University, Albertov 6, Praha 2, 12843, Czech Republicb Université Louis Pasteur, EOST, UMR 7517, 1 Rue Blessig, Strasbourg 67084, Francec Department of Earth Science & Engineering, Royal School of Mines, Imperial College of Science, Technology & Medicine, Prince Consort Road,London SW7 2BP, UKReceived 8 June 2004AbstractThis work shows lateral variations in fold geometry within an amphibolite unit of constant mineralogical composition under increasingmetamorphic grade. Analysis of the fold geometries indicate: (1) medium amplification associated with low post-buckle flattening in thegarnet zone, (2) high amplification coupled with medium post-buckle flattening in the staurolite zone, and (3) passive amplificationdominated by intense post-buckle flattening in the sillimanite zone. A systematic decrease in the mechanical anisotropy of the folded fabric isobserved with increase in metamorphic grade. Quantitative microstructural study shows contrasting deformation micro-mechanismsassociated with folding manifested by: (1) brittle dominated deformation of amphiboles that form a stress supporting network with a highcompetence contrast with respect to plagioclase in the garnet zone, (2) ductile dominated heterogeneous deformation of an interconnectedweak layer structure with low competence contrast in the staurolite zone, and (3) homogeneous deformation of a stress supporting frameworkwith low competence contrast in the sillimanite zone. The difference in the folding style between the garnet and staurolite zones is associatedwith the lateral variations in microstructure of the amphibolites inherited from a pre-folding metamorphic zonation and with differentdeformation mechanisms in the hinge zones. However, the change in fold style observed as one moves into the sillimanite zone is controlledby the recrystallization associated with an important syn-folding heat input from an adjacent granite intrusion.q 2005 Elsevier Ltd. All rights reserved.Keywords: Fold geometry; Quantitative link; Amphibolite; Metamorphic gradient; Deformation microstructures1. IntroductionThe fold shape in a folded bilaminate is determined bythe ratio of competent to incompetent layer thickness andthe viscosity contrast (m 1 /m 2 ) between them (Ramberg,1961). Ramsay and Huber (1987, p. 405) discuss the factorsthat influence the fold geometry of such buckled bilaminates.The most important factors are the primaryrheological properties of the layers, which depend on themineralogical composition and grain size, external parameterssuch as temperature and pressure, and the* Corresponding author. Tel.: C420-257189509E-mail address: lka@natur.cuni.cz (L. Baratoux).0191-8141/$ - see front matter q 2005 Elsevier Ltd. All rights reserved.doi:10.1016/j.jsg.2005.01.001development of a preferred orientation of minerals duringdeformation. The seminal work of Biot (1961) provided atheoretical treatment capable of analysing these systems andhis work has been successfully applied by geologists toexplain the buckling behaviour of both discretely layeredsystems and mineral fabrics. In Biot’s work the mostimportant factor controlling the mode and style of folding isthe mechanical anisotropy of the material. It can be shownthat there is a direct link between the mechanical anisotropyand the competence contrast of a bilaminate (m 1 /m 2 ) and thisallows one to use both theories to study their foldingbehaviour (Price and Cosgrove, 1990). By analysing thegeometries of the studied folds it is possible to determinewhich of the two theoretical approaches is more appropriate.Although a large quantity of work has been carried out onthe folding mechanics of layered systems (e.g. thetheoretical work by Biot (1961, 1963a,b), Ramberg (1963,109

DTD 5ARTICLE IN PRESS2L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–241964) and Fletcher (1995), the experimental studies byGhosh (1968), Dubey and Cobbold (1977), Dubey (1980),Mancktelow and Abbassi (1992) and Hobbs (1971) and thefield studies by Oertel and Ernst (1978) and Fowler andWinsor (1996)), this work is not directly applicable to thefolding of mineral fabrics where the concept of layerthickness and competence contrast are more difficult todefine. Fortunately a theory of the deformation ofmechanically anisotropic materials has been developed byBiot (1967), which enables the folding of mineral fabrics tobe considered.The application of modern computer graphics, inparticular geographic information systems (GIS) in quantitativepetrography, and computer techniques, which allowthe precise evaluation of grain shapes (Panozzo, 1983,1984), grain sizes and grain spatial distribution, haveprovided tools for the assessment of the relationshipbetween mineral fabrics and the folding that developswithin them. In this paper we present a method ofquantitative microstructural analysis that allows the semiquantitativeassessment of the mechanical anisotropy ofmineral fabrics. We demonstrate that quantitative computeraidedanalysis of natural fold shapes in conjunction with astudy of the internal microstructure allows a better understandingof the mechanical behaviour of naturally foldedmultilayers. Our observations show that it is possible todistinguish folds formed by active buckling from thosedominantly controlled by passive amplification. This can beclearly correlated with major variations in mineral microstructures,semi-quantitatively estimated mechanical anisotropyand the deformation mechanisms associated withfolding.The rock unit studied in this analysis consists ofmetabasites of relatively constant composition that showmetamorphic zonality ranging from the lower amphibolitefacies conditions in the east to the upper amphibolite faciesconditions in the west. The metamorphism wasaccompanied by the development of progressively evolvingmineral microfabrics forming a gently dipping to subhorizontalmetamorphic foliation. This consistently orientedbut systematically varying rock anisotropy has beensubsequently affected by horizontal compression responsiblefor the development of upright post-metamorphicfolds. In addition, a granite intruded into the western part ofthe amphibolite sheet and thermally influenced the host rockwithin a zone several kilometres wide. Field observationsshow a remarkable change in the style of the post-peakmetamorphic folding approaching the intrusion (Schulmannand Gayer, 2000) and the question arises as to whether thechange in fold style reflects the variation in theinherited mechanical anisotropy associated withthe east–west metamorphic–microfabric zonation,whether it is related to the lateral heat input from thewesterly granite, or both.The aims of this contribution are: (1) the quantitativecharacterization of the fold shapes and semi-quantitativedetermination of mechanical anisotropy based on quantitativemicrostructural analysis, (2) the determination ofpossible folding mechanisms using fold shape analysis aswell as the micro-deformational mechanisms associatedwith folding, and (3) the evaluation of the relative role of theinherited paleometamorphic microstructural zonation(mechanical anisotropy) and lateral heating, on folding.2. Geological settingThe northeastern margin of the Bohemian massif is builtup of the Neo-Proterozoic basement called Brunia and itssedimentary cover (Dudek, 1980) in the east, and a belt ofBrunia derived Variscan nappes (Silesian domain) with highgrade rocks of the orogenic root belonging to the Lugiandomain in the west (Fig. 1).The studied area is located in the Brunia derived parautochthoncalled the Desná dome, which consists ofmetamorphic Neo-Proterozoic basement core surroundedby a volcano-sedimentary Devonian envelope (Fig. 1b).Metabasites of the Jeseník amphibolite massif, the subjectof this study, form part of this Devonian sequence. The hugeaccumulation of basic rocks was formed during the EarlyDevonian rifting, as indicated by their tectono-stratigraphicposition and back-arc tholeiitic composition (Patočka andValenta, 1990). Metabasites are present on both the easternand western flanks of the Desná gneissic core. The basites atthe eastern margin show more tuffitic character while themain amphibolite body situated in the west developedmostly from basalts and possibly from gabbros as indicatedby preserved textures in the less deformed domains. Themetabasites are locally stratigraphically or tectonicallyintercalated with metapelites and quartzites of Pragian age(Chlupáč, 1994), which show the same tectono-metamorphichistory.2.1. Structural geologyThe complex polyphase tectonic evolution of the regioncomprises the early pre-Variscan deformations termed D 1 .The first Variscan structures are associated with theBarrovian metamorphism (termed D 2 ), and the late Variscanstructures (D 3 ), developed during the transpressionalregime. The most important D 2 structure is represented bya planar metamorphic fabric expressed as localized shearzones in the basement orthogneisses and as a penetrative,originally sub-horizontal metamorphic foliation (S 2 ) in allthe rocks of the Devonian cover. This fabric is now dippingto the NW or SE at moderate to steep angles as a result ofsubsequent large-scale folding (Fig. 2a). Within the cover,syn-schistose folds (F 2 ) as well as mineral lineation (L 2 ) arelocally well developed.The D 3 deformation had various effects on differentlithologies, depending mainly on the intensity of theprevious S 2 anisotropy (Schulmann and Gayer, 2000).110

DTD 5ARTICLE IN PRESSL. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 3Fig. 1. (a) Geological map of the eastern margin of the Bohemian Massif. Inset shows location of the studied area in the frame of European Variscides. (b).Geological map of the Silesian domain based on the geological map 1:200,000 by Pouba (1962) and Roth (1962). Important thrust faults and normal faults areindicated. A–A 0 indicates the position of the cross-section in Fig. 2.Basement orthogneisses display rare upright folding.Metabasites and metasediments of the Devonian volcanosedimentarycover show well-developed mostly asymmetricF 3 folds with axial planes dipping to the NW or SE andsubhorizontal NE–SW-trending hinges (Fig. 2a). The foldsize ranges from the microscopic grain-scale up to metresscalefolds. In the eastern and central parts of the Desnádome the D 3 deformation was post-peak metamorphic andtook place under greenschist conditions as marked bygrowth of chlorite, muscovite and albite along axialcleavage of F 3 folds in metapelites (Schulmann andGayer, 2000). However, it was contemporaneous with theemplacement of the Žulová granite in the west (Cháb andŽáček, 1994; Schulmann and Gayer, 2000), which resultedin the overprinting of the S 2 foliation by a NE–SW-trendingsteep and penetrative foliation S 2–3 (Fig. 2a).111

DTD 5ARTICLE IN PRESS4L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24Fig. 2. Geological cross-section A–A 0 (shown in Fig. 1b) of the Desná dome showing major structures, lithology of individual units, and major tectonicboundaries (a). Metamorphic zones extended from metapelites are indicated as well as approximate actinolite-out and clinopyroxene-in isograds (for Ca-richmetabasites). Equal area lower hemisphere stereoplots are shown for D 2 , and D 3 planar and linear structures. Each stereoplot contains between 50 and 100 polescontoured as multiples of uniform distribution. The cross-section is constructed from photographs and field notes and shows the principal structural features.The vertical axis is not to scale. Note that the metamorphic isograds cross-cut lithological boundaries and their apparent steep attitude is associated with a largescale F 3 folding of their original flat dip as shown by Schulmann and Gayer (2000) and Štípská et al. (2000). (b) Lower left inset shows PT plot with indicatedages of M 3 and M 2 metamorphisms and associated fabrics; a—Schulmann and Gayer (2000); b—Jehlička (1995); c—Maluski et al. (1995). PA — plutonaureole.112

DTD 5ARTICLE IN PRESSL. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 52.2. Metamorphic zonalityThe Variscan metamorphic evolution of the study areainvolves two main phases; an M 2 metamorphism ofBarrovian character with the intensity increasing westwardsand an M 3 metamorphism induced by the Žulová graniteintrusion in the western margin of the Desná dome. TheBarrovian M 2 metamorphic grade ranges from chlorite zonein the eastern margin of the Desná dome up to staurolite andpossibly sillimanite zone in the west. The M 3 periplutonicHT/LP overprint also attains its maximum in the westernpart of the studied area, where it is documented by thepresence of K-feldspar–cordierite migmatites (Rozkošnýand Souček, 1989; Cháb and Žáček, 1994), as well as by thegrowth of sillimanite and new garnet in the staurolitemicaschists in the pluton aureole (Fig. 2).Although the metabasites cannot be precisely dividedinto particular metamorphic zones once they have reachedamphibolite facies conditions, it is nevertheless necessaryfor the purpose of this work to establish a metamorphiczoning of the amphibolite massif based on the increasingmetamorphic conditions. The metamorphic zones determinedin intercalated metasedimentary rocks by Souček(1978) are therefore extended into the adjacent metabasitesand used as a reference for the definition of the metamorphicgrade therein. The justification for this procedure is based onthe mutual field relations of the amphibolites and metasediments,which indicate that both lithologies have experiencedthe same tectonic and metamorphic history(Schulmann and Gayer, 2000). Therefore, PT conditionswere determined in metasediments applying variousthermobarometry methods (Baratoux, 2004). The degreeof metamorphism in the east of the metabasite massifcorresponds to the garnet zone in the contiguous metasedimentsand the mineral assemblage in the metabasitesincludes hornblende, plagioclase, actinolite, chlorite, andepidote corresponding to PT conditions of 540G10 8C and5G1 kbar. Further to the west actinolite and epidotedisappear and HblCPlGQtzCIlmCTtn, which correspondsto the staurolite zone, appears with metamorphicconditions of M 2 estimated to be 570G30 8C and 5.5G1 kbar (Baratoux, 2004). The mineral assemblage in themetabasites corresponding to the sillimanite zone isrepresented by HblCPlGQtzCIlmCTtn and AmpCPlGQtzGCpxGCalCIlmCTtn in calcium-rich lithologies andmetamorphic conditions of this zone were estimated to590G20 8C and 5.5G1 kbar. In the pluton aureole,sillimanite–cordierite–K-feldspar assemblage occur inmetapelites and the PT conditions of periplutonic M 3metamorphism reached 700G15 8C and 4.2G0.8 kbar (Fig.2b). The mineral assemblage in the amphibolites correspondsto that of the sillimanite zone.The age of the main fabric-forming M 2 metamorphicevent is difficult to establish in the studied area, but it issupposed to have occurred during the main collisionalevent, which is dated elsewhere at w340 Ma (Schulmannand Gayer, 2000; Štípská et al., 2004). The termination ofthe D 2 –M 2 event can be constrained by Rb–Sr dating (335G7.5 Ma) of the crystallization of the Žulová granite(Jehlička, 1995). However, 40 Ar/ 39 Ar dating of muscoviteand biotite from mylonitic gneisses of the Desná dome andof the Žulová granite (Maluski et al., 1995) suggest thatcooling through the white mica and biotite closuretemperatures (350 and 300 8C, respectively) occurredbetween 310 and 300 Ma. Therefore, this age maycorrespond to the greenschist facies F 3 folding activity inthe east and to the cooling of the Žulová pluton to the west.3. D 2 and D 3 amphibolite microstructures acrossmetamorphic zones3.1. Eastern part of the massif (the garnet zone)In the eastern part of the massif, i.e. in the garnet zone,some weakly deformed metagabbros with large grains ofhornblende and plagioclase (1–2 mm) are still present. Themain metamorphic fabric S 2 of amphibolites is preserved indomains unaffected by F 3 folding or in F 3 fold limbs and it ischaracterized by a strong mineral shape preferred orientation(SPO) of amphibole 0.2–3 mm in size (Fig. 3a). Thesegrains often show a strong chemical zonality with actinoliticcores and tschermakitic rims (Fig. 4a). Fine-grained (30–60 mm) sub-equant plagioclase (An 25–35 ) forms elongatepolycrystalline aggregates surrounded by laths of amphibole.These new grains develop from large relict clasts(An 40–50 ) (Fig. 4d). The plagioclase in these domainsexhibits features typical for dynamic recrystallization (in thesense of Poirier and Guillopé (1979)) such as undulatoryextinction, the development of sub-grain boundaries, and auniform grain-size distribution. However, a contribution ofmetamorphic nucleation is evidenced by the differentchemical composition of the host and the new grains(Rosenberg and Stünitz, 2003).In the hinge zones of F 3 microfolds, amphibole grainswith irregular boundaries are commonly bent and broken(Fig. 3b). Microfractures associated with domainal undulatoryextinction fragment large grains into smaller elongatesegments. There is no difference in plagioclase microstructurein the non-folded foliation and in the crenulateddomains. It is concluded, therefore, that the main D 3deformation mechanism for amphibole is fracturing (in thesense of Nyman et al. (1992) and Stünitz (1993)) andpassive grain rotation within the plagioclase matrix.3.2. Central part of the massif (the staurolite zone)In the eastern part of the central zone of the massifactinolite is still locally present in the cores of dark greengrains of magnesio-hornblende. Amphibole compositionfollows the pargasitic line (Fig. 4b) within the S 2 foliation.The size of the dark green grains of magnesio-hornblende113

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DTD 5ARTICLE IN PRESSL. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 7Fig. 4. The evolution of amphibole (a)–(c) and plagioclase (d) compositions in the garnet, staurolite and sillimanite zones. In (a)–(c) the compositions situatedbelow the pargasitic line (PRG) in the garnet and straurolite zones are interpreted as relics of previous metamorphic or magmatic stage while those following orsituated above the pargasitic line are thought to be of metamorphic origin (for further discussion see Baratoux (2004)). The numbers next to the symbolscorrespond to the analysed samples. Sample 1 in the sillimanite zone is situated in the pluton aureole. Compositions of end members edenite (ED), pargasite(PRG), tschermakite (TS), actinolite (AC), and tremolite (TR) are indicated.varies from fine-grained (0.01–0.1 mm), through mediumgrained(0.1–1 mm), up to coarse-grained (0.5–3 mm)porphyroblasts. The finer the amphibole grains, the strongertheir SPO and alignment are. The sub-equant, weaklyelongate plagioclase of An 25–35 composition and 20–50 mmin size is either randomly distributed among the amphibolesor forms either elongate polycrystalline aggregates or layersalternating with the amphibole planar fabric (Fig. 3c).Straight boundaries of sub-equant plagioclase grainscommonly meet in triple junctions.In the hinge zones of F 3 folds, large amphiboles aregenerally bent and fractured (Fig. 3d). The main differencewith respect to the previous zone is an important grain sizereduction of amphibole down to 0.01–0.05 mm as a result offracturing. New grains attain lower aspect ratios andbecome mixed with plagioclase of the same grain size.There is no difference in chemical composition of either theplagioclase or amphibole in the hinge zones compared withthe unfolded S 2 fabric.3.3. Western area of the massif (the sillimanite zone)In the western area of the massif, i.e. in the sillimanitezone, there is a major change in the microstructuralFig. 3. Photomicrographs of representative microstructures for each zone. (a) Large amphibole grains parallel to the S 2 foliation in the garnet zone aresurrounded by fine-grained recrystallized plagioclase. (b) In the hinge area, amphiboles included within a sub-equant plagioclase matrix grains are bent andbroken as a result of micro-folding. (c) Amphibole and plagioclase in the limb area of the staurolite zone tend to form alternating aggregates elongated parallelto the foliation. (d) Microcrenulation leads to strong grain size reduction via brittle deformation of amphibole resulting in the mixing of amphibole andplagioclase in the hinge areas. (e) Plagioclase and amphibole in the sillimanite zone show straight equilibrated boundaries. Both minerals are arranged parallelto the S 2 foliation. (f) Amphibole and plagioclase grains of the hinge area deformed by microcrenulation achieve a high aspect ratio. They reorient sub-parallelto the axial plane without any internal deformation. (g) Foam-like texture with straight equilibrated grain boundaries meeting at triple junctions of 1208 as wellas a high grain size in both minerals is typical of the metabasites of the pluton aureole. Amphibole and plagioclase tend to form separate monomineral layers.(h) Amphibole and plagioclase texture in the hinge zone of microfolds is similar to that in the limbs.115

DTD 5ARTICLE IN PRESS8L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24character of the metabasites compared with those in thegarnet and staurolite zones. Brownish green ferroanpargasitic grains of hornblende with aspect ratios of 2–4are arranged parallel to the S 2 fabric (Fig. 4c). They displaymutual equilibrated straight grain boundaries typical ofhigh-grade amphibolite textures (Brodie and Rutter, 1985).Unlike the staurolite zone where the plagioclase (An 30–60 )tends to form layers, we observe isolated grains andaggregates of plagioclase surrounded by elongate and welloriented crystals of amphibole (Fig. 3e). Chemical zonalityand a large span of plagioclase compositions evolvingtowards anorthite document prograde metamorphic growthassociated with dynamic recrystallization (Yund and Tullis,1991) (Fig. 4d).The hinge zones of micro-crenulations are very narrowand the amphibole grains tend to reorient parallel with thefold axial planes without any bending and fracturing (Fig.3f). The hornblende grains show straight boundaries, localdecussate structures, and a similar aspect ratio and size tothose within limb zones. Plagioclase reaches slightly higheraspect ratios and grain size compared with the limbs. Allthese features together with the absence of any compositionalzoning of the pargasitic hornblende (Fig. 4c)indicate grain growth related to prograde metamorphism(Vernon, 1976).3.4. Western area of the massif (the sillimanite zone of thepluton aureole)Unlike the zones described above, brown pargasitichornblende rich in Ti within the sillimanite zone of thepluton aureole, exhibits straight or slightly concaveequilibrated boundaries and fairly low aspect ratios (1.5–2). It is difficult if not impossible to distinguish in thinsection the S 2 and S 3 fabrics. In most of the studied samplesthe pargasitic hornblende is associated with plagioclase(An 30–60 ) of similar aspect ratios and size and arranged intoa regular ’static’ foam-like structure (Fig. 3g). The latterexhibits straight growth-related twins. Plagioclase-rich andamphibole-rich compositional bands are locally developed.Both minerals attain the same size ranging between 0.05 and0.5 mm. 1208 triple point junctions are formed byamphibole–amphibole, plagioclase–plagioclase, and evenamphibole–plagioclase grain boundaries. All of the featuresdescribed above are consistent with re-equilibration underhigh temperature conditions corresponding to the M 3 HT/LPmetamorphic overprint. The hinges of F 3 microfolds areextremely rare but if present they show very similarmicrostructural relations of hornblende and plagioclase tothat of the main fabric (Fig. 3h). A characteristic feature isthe growth of large elongate hornblende crystals parallel tothe axial plane of the F 3 microfolds. The texture of theserocks bears a remarkable resemblance to the upperamphibolite to granulite facies example of Brodie andRutter (1985, p. 155).4. Fold shape analysis4.1. Methods of quantitative analysisFor the purpose of the fold analysis, 3–6 photographs ofrepresentative fold types from each metamorphic zone(garnet, staurolite and sillimanite with a low degree of HToverprint) were selected (Fig. 5). The fold analysis could notbe carried out on the rocks from the pluton aureole becauseof the scarcity of macroscopic folds. The fold shapes,redrawn from photographs, were transformed by parallelprojection onto a plane perpendicular to the fold axis. Twoquantitative methods have been applied to quantify the foldshape: the method of Lisle (1997) based on Ramsay’s(1967) classification and the harmonic fold shape analysis ofHudleston (1973). The principles of these methods are givenin Appendix A.The method of Lisle is based on the polar projection ofthe normalized thickness of a folded layer. Each fold can becharacterized by a single number (the index F), whichexpresses the amount of flattening within each fold limb.Class 1 folds are characterized by positive F values (0 to N)and these give a measure of the amount of homogeneousflattening perpendicular to the axial plane required togenerate this fold shape from a parallel fold (class 1C forFO1, class 1B for FZ1 and class 1A for 0!F!1). Class 3folds are folds with strong thinning of the limbs (attenuatedfolds), typically developed in incompetent layers.These folds are characterized by negative F values (0 toKN). Lisle (1997) divided the field of class 3 folds intothree subfields with fold shape 3B marking the boundarybetween the newly defined 3A and 3C classes. Class 3Bfolds are defined as a ‘pure’ class 3 fold geometry fromwhich the other types (3A and 3C folds) develop by thesuperposition of flattening strains. Class 3A folds aregenerated from 3B by flattening in the direction normal tothe axial plane and have F values ranging from K1toKN.Class 3C folds are the result of flattening parallel to the axialplane and are relatively rare in the nature.Hudleston’s (1973) fold classification, based on theFourier harmonic analysis of folds, exploits the fact that thegeometry of a quarter wavelength is sufficient to characterizethat of the whole fold. The shape is approximated interms of its harmonics and it is found that the first two oddharmonics (coefficients b 1 and b 3 ) are sufficient toadequately describe the fold’s profile shape. b 1 expressesthe amplitude of the fold profile, b 3 is a measure of the‘angularity’ or ‘sharpness’ of the fold’s hinge. For sinewaves, b 3 Z0, for box-like folds b 3 O0 and for chevron-likefolds b 3 !0. The ratios of b 3 /b 1 describe a continuous seriesof shapes between the chevron and box-end members. Themethod for calculating b 1 and b 3 values is given inAppendix A.In the present analysis of fold shape, these two methodshave been combined by plotting b 1 values (a measure of theactive amplification of the fold) against F values (a measure116

DTD 5ARTICLE IN PRESSL. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 9Fig. 5. Field photographs of typical fold shapes: (a) garnet zone, (b) staurolite zone, (c) sillimanite zone and (d) sillimanite zone of the pluton aureole. Chevronfolds are typical for the garnet zone (a), flattened folds with high amplitude are characteristic of the staurolite zone (b), and similar folds with rather lowamplitudes are present in the sillimanite zone (c). Steep foliation with relic F 3 folds characterize the sillimanite zone in the pluton aureole (d).of the post buckle flattening). This graph shows therelationship between these two different expressions ofshortening. In order to illustrate the geometric implicationsof this graph we have plotted the fold patterns of the modelmultilayer sequences given by Ramsay and Huber (1987,pp. 414–415) (Fig. 6). This shows clearly the effect of theviscosity ratio (m 1 /m 2 ), layer thickness ratio (nZd 1 /d 2 ) andthe amount of superimposed strain on the folds positionon the graph. Folded incompetent layers plot in the left halfof the diagram and the slope of the data points shows thestrong negative correlation between the b 1 and F parameters.Folded competent layers plot in the right half of thediagram and the steeper slope of the data point trend shows aless pronounced positive correlation between the b 1 and Fparameters (Fig. 6).We have used these plots as standards with which wecompared the plots representing the natural folds from thestudy area. In addition to the qualitative estimate of theviscosity contrast and the proportion of incompetent tocompetent fraction, we can deduce semi-quantitatively theactive amplification of the buckle fold and the amount ofpost-buckle shortening. The histograms of the F values andb 3 /b 1 ratios for each of the folds studied are given in Fig. 7.4.2. Quantitative analysis of fold styles in the garnet zoneThe post peak metamorphic folds in the eastern part ofthe massif, i.e. in the garnet zone, display only a smallvariation in shape. The post buckle flattening is small, asdocumented by relatively low positive values of F (0.8–3)(Fig. 7a), which approximates to a parallel or weaklyflattened parallel fold. Very few folds in this domain belongto class 3A, i.e. folds that exhibit relatively small negativevalues of F (Fig. 7a). The dominance of class 1B (i.e.parallel) folds is important as it documents that activebuckling played a major role during the fold amplificationand that post-buckle flattening was subordinate. The shapesacquired from the harmonic Fourier analysis lie betweenchevron and sine wave, as can be seen from the histogram ofb 3 /b 1 (Fig. 5a), which has an average value of b 3 /b 1 ZK0.05. The absolute values of both F and b 1 are low(F max Z3; b 1max Z3), suggesting a low amount of deformationof the analysed rocks. A comparison of the measuredfold shapes plotted on the b 1 vs. F diagram with the modelmultilayer sequences (Fig. 6) indicates that the fieldexamples show a close resemblance to model folds (type5, i.e. folds characterized by a high viscosity ratio (m 1 /m 2 )117

DTD 5ARTICLE IN PRESS10L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24Fig. 6. The position of representative multilayer sequences (modified after Ramsay and Huber (1987, pp. 415–416)) in the diagram b 1 vs. F. The relativelycompetent layers are dark in colour. Corresponding symbols are given to the left of the schematic drawings; open symbols represent the more deformed stage.The boxed area (F from 4 to K4 and b 1 from 0 to 3) has been enlarged. The properties of the multilayer sequences are (following Ramsay and Huber, 1987): (1)m 1 /m 2 is low, n (the number of layers) is high, _A (fold amplification rate) is low, _e is high; (2) m 1 /m 2 is low, n is moderate, _A is moderate, _e is moderate; (3) m 1 /m 2is low, n is low, no characteristic initial wavelength is established; (4) m 1 /m 2 is high, n is high, _A is high, _e is low; (5) m 1 /m 2 is high, n is moderate, _A is high, _e islow. (6) m 1 /m 2 is high, n is low, _A is high, _e is low. Multilayer sequence (3) is missing in our diagram because no folding occurs and the deformation ispredominantly layer parallel shortening. (a) and (b) in the text refer to the less and more deformed stages, respectively.and moderate n (d 1 /d 2 ) values). This fold assemblage hasgeometry close to that of chevron folds and is characterizedby high fold amplification and a low degree of post-buckleflattening.4.3. Quantitative analysis of fold styles at the staurolite zoneThe histogram of F indexes for the post peak metamorphicfolds in the staurolite zone (Fig. 2) shows that thefolds of class 1B (FZ1) are no longer present in the centralpart of the massif (Fig. 7b). Class 1C folds are the mostcommon with F values ranging between 1 and 5 with anaverage value around 4. There are a number of foldsbelonging to class 3A, with F values ranging from K10 toK1 with an average of around FZK4. These F valuesindicate that a high amount of post-buckle flattening occursin both the competent and incompetent layers. Three peakscan be distinguished in the histogram of the b 3 /b 1 ratio. Onemaximum is situated close to the chevron shape, a second isclose to the sine wave and the third one lies between theshapes of a parabola and a semi-ellipse (Fig. 7b). However,the majority of the folds are situated between the chevronand sine wave shapes. This shape diversity indicates thatfold geometries characteristic of both competent andincompetent materials occur. In the context of the mineralfabric in which these folds develop, this indicates that thereare compositional and modal differences implying competencecontrasts between layers. The diagram b 1 vs. F alsodemonstrates that the amphibolites in this zone havesuffered an important amount of deformation, which is118

DTD 5ARTICLE IN PRESSL. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 11Fig. 7. Graphs of the fold analysis b 1 vs. F with the inset histograms of F and b 3 /b 1 coefficients: (a) garnet zone; (b) staurolite zone; (c) sillimanite zone. Typicalfold shapes (sine wave, parabola etc.) expressed by the ratio of b 3 /b 1 (Hudleston, 1973) are depicted as vertical lines. See text for further discussion.documented by higher values of the b 1 and F coefficients (Fig.7b). This fold assemblage corresponds to that of model 5b(Fig. 6) a multilayer sequence marked by high viscosity ratio(m 1 /m 2 ) and moderate n (d 1 /d 2 ) values. This implies that highfold amplification and an important post-buckle flatteningdeveloped in both competent and incompetent layers. Thesefolds can be regarded as a more deformed equivalent of thefold assemblage observed in the garnet zone (cf. Fig. 7aandbwith Fig. 6—model fold types 5a and 5b).4.4. Quantitative analysis of fold styles in the sillimanite zoneWithin the sillimanite zone (Fig. 7c) most of the folds areof class 1C and have F values that are concentrated between119

DTD 5ARTICLE IN PRESS12L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–241 and 5 with an average value of around 3. Very few class3A folds are present. This indicates the important fact thatthere are relatively few incompetent layers and that a highdegree of post buckle flattening has occurred (as indicatedby the absence of parallel folds). The broad histogram of theb 1 /b 3 ratios, with no unequivocal peaks, reveals that theaverage fold shape approximates to that of a sine curve. Thegraph of b 1 vs. F (Fig. 7c) confirms the low number of class3A folds and the dominance of class 1C. It can be seen fromthis diagram that class 1C folds (FO1) have a low range ofb 1 values (between 1 and 4). In addition, we observe apositive correlation between the F and b 1 parameters. Theimplications of these observations are that although the foldamplification increases with the degree of post-buckleflattening, the maximum value of amplification is relativelylow. The measured fold assemblage may be compared withthe model folds of type 2b in Fig. 6, which indicate a lowviscosity ratio (m 1 /m 2 ) and a moderate value of n (d 1 /d 2 ).Compared with the folds in the staurolite zone where highamplification and low post-buckle flattening dominate, inthe sillimanite zone we observe limited amplification andpronounced post-buckle flattening. This indicates a changefrom folding dominated by active amplification (in thestaurolite zone) to dominantly passive fold amplification inthe sillimanite zone. We interpret this as reflecting animportant change in the rheological properties of the foldedmaterial rather than the result of differing amounts of finitestrain.4.5. Active buckling vs. post buckle flattening in the studiedfoldsThe fold geometries were examined in order to determinethe relative importance of active buckling and post-buckleflattening across the metamorphic zones of the studiedamphibolite unit. It has been shown that in the garnet zonefold amplification was dominated by active buckling withonly a small contribution from post-buckle flattening. In thestaurolite zone, although the folds have also experienced ahigh degree of amplification, this involved both activebuckling and post-buckle flattening. By comparing theseresults with those obtained from the model folds it can beargued that the folds in the staurolite zone are the equivalentof those in the garnet zone (model fold types 5a and 5b) buthave experienced a higher degree of finite strain. In thesillimanite zone the folds show a relatively lowamplification but important post-buckle flattening eventhough field observations suggest that the strainintensity of these folds and those of the staurolitezone are very similar. It is concluded that thedominance of flattening in the folds developed in thesillimanite zone, indicating that the folding wasdominated by passive amplification as opposed to thatwhich occurred in the garnet and staurolite zones.5. Quantitative analysis of rock anisotropyIn order to understand more fully the results of themesoscopic fold analyses discussed above, it is useful tostudy the petrofabrics of the folded units. The main goals ofthis study are to evaluate the mechanical anisotropy offolded systems and to determine the micro-deformationalmechanisms associated with folding.5.1. Methods of quantitative microstructural analysisApproximately 100 thin sections, collected from all threemetamorphic zones, have been studied in an attempt to showthe relationship between the microstructures and thefolding. Quantitative microstructural analysis has beenapplied to eight representative samples collected from thegarnet, staurolite and sillimanite zones and from the graniteaureole. Two thin sections, cut perpendicular to the F 3 foldaxes (YZ sections), were taken from the fold hinge and foldlimb in each zone, respectively. As noted above, macroscopicfolds are only present in the garnet to sillimanitezones. The degree of transposition of the original metamorphicfabric within the contact aureole was so high thatthe macroscopic folds are preserved only in domains withhigh lithological contrast. However, relics of rootlessmicroscopic folds can be seen in massive amphibolites inthin sections and, in this zone, it is these that are analysedfrom a microstructural point of view.A quantitative microstructural analysis of grains andgrain boundaries was carried out on the representativesamples by tracing and digitising the outlines of individualgrains using the ESRI ArcView 3.2 Desktop GIS environment.The map of grain boundaries was generated usingArcView extension Poly (Lexa, 2003). These data havebeen treated by MATLABe PolyLX Toolbox (Lexa, 2003)in which grain boundary and grain SPO were analysed usingthe moments of inertia ellipse fitting and eigen-analysis ofthe bulk orientation tensor techniques (Lexa, 2003).Digitised drawings of representative samples are shown inFig. 8 and the results from the quantitative microstructuralanalyses are presented in Table 1. The grain size of theminerals was calculated in terms of their Ferret diameter andthe resulting grain size distributions were statisticallyevaluated. The grain size statistics are summarized in Fig.9, which shows median values and the quartile difference ofthe Ferret diameters. In addition, a method of determiningthe orientation of grain boundaries and grain shapes,using the eigen-analysis technique, was applied in anattempt to quantify the bulk rock anisotropy. The results ofthe analysis of the bulk SPO of grains vs. their aspect ratio(R) are presented in Fig. 10. Grain boundary preferredorientation (GBPO) is presented using a diagram ofeigenvalue ratios (rZe 1 /e 2 ) and the orientation of thelargest eigenvector of the different types of grain boundaries(Fig. 11).The grain size analysis is a powerful technique for120

DTD 5ARTICLE IN PRESSL. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 13describing the degree of grain coarsening related to theintensity of metamorphism (Kretz, 1994). Moreover, inpolymineralic tectonites the relative grain size distributionfor different minerals may indicate the degree of strain–stress partitioning (Handy, 1990; Schulmann et al., 1996).The SPO and elongation of minerals with a low degree ofcrystallographic anisotropy is generally attributed to theamount of strain in a rock (recrystallized quartz, calcite andfeldspar). However, the degree of elongation of stronglyanisotropic minerals, such as amphiboles, reflects the degreeof metamorphism rather than the degree of strain. Consequentlya typical feature of amphiboles is the decrease ofaxial ratio with increasing metamorphic grade (Brodie andRutter, 1985). Although the degree of GBPO is the factormost affiliated to the microstructural anisotropy, it isnecessary to discuss all of the above-mentioned parametersin order to evaluate the bulk mechanical anisotropy of therock.5.2. Grain size distributionThe grain size is expressed as the Ferret diameter of thegrain section without stereological corrections. The grainsizes of amphibole and plagioclase show slightly differentevolutionary trends (Fig. 9). In the limbs (full symbols),amphibole grain size decreases from 31 to 21 mm from thegarnet to the staurolite zone, respectively. The size thenincreases with increasing metamorphic grade via sillimanitezone reaching 80 mm in the contact aureole of the granite.The grain size distribution of amphiboles in the hinge zones(open symbols) shows a similar trend but the grain sizedifferences between the metamorphic zones are not as highas in the fold limbs. The amphiboles from the fold hinges inthe contact aureole (median value 43 mm) are slightlysmaller than those from fold hinges in the sillimanite zone(median value 46 mm).In the limbs, dynamically recrystallized plagioclase hasthe same grain size in both the garnet and staurolite zones(median value 17 mm). The grain size then increases withincreasing metamorphic grade up to a median value of65 mm in the contact aureole. As with the amphiboles, thegrain size is always smaller in the fold hinges reaching amaximum median value of 35 mm in the contact aureole.Both the average grain size of amphibole and the varianceare slightly higher than those of plagioclase for allmetamorphic zones.In the garnet zone, amphibole shows a larger grain sizeand grain size spread than plagioclase (Fig. 9). Moving tothe staurolite zone the grain size of the amphibolesdecreases, resulting in a decreasing difference in grain sizedistribution between the plagioclase and amphiboles. Themain characteristics of the sillimanite zone are the increasein grain size of both minerals with respect to the previouszones coupled with an increase in difference of grain sizedistributions. Within the contact aureole the grain sizedistributions of coarse-grained plagioclase and amphibolebecome equal.5.3. Aspect ratio and shape preferred orientationThe microstructures of the fold limbs are characterizedby a strong SPO of amphibole grains as well as a highaverage aspect ratio (RZ2.05, 2.17 and 2.03 in the garnet,staurolite, and sillimanite zones, respectively) (Fig. 10).Only in the highest metamorphic grade does the SPOdecrease, together with the aspect ratio (down to RZ1.57).Plagioclase displays the same trend as the amphibole but theabsolute values of their aspect ratios and SPO are noticeablylower (RZ1.46, 1.64 and 1.59 in the garnet, staurolite andsillimanite zones, respectively). As with the amphiboles, theaspect ratio and SPO of the plagioclase decreases in thecontact aureole (RZ1.49, i.e. similar to the R value ofamphibole in this zone).The SPO of amphibole in the hinge zones shows acomplex evolution. In the garnet zone the SPO is still high,but the aspect ratio is rather low (RZ1.81). In the staurolitezone both the SPO and the aspect ratio (RZ1.89) decreasewith respect to the limb zone. In contrast the fold hinges inthe sillimanite zone are associated with rather high SPO andvery high aspect ratio with respect to the previous zone (RZ2.36). Within the contact aureole it is found that the SPO ofamphibole and the aspect ratio (RZ1.60) diminish in thehinge areas of the microfolds. A similar trend can beobserved in the plagioclase (RZ1.48, 1.53, 1.76 and 1.41for the garnet, staurolite, sillimanite zones and the contactaureole, respectively).5.4. Grain boundary preferred orientationA study of the evolution of amphibole–amphibole grainboundary preferred orientation (GBPO) shows two importanttrends. The first is the GBPO in the fold limbs which,apart from the staurolite zone, decreases with increasingmetamorphic grade (rZ2.12, 2.39, 1.76 and 1.23 for thegarnet, staurolite, sillimanite zones and the contact aureole,respectively) (Fig. 11). The second trend shows that in thegarnet and staurolite zones the GBPO in the hinge domainsis at a high angle to that of the axial plane representing S 3 .Itcan be seen that in the garnet and staurolite zones the degreeof GBPO decreases noticeably in the hinge areas of foldscompared with the limb areas (rZ1.47 and 1.43, respectively).In contrast, in the sillimanite zone, the GBPO in thehinge zone becomes parallel or sub-parallel to the axialplane (S 3 ) and shows the same intensity (rZ1.7) to that inthe limbs. Similarly, in the contact aureole, the GBPO in thehinge zones shows parallelism to the axial plane (S 3 ), andhas a similar intensity (rZ1.22) to that on the limbs.The r values for plagioclase–plagioclase boundaries areweak for all the zones (1.06–1.21; see Fig. 11 and Table 1)indicating a very weak or even a lack of GBPO. In contrast,the amphibole–plagioclase GBPO, which is controlled by121

DTD 5 ARTICLE IN PRESS14L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24Fig. 8. Typical microstructures from the four metamorphic zones, redrawn and digitized in the ArcView GIS environment. The rose diagrams represent grain long axis distribution. Full symbols correspond tofold limbs, open symbols to fold hinges. These symbols have the same significance in Figs. 9 and 11. The minerals can be identified as follows: plagioclase—white; amphibole—medium grey; spheneCilmenomagnetite—black; quartz—vertically hatched; calcite—dotted; clinopyroxene—light grey.122

Table 1Statistical values of the quantitative textural analysisZone Grt St SilSil (plutonaureole)Sample 162 163 84 84 10 10 149 4Limb Hinge Limb Hinge Limb Hinge Limb HingeGBPO Pl–pl 1.06 1.10 1.21 1.21 1.25 1.12 1.10 1.22Eigenvalue rZ Amp–amp 2.12 1.47 2.39 1.43 1.76 1.70 1.23 1.22e1/e2Amp–pl 1.45 1.17 1.95 1.26 1.48 1.52 1.33 1.08Eigenvector Pl–pl K25 K74 22 88 K9 K65 28 83orientation (8) a Amp–amp K11 K22 K3 K76 K11 K77 74 K83Amp–pl K16 K2 K4 89 K14 K78 91 K73SPO Pl 1.21 1.06 1.57 1.25 1.42 1.44 1.23 1.16Eigenvalue r Amp 2.39 1.57 3.06 1.44 2.16 1.88 1.44 1.31Eigenvector Pl K16 K38 1 K90 K9 K81 89 K90orientation (8) a Amp K11 K3 K6 K88 K15 K79 79 K86Aspect ratio Pl 1.46 1.48 1.64 1.53 1.59 1.76 1.49 1.41(median)Amp 2.05 1.81 2.17 1.89 2.03 2.36 1.57 1.60Grain size—FerretdiameterMedian (mm) Pl 17 22 17 15 39 28 65 35Amp 31 33 21 20 64 46 80 43Q1 (mm) Pl 11 14 10 10 27 20 36 21Amp 16 18 13 13 38 29 45 29Q 3 (mm) Pl 25 34 25 21 52 39 114 53Amp 75 59 35 32 104 79 130 65Q3KQ1 (mm) Pl 14 20 15 12 25 19 79 33Amp 59 41 22 19 66 49 85 37a Positive values are oriented anticlockwise with respect to the horizontal.L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 15DTD 5 ARTICLE IN PRESS123

DTD 5ARTICLE IN PRESS16L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24Fig. 9. Plot of amphibole and plagioclase median of Ferret diameter and variances for all the metamorphic zones studied. Full symbols correspond to fold limbsand open symbols represent hinge areas. Samples from the pluton aureole are assigned as Sil*.the difference between the grain size of the large elongategrains of amphibole and small grains of plagioclase as wellas by the SPO of the amphibole, is higher. The highestGBPO for these boundaries is developed in the limb zonesof the folds in the staurolite zone (rZ1.95). The degree ofGBPO is low in the limb zones of folds in the garnet andsillimanite zones (rZ1.45 and 1.48, respectively) andfurther decreases to rZ1.33 in the contact aureole. In thehinge regions the GBPO of plagioclase–amphibole boundariesis very low in the garnet, staurolite zones and contactaureole (rZ1.17, 1.26 and 1.08, respectively; Fig. 11). It isonly in the hinge zones of the folds in the sillimanite zonethat the GBPO of this boundary exceeds 1.5 (rZ1.52).5.5. A comparison of the microstructures in the hinge andlimb areas for folds from the different metamorphic gradesThe analysis of the limb areas of the folds in the garnetzone reveals a large grain size and aspect ratio differencebetween the amphiboles (elongate and large) and plagioclase(sub-equant and small) (Figs. 9 and 10). This workalso shows a relatively small SPO and GBPO for plagioclasecompared with strong SPO and GBPO for amphibole. Thismicrostructure represents a network of large amphibolecrystals surrounding pockets of fine-grained plagioclase. Inthe hinge regions of folds in the garnet zone the amphibolesare oriented at high angles to the axial plane. Compared withthe limb there is a decrease in the aspect ratio, grain size,SPO and GBPO of both minerals.In the limbs of the folds in the staurolite zone both theamphiboles and the plagioclase display a very intense SPO,high aspect ratio and a strong GBPO coupled with adecreasing difference in grain size between the two mineralscompared with the fold limbs in the garnet zone (Figs. 9–11). This reflects the fact that both minerals forminterconnecting aggregates. The hinge zone microstructure,like that in the garnet zone, is marked by rotation of thefabric into an orientation at a high angle to the fold axialplane and, compared with the limbs, the mineral aspectratio, grain size, SPO, and GBPO are all reduced.The sillimanite zone and the contact aureole both showan increase in grain size for both minerals compared withthe garnet and staurolite zones (Figs. 9–11). However, themain difference is the parallelism of the mineral preferredorientation in the hinge zones with the fold axial plane,which is expressed by both the SPO and the GBPO (Figs. 8and 11). This is connected with an increase in aspect ratiofor both plagioclase and amphibole in the hinge comparedwith the limb. The other important feature is the similarGBPO for like–like and unlike boundaries for both the hingeand limb zones. The only difference between the sillimanitezone and contact aureole is a weakening of the SPO andGBPO in the latter where the aspect ratio of both mineralsapproaches 1. The distribution of amphibole and plagioclaseaggregates in the limbs is rather random in the sillimanitezone and tends to be layered in the contact aureole.In summary, we note that the rocks of the study area fallinto two distinct groups. Those in the garnet and staurolitezones, which show mineral preferred orientation in thehinge zone at a high angle to the fold axial plane and to the124

DTD 5ARTICLE IN PRESSL. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 17Fig. 10. The plot of grain shape preferred orientation (SPO) of amphibole and plagioclase displaying the weighted ratio of the eigenvalues of inertia. The resultsare summarized in a boxplot-type diagram of aspect ratios vs. eigenvalue ratio of bulk matrix of inertia of the individual minerals. The individual boxes showthe median, first and third quartiles of the aspect ratio. The ‘whiskers’ represent statistical estimates of the data range. Outliers are not plotted. Boxes of thelimbs are shaded while those of hinge areas are white. The number of the analysed sample is indicated for each zone. Samples from the pluton aureole areassigned as Sil*.SPO of amphibole and plagioclase on the limbs, and those inthe sillimanite zone and contact aureole, which showmineral preferred orientation in the hinge zone sub-parallelto the axial plane.6. DiscussionIn this paper we have attempted to use the theory ofbuckling of multilayers and the theory of buckling ofmechanically anisotropic materials to explain the foldingmechanisms of some natural folds, which show lateralvariations in their shape. The folds are developed in a rockunit of constant mineral composition but which exhibits alateral variation in rock microstructure related to metamorphicgrade. Systematic variations in fold shapes acrossthe whole profile suggest that the original variations in rockprotolith are less important for folding mechanisms than themetamorphic recrystallization associated with developmentof mineral fabrics prior to the folding event.6.1. Competence contrast vs. mechanical anisotropyThe question arises as to whether the differences betweenthe fold styles in the garnet and staurolite zone, and folds inthe sillimanite zone, are the result of differences in themechanical anisotropy of the mineral fabric or of rheologicalvariations between adjacent layers. As is discussed inthe following section, an inspection of Figs. 6 and 7 can helpanswer this.The buckling behaviour of a multilayer is controlled bythe ratio of the competencies, e.g. the viscosities (m 1 /m 2 )ofthe competent and incompetent layers and their relativethicknesses (e.g. Ramberg, 1963). The most important125

DTD 5ARTICLE IN PRESS18L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24Fig. 11. Plot of grain boundary preferred orientation (GBPO) of amphibole–amphibole, amphibole–plagioclase and plagioclase–plagioclase boundariesdisplaying weighted ratio of eigenvalues of inertia. The numbers represent the orientation of the eigenvector V 1 of GBPO with respect to the horizontal.Positive values indicate anticlockwise deviation. Samples from the pluton aureole are labelled Sil*.factor controlling the mode and style of folding of amineral fabric is the mechanical anisotropy of thematerial and this is defined by the ratio of two moduli,one a measure of the resistance to layer or fabricparallel compression (M) and the other to the resistanceto shear in the same direction (L).It can be shown that there is a direct link betweenthe anisotropy (M/L) and the competence contrast of abilaminate (m 1 /m 2 )(seePrice and Cosgrove, 1990) andthis allows one to use both theories to study foldingbehaviour. By analysing the geometries of the studiedfolds it is possible to determine which of the twotheoretical approaches is more appropriate. As shownearlier, the low F values of the folds from the garnet(low b 1 ) and staurolite zones (high b 1 ) are similar tothose in the range type 5a and 5b to type 4b (Fig. 6).The trend of the arrows in Fig. 7b (which indicates theratio of active buckling amplification and fold flatteningin the staurolite zone) is comparable with the trendsindicated in Fig. 6 for a multilayer (type 5b) with ahigh competence contrast between adjacent layers. Theconstruction of the dip isogons for the folds from thiszone shows alternations of layers with differentgeometries (classes 1B and 1C alternating with class3) and different curvature profiles of the hinge regions(Fig. 12), a pattern compatible with the folds of type5b. Using the classical theory of Ramberg (1963), thispattern can be interpreted in terms of folding of amultilayer sequence marked by alternation of layerswith a high competence contrast.However, as was shown above, the garnet zoneamphibolites are more or less compositionally homogeneousand do not show distinct compositional layering.Inspection of Fig. 7a shows that the folds in the garnet zonecould be considered to represent the early stages of activeand passive amplification observed in the staurolite zone.However, there is a marked difference in the density of thedata relating to the ‘competent’ and ‘incompetent’ membersof the multilayer. When the pattern of dip isogons for arepresentative fold profile for these folds is examined (Fig.12), it can be seen that there is an important increase of foldflattening component resulting in a convergence of shapes(and therefore dip isogon patterns) of adjacent layers.A comparison of the natural data from folds in thesillimanite zone (Fig. 7c) with the model folds in Fig. 6shows that the best fit is with type 2b folds. Thisgeometry is traditionally interpreted as indicating abilaminate with a low competence contrast. However,we note the strong asymmetry in the density of pointsindicating a lack of incompetent ‘layers’. This isconfirmed by inspection of the dip isogon pattern ofrepresentative folds from this zone (Fig. 12). Althoughthe pattern shows alternations of fold shapes of class 1Cwith those of class 3, it is obtained by analysing‘layers’ that are composed of several units. The absenceof alternations of clearly defined individual layerscombined with the high degree of flattening andrelatively low fold amplification indicate that thissystem approximates more closely to the model analysedby Biot than to a bilaminate.126

DTD 5ARTICLE IN PRESSL. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 19Fig. 12. Selected fold assemblages from the garnet, staurolite and sillimanite zones, respectively. (a) Dip isogon patterns. (b) Graphs showing the changes in F(degree of fold flattening) across successive folded layers for each fold assemblage. (c) b 1 vs. F plots constructed for the presented folds showing the relativeimportance of fold flattening and active fold amplification.6.2. Interpretation of the deformation micro-mechanismsassociated with foldingExperimental work on amphibole and plagioclase as wellas field observations have shown that the former mineral isstronger than the latter under the complete range ofhomologous temperatures (see Brodie and Rutter (1985)for a review). Observations of the microstructures in thehinge zone show that the micro-folding in the garnet zonehas been achieved by the bending and fracturing of thestrong amphiboles whilst the relatively weak plagioclaserecrystallizes and accommodates space modificationsassociated with the rigid body rotation of the amphibolecrystals in a weak matrix by mechanism modelled byArbaret et al. (2001), for example (Table 2). However, nofracturing or bending of amphiboles is observed in limbzones, which implies that the fabric on the limbs stillrepresents the original, i.e. pre-folding, fabric. Folding ofamphiboles in the garnet zone is thus similar to thatdescribed from low-grade metasediments and which leadsto the occurrence of metamorphic differentiation (Cosgrove,1976). In this work, Biot’s (1961) theory was used toaccount for the buckling of a greenschist facies mineralfabric made up of a rigid stress supporting framework ofmica containing interstitial weak quartz. In the hinge theframework of mica protects the quartz from the buckling127

DTD 5ARTICLE IN PRESS20L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24stress but on the limbs it does not. The resulting stressgradients cause the migration of quartz from the limb to thehinge and mica from the hinge to the limb.Microstructural examination of the rocks in the presentstudy area shows an analogous situation in the amphiboleplagioclase rock. The resulting stress gradients establishedin both the amphibole and the plagioclase cause a migrationof plagioclase to the hinge and amphibole to the limbs,leading to metamorphic differentiation.We suggest that this process operated in the hinge zonesof some macrofolds leading to their thickening. Because ofstrain compatibility, this hinge thickening develops inalternating layers, thus producing patterns similar to thefolding of bilaminate (Fig. 12).In the hinge zones of folds in the staurolite zone aconspicuous reduction in the grain size of the amphiboleswith respect to the limb zones can be observed. This islocally connected with the reorientation of amphibolefragments, and we interpret this microstructure as beingthe result of the fracturing and rigid body rotation ofamphiboles as described by Nyman et al. (1992). Thedynamically recrystallized plagioclase in the hinge zonesalso shows a decrease in grain size with respect to the limb.In the hinges, both minerals have similar grain size, lowaspect ratio, very weak SPO and GBPO and exhibit animportant degree of mixing. We suggest that these featuresmay stimulate a change in deformation mechanism in thesedomains and that in the highly attenuated folds thefracturing of amphiboles and the dynamic recrystallizationof plagioclase switch to a type of granular flow facilitatingthe development of high strain intensities in these areas(Table 2). The changeover from dislocation creep dominatedflow to granular flows connected with a reduction ofgrain size and the mechanical mixing of minerals has beendescribed by several authors in greenschist facies metabasitemylonites (e.g. Stünitz, 1993; Berger and Stünitz,1996). As with the folds in the garnet zone we argue that thefabric on the limbs of the folds in the staurolite zonerepresents the original, pre-fold fabric.In contrast to the garnet and staurolite zones no bendingor fracturing of the amphibole lattice was noted in the hingezones of folds in the sillimanite zone. Instead the fold hingeareas in the sillimanite zone display crossover growths ofamphiboles and straight, well-equilibrated grain boundariesof all minerals. In addition, the amphibole grains show nosigns of internal deformation and the grain size is alwayshigher than that observed in the garnet and staurolite zones.All these criteria indicate that these grains developed by themechanism of nucleation and possibly syn-deformationalgrowth (Vernon, 1976; Rosenberg and Stünitz, 2003). Theplagioclase also shows well-equilibrated grain boundariesmeeting at triple point junctions and an almost uniformdistribution in the rock, features consistent with the highgradecrystallization of amphiboles and plagioclase (Brodieand Rutter, 1985). Comparison of the mineral microstructuresof the hinge zones and the limbs shows that the hingezones display higher aspect ratios, smaller grain sizes,similar SPO and like–like and unlike GBPOs (Table 2). Thisimplies that, during the development of the folds, recrystallizationoccurred simultaneously in both the hinge zonesand the limbs, producing an increase in elongation of theoriginal crystals in the hinge zone. This is because the grainsin both the limb regions and the hinge regions are parallel tothe axial plane, i.e. are perpendicular to the largest principalcompressive stress axis. This resulted in an increase inaspect ratio by the process of heterogeneous dissolutionaccompanied by recrystallization and grain growth, themechanism of schistosity transposition well known in highgrade schists and described by a number of authors (seePasschier and Trouw (1996, fig. 4.17) for a review). Unlikethe folds in the garnet and staurolite zones where weconsider that the fabric on the fold limbs represents the prefoldingmicrostructure, in the sillimanite zone it is modifiedby dissolution and growth process.In the contact aureole a similar type of microstructuralpattern to the sillimanite zone occurs but it is marked by acomplete loss of SPO and GBPO coupled with an importantgrain size increase (Table 2). These features are consistentwith there having been an important contribution from highgrade “static” recrystallization in both the hinge and limbdomains. However, we emphasise that one can still observesmaller grain sizes and aspect ratios for both amphibole andplagioclase in the hinge zones compared with the limbs.7. ConclusionsA summary of the fold analysis and the microstructuralstudies supporting the fold mechanics models is given inTable 2.The garnet grade region. The quantitative microstructuralanalysis of the limb zones in the folds in the garnetzone reveals the possible existence of a stress-supportingnetwork. This type of structure implies a relativelyhomogeneous stress distribution in the rock (Jordan,1988), which is controlled by the rheologically resistantamphiboles, with the plagioclase representing only weakpockets that deform to accommodate the deformationimposed by the strong amphibole. This is supported by thelarge grain size difference between the recrystallizedplagioclase and amphibole crystals, which can be interpretedin terms of a stress-supporting network with a highrheological contrast between the plagioclase and amphibole(Handy, 1990). In addition, the microstructural analysisshows that the amphibole stress-supporting frameworkcollapses in the hinge zones by brittle failure. This brittledeformation, which represents an extreme example of strainlocalization, tends to produce sharp-hinged chevron folds ona grain scale (Fig. 3a and b).The observed folding mechanism also gives rise tochevron folds on a macroscopic scale indicating that themechanical anisotropy of the rock was high. To interpret the128

Table 2Summary of parameters derived from fold shape and microstructural analyses in all zones. Used symbols are: A, Amp—amphibole; P, Pl—plagioclase; S—grain size; R—aspect ratio; SPO—shape preferredorientation; GBPO—grain boundary preferred orientationFoldingmechanismsDegree ofmechanicalanisotropyQuantitavivemicrostructureparametersMicrostructureDeformationmechanismsGarnetzoneStaurolitezoneSillimanitezone andpluton aureoleActive amplification Active amplification and post-buckle flattening Passive amplificationLow to medium b1. Low positive F Medium to high b1. High F Low to medium b1. High positive FHigh mechanical anisotropy Bilaminate Low mechanical anisotropyS R SPO GBPO S R SPO GBPO S R SPO GBPOLimb AOP AOP AOP AOP AZP ARP ARP ARP AZP AZP AZP AZPHinge AOP AOP AOP AOP AZP ARP ARP AOP AZP AZP AZP AZPAmphibole supported LBF (large elongate and interconnectedIWL microstructure with low-viscosity contrast (alter-Amphibole supported LBF structure with low-viscosityAmp grains surrounding pockets of finenationof Pl-rich and Amp-rich aggregates)contrast (mixture of Amp and Pl of equal size and aspectgrained Pl)ratio)Hinge Limb Hinge Limb Hinge and limbAmp fracturing Brittle reactivation of S 2fabricAmp fracturing and granularflow of Amp–PlmatrixDynamic recrystallizationof Pl and ductile flow inPl-rich layersNucleation and syndeformational growthL. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 21DTD 5 ARTICLE IN PRESS129

DTD 5ARTICLE IN PRESS22L. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24isogon patterns (Fig. 12) we propose an explanation ofalternations of hinge zones thickened by this mechanism ofmicrofolding with those without any thickening and distinctmicrostructural changes.The staurolite grade region. In the limbs of foldsgenerated in the staurolite zone, the pre-folding fabric ismarked by fairly well developed, alternating elongateaggregates of plagioclase and amphibole. The intensity ofmineral preferred orientation is very strong, as documentedby the high aspect ratio and strong GBPO and SPO values.Inspection of the fabric in thin sections (Fig. 3c and d)suggests that this structure approximates to the so-called‘interconnected weak layer structure’ (IWL of Handy(1990)) characterized by an alternation of relatively strongamphibole rich domains and relatively weak domains rich inplagioclase. Because the alternating domains differ only insmall modal differences in amphiboles and plagioclase thisstructure represents an IWL structure with a low viscositycontrast as defined by Handy (1994). Such a systempossesses the geometrical characteristics of a bilaminatewith diffuse boundaries between the layers. In the hingezones the microfabric shows areas of distributed deformation.During fold development these zones of distributeddeformation (granular flow in the hinge zones) wouldcontribute to strain softening in this region leading tocontinuous amplification of the folds. We note that even ifthe deformation of the amphiboles is by brittle failure, itresults in distributed ductile flow in the highly deformedhinges. Unlike the folds in the garnet zone where probablyno slip on the limbs occurred, in the staurolite zone the slipis distributed through the relatively weak plagioclase richzones increasing the tendency for active amplification of thefold. In addition, on the limbs, because of the presence ofrelatively weak ‘layers’, fold amplification can be furtherassisted by flattening. This is well documented by the dipisogons pattern and the b 1 vs. F graph of Fig. 12.The sillimanite grade zone. In the sillimanite zone theamphibole and plagioclase show a relatively high aspectratio connected with a low degree of GBPO of like–like andunlike boundaries. In this zone, unlike the staurolite zonewhere the plagioclase was interconnected, isolated elongategrains or aggregates of plagioclase surrounded by highlyelongate and well-oriented crystals of amphibole occur(Figs. 3e and f and 8). This structure and the flattening ofboth minerals can be interpreted in terms of a stresssupportingnetwork with a low viscosity contrast betweenweaker plagioclase and stronger amphibole (Handy, 1990).Thus, in the sillimanite zone, it was the strength of ‘weaker’plagioclase that dominated the rheological properties of thesystem, which therefore acted as a relatively weak,homogeneous material.In contrast to the garnet zone, where buckling wascontrolled by localized microfolding and to the staurolitezone where it was controlled by ductile shearing alongweak, plagioclase rich zones, no such zones of weakness,which facilitate the active amplification of the folds, areobserved in the sillimanite zone. Instead the fold shapeanalysis shows an importance of post-buckle flattening overactive fold amplification. The micro structural analysis ofboth the limb and hinge areas shows features consistent withhomogeneous flattening (i.e. higher aspect ratio and smallergrain size in the hinge domains than in the limbs) anobservation entirely consistent with passive amplification ofa material with a low mechanical anisotropy.The contact metamorphic aureole. The deformation inthe contact aureole (Figs. 3g and h and 8) is an extremeexample of flattening dominated deformation as shown bydifferences in grain size and grain shapes in the hinge andlimb areas. The lack of macroscopic folds in theamphibolites within the aureole is taken as further evidencethat the amplification was almost entirely passive.AcknowledgementsThe project was funded by grants of Czech NationalGrant Agency No. 42-201-204 to K.S. and 42-201-318 to P.Štípská, by Czech Geological Service assignment No. 6327to P. Mixa, and by a Ph.D. financial support attributed by theFrench Government to L.B. We thank R. Vernon and D.Grujic for constructive reviews and J. Hippertt for editorialhelp.Appendix AA polar graph aimed to represent the variation oforthogonal thickness t around folded layers was first utilizedby Lisle (1997). A fold can be represented by a series ofpoints with polar coordinates (1/t 0 , a), where 1/t 0 is thereciprocal normalized thickness (t 0 Zt/t h , where t h is theextreme value of t, which is generally situated in the foldhinge), and a is the orientation of the layer tangent. Usingthis technique, Ramsay’s (1967) fold types give rise tovarious conic sections (ellipses, hyperbolas) in the polargraph with horizontal semiaxis equal to unity (see Lisle,1997).Flattening index F, or axial ratio of strain ellipse, expressthe amount of post-buckle flattening superimposed on theparallel fold. We used a numerically stable direct leastsquaremethod (Halír` and Flusser, 1998) to fit either ellipsesor hyperbolas onto points in a polar graph of normalizedthickness. This technique of evaluation of the flatteningindex F poses a robust estimate and is preferred in this work.The method of analysing fold shapes in terms of theharmonic coefficients of a Fourier series was originallydevised by Stabler (1968) and subsequently elaborated byHudleston (1973).The most basic and suitable segment of a folded surfacefor analysis is a ‘quarter-wavelength’ unit between adjacenthinge and inflexion points. Such a choice of unit leads to aharmonic series consisting only of the odd terms of a sine130

DTD 5ARTICLE IN PRESSL. Baratoux et al. / Journal of Structural Geology xx (xxxx) 1–24 23series and the fold profile is thus approximated by theequation: f ðxÞ Z XN ð2n K1Þpxb 2nK1 sin2LnZ1The first few harmonic coefficients are sensitiveparameters of fold shape and most information about thefold shape can be gained from the first two coefficients, b 1and b 3 .To obtain these coefficients, a ‘quarter-wavelength’ unitof length L of the fold is divided into equal sectors on a line,which is perpendicular to the axial surface and passesthrough the inflection point. Pairs of f(x n ) and x n coordinatesare measured at each of these points. A system of linearequations is assembled from pairs {x n , f(x n )} and solved forunknown coefficients, so2 sin px 1sin 3px 1/ sin ð2n K1Þpx 312L 2L2L2 3sin px 2sin 3px 2/ sin ð2n K1Þpx b 122L 2L2Lb 3« « 1 «6«74 56 sin px nsin3px nsinð2n K1Þpx 745nb 2nK122L3 2L2Lf ðx 1 Þf ðx 2 ÞZ64 «75f ðx n ÞReferencesArbaret, L., Mancktelow, N.S., Burg, J.-P., 2001. 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J. metamorphic Geol., 2008, 26, 273–297 doi:10.1111/j.1525-1314.2007.00755.xVertical extrusion and horizontal channel flow of orogenic lowercrust: key exhumation mechanisms in large hot orogens?K. SCHULMANN, 1 O. LEXA, 1,3 P. ŠTÍPSKÁ, 1 M. RACEK, 2 L. TAJČMANOVÁ, 2,4 J. KONOPÁSEK, 2,3J.-B. EDEL, 1 A. PESCHLER 1 AND J. LEHMANN 11 Université Louis Pasteur, EOST, UMR 7516 - 7517, 1 Rue Blessig, Strasbourg 67 084, France2 Czech Geological Survey, Klárov 3, 118 21 Praha 1, Czech Republic3 Institute of Petrology and Structural geology, Charles University, Albertov 6, 128 43, Prague, Czech Republic4 Dipartimento di Mineralogia e Petrologia, Università di Padova, Corso Garibaldi 37, I-35 137 Padova, ItalyABSTRACTA large database of structural, geochronological and petrological data combined with a Bougueranomaly map is used to develop a two-stage exhumation model of deep-seated rocks in the eastern sectorof the Variscan belt. An early sub-vertical fabric developed in the orogenic lower and middle crustduring intracrustal folding followed by the vertical extrusion of the lower crustal rocks. These eventswere responsible for exhumation of the orogenic lower crust from depths equivalent to 18)20 kbar todepths equivalent to 8)10 kbar, and for coeval burial of upper crustal rocks to depths equivalent to 8–9 kbar. Following the folding and vertical extrusion event, sub-horizontal fabrics developed at mediumto low pressure in the orogenic lower and middle crust during vertical shortening. Fabrics that record theearly vertical extrusion originated between 350 and 340 Ma, during building of an orogenic root inresponse to SE-directed Saxothuringian continental subduction. Fabrics that record the later subhorizontalexhumation event relate to an eastern promontory of the Brunia continent indenting into therheologically weaker rocks of the orogenic root. Indentation initiated thrusting or flow of the orogeniccrust over the Brunia continent in a north-directed sub-horizontal channel. This sub-horizontal flowoperated between 330 and 325 Ma, and was responsible for a heterogeneous mixing of blocks andboudins of lower and middle crustal rocks and for their progressive thermal re-equilibration. Theerosion depth as well as the degree of reworking decreases from south to north, pointing to an outflow oflower crustal material to the surface, which was subsequently eroded and deposited in a foreland basin.Indentation by the Brunia continental promontory was highly noncoaxial with respect to the SEorientedSaxothuringian continental subduction in the Early Visean, suggesting a major switch of plateconfiguration during the Middle to Late Visean.Key words: Bohemian Massif; channel flow; exhumation; orogenic lower crust; Variscan belt.INTRODUCTIONCurrent concepts of exhumation of deep-seated rocksin convergent orogens are generally based on the styleof the pressure–temperature–time (P–T–t) pathretrieved from high-pressure (HP) to ultra-highpressure(UHP) rocks (e.g. Duchene et al., 1997). Onegroup of exhumation mechanisms for these rocks hasbeen inferred from conceptual or numerical modelsdriven by subduction–accretion processes that result ineither corner flow circulation within an accretionarywedge (Platt, 1986, 1993; Allemand & Lardeaux, 1997;Gerya & Stockhert, 2006) or buoyancy-driven exhumationof subducted continental crust (Chemendaet al., 1995). Another group of conceptual models hasbeen developed for gravity-driven exhumation of HProcks in thickened orogenic root systems. In thesemodels, processes such as convective removal of atectospheric root (England & Houseman, 1988;Andersen et al., 1991) or lateral variations in gravitationalpotential energy of thickened continentallithosphere (Milnes & Koyi, 2000; Rey et al., 2001;Vanderhaeghe & Teyssier, 2001) drive the exhumation.Recently, an alternative model has been developed forthe large-scale horizontal movement of melt-bearingmiddle crust based on channel flow (Beaumont et al.,2001, 2006; Godin et al., 2006), and has been appliedto explain the ductile extrusion of medium-pressuremetamorphic rocks along the Himalayan front (Grujicet al., 1996; Jamieson et al., 2002, 2004). Whereas thefirst group of models focuses on explaining verticaldisplacements of UHP and HP rocks, the second groupof models emphasizes the importance of large-scalehorizontal displacements in orogens.Monocyclic continuous models of orogenic lowercrust exhumation may be supported by structural andkinematic field studies that emphasize the 2D characterof the exhumation process (e.g. Milnes & Koyi, 2000).However, more commonly there are complexities inthe 3D character of the tectonic evolution of largeÓ 2007 Blackwell Publishing Ltd 273133

274 K. SCHULMANN ET AL.orogenic root systems that arise from the polyphasenature of vertical and horizontal material and heattransfer during long-lasting orogenic events that makethese systems more difficult to unravel. The polyphaseand discontinuous character of orogeny may be due tomajor changes in plate configurations (e.g. Deweyet al., 1989) or to the existence of inherited rheologicalheterogeneities and variations in mechanical anisotropy(e.g. Burg, 1999). Therefore, detailed regionalstructural, petrological and geochronological studiesof orogenic fabrics may provide a key to decipheringthe succession, and length and time scales of processesresponsible for material and heat transfer within theselarge orogenic root systems.Classically, the Palaeozoic Variscan orogen inWestern and Central Europe (Fig. 1) is a large, hot,bivergent orogen that is interpreted to have developed50S.ArmoricanF.BrayF.Cadomian0TBBav. F.ElbeF.hercynianRhenohercynianthuringianSaxothuringianMoldanubian14 1Bruni runia48-40N50 kmMOLDANUBIANMonotonous andVaried GroupGföhl unitTEPLA-BARRANDIANProterozoicEarly Palaeozoic4WEST SUDETES52 oLUGIANPRAHA50 o 48 o12 oSAXOTHURINGIANduring prolonged convergence between Laurussia andGondwana (e.g. Ziegler, 1986; Matte et al., 1990).Traditionally, burial and exhumation of UHP and HProcks are thought to be the result of a kinematic continuumof coaxial subduction and subsequent collisionprocesses (e.g. OÕBrien & Carswell, 1993; Konopa´ sek& Schulmann, 2005), and as a consequence petrologicaland geochronological data may fit in one of the2D conceptual models discussed above.In this study, we present structural, petrological andgeochronological data acquired during the last twodecades from the eastern termination of the EuropeanVariscan front within an area of about 20 000 km 2(Figs 1 & 2). It is shown that exhumation of the orogeniclower crust occurred during two distinct periodsrelated to a major change in the configuration oflithospheric plates during the Visean. The first exhu-TEPLA-BARRANDIANELBE -ZONESaxothuringiansubductionSouth Central Bohemian Bohemian pluton plutonC retaceousBasinMOLDANUBIAN DOMAINMORAVO-SILESIANBRUNIA14 o 16 o BrnoFig. 1. Outline geological map of the Bohemian Massif with major units shown schematically (modified after Franke, 2000).Upper inset is the position of the Bohemian Massif in the framework of the European Variscides (after Edel et al., 2003). The areadiscussed in this review is marked by a rectangular box.Ó 2007 Blackwell Publishing Ltd134

EXHUMATION IN LARGE HOT OROGEN 275mation event was related to almost vertical materialtransfer, which was driven by S-E directed oceanic andcontinental (Armorican–Saxothuringian) subductiondynamics during the Devonian and Tournasian toEarly Visean (Fig. 1). Following this early evolution,during the Middle to Late Visean indentation from theeast by a promontory of the Brunia continent at a highangle to the Saxothuringian subduction directionresulted in horizontal flow and transport of theextruded orogenic lower crust over the rigid continentas a hot fold nappe.GEOTECTONIC SETTINGSuess (1912, 1926) described the geology of the easternVariscan front and divided the crystalline complexes ofthe Bohemian Massif into two parts, an internalMoldanubian–Lugian domain and an external Moravo–SilesianZone (Figs 1 & 2). Dudek (1980) completedthis subdivision and defined a Brunia continentwith the Moravo–Silesian Zone as its westerndeformed margin. The eastern segment of the Bruniacontinent is built up of unmetamorphosed to weaklymetamorphosed Neoproterozoic granites, high-gradeschists and migmatites. This basement is unconformable,covered by Lower Carboniferous foreland basinsedimentary rocks, by Devonian shallow marine sedimentaryrocks, and locally by Cambro-Ordovicianclastic, pelitic metasediment rocks and bimodal metavolcanicrocks (Franke, 2000; Hartley & Otava, 2001).The Moravo–Silesian Zone represents a NE–SWtrendingbelt of sheared and metamorphosed rocksderived from the Brunia continent. This 300 km long,30)50 km wide belt consists of three NE–SW-elongatedtectonic windows emerging through structurallyoverlying high-grade rocks of the Moldanubian–Lugian domain: a southern Thaya window; a centralSvratka window; and a northern Silesian domain(Fig. 2). Schulmann et al. (2005) identified the Moldanubian–Lugiandomain in the Bohemian Massif asthe deep orogenic root system of the Variscan orogen(Fig. 2). The Elbe zone (Fig. 2) divides rocks of theMoldanubian–Lugian domain into two: a larger highgradeMoldanubian domain to the south and a smallerhigh-grade Lugian domain to the north (Suess, 1926).The Moldanubian domain (Figs 1 & 2) has beensubdivided into three major lithotectonic units (e.g.Fuchs, 1986), namely the amphibolite facies Monotonousand Varied Groups, which together with gneissesmake up the orogenic middle crust, and the predominantlygranulite facies Gfo¨ hl Unit, which is inferred tobe the orogenic lower crust (Fig. 2). The Varied Groupoutcrops structurally above the Monotonous Group,with a contact that is commonly marked by bodies ofgranite gneiss. The Monotonous Group consists ofmigmatitic paragneisses (metagraywacke) interlayeredwith granite orthogneisses, quartzites and rare eclogites(Dudek & Fediukova´ , 1974; Petrakakis, 1997). TheVaried Group consists of a thick sequence of paragneissesinterlayered with calcsilicate rocks, marbles,quartzites, graphite schists, amphibolites and felsicmetavolcanic rocks. The Gfo¨ hl Unit is composed oflarge areas of migmatitic granite gneiss, called theGfo¨ hl gneiss, and of areas of various highly anatecticmigmatites and paragneisses that are in placesaccompanied by migmatitic amphibolites at the base.The Gfo¨ hl Unit includes numerous bodies of Ky–Kfsfelsic granulite as well as tectonic lenses of eclogite andgarnet and ⁄ or spinel peridotite.The main part of the Lugian domain is composed ofmedium- to high-grade granite gneisses and metamorphosedvolcano-sedimentary rocks with Cambro–Ordovician protolith ages (Kro¨ ner et al., 2001, andreferences therein). The granite gneisses contain boudinsof eclogite and a belt of garnet–omphacite granulite,and are considered an equivalent of the Gfo¨ hlUnit, whereas the belts of medium-grade schists ofvolcano-sedimentary origin are regarded as an equivalentof either the Monotonous or the Varied Groups(Fig. 2). An Ordovician leptyno-amphibolite lowercrustal complex (the Stare´ Město belt) occurs in theeastern part of the Lugian domain (Sˇtı´ pska´ et al.,2001).DEFINITION OF BASEMENT AND OROGENICCRUSTAL LEVELSIn this section, multiple criteria, such as the peakpressure conditions attained by individual units, thecharacter and the age of the protolith and chronologyof metamorphic zircon, are used to decipher the relativevertical position of crustal units during the LowerPalaeozoic and the Devonian. This information is thenused to propose a model for the stratification of theorogenic crust along the eastern Variscan front.The Brunia continentBased on zircon protolith ages (Fig. 3a) and 40 Ar ⁄ 39 Arcooling ages ranging from 600 to 540 Ma, granites andmigmatites forming the basement of the Brunia continentare inferred to have originated during the Pan-African orogenic events (Van Breemen et al., 1982;Fritz et al., 1996; Finger et al., 2000; Friedl et al.,2004). Tectonic imbrication of Devonian metasedimentaryrocks and basement, and metamorphism tolower greenschist facies occurred at the western marginof the Brunia continent during the Carboniferous(Francu˚ et al., 2002). In the northern Silesian domain(Fig. 2), the metamorphic pattern of nappes derivedfrom the Brunia continent shows an inverted Barrovianzonation ranging from chlorite grade in the east tokyanite ⁄ sillimanite grade in the west (Sˇtı´ pska´ &Schulmann, 1995; Schulmann & Gayer, 2000). Thewestern termination of the Moravo–Silesian Zone inthe north reached eclogite facies conditions (Fig. 4;Sˇtı´ pska´ et al., 2006), and it is likely that the western tipof the Moravo–Silesian Zone in the south has beenÓ 2007 Blackwell Publishing Ltd135

276 K. SCHULMANN ET AL.Moldanubian-LugianOROGENIC ROOTOrogenic upper crustProterozoic andPaleozoic sedimentsOrogenic middle crustMonotonous groupVaried groupGneissesOrogenic lower crust(Gföhl Unit in Moldanubian)Granulites, gneissesMetasedimentsOrdovician mafic lower crustLUGIANBELBE ZONEC′Fig. 7SilesianB’ B′MOLDANUBIAN0 20 40 60MAGMATIC ROCKSSyenitesTonalitesCMPAS outhernC entralNorthernCSvratk vratka aThay haya aA′Moravo-SilesianBRUNIABRUNIAFig. 6Fig. 5Foreland basinCulmPara-autochtonous unitsCoverBasementAlochtonous unitsCoverBasementMoravianMicaschist ZonePermian coverLate granitesLetovice ophioliteFig. 2. Map of the eastern margin of the Bohemian Massif to show inferred orogenic crustal levels and the locations of Figs 5, 6 & 7,which are indicated by rectangular boxes. Also shown are the positions of interpretative cross sections A–A¢, B–B¢ and C–C¢ presentedin Fig. 10.Ó 2007 Blackwell Publishing Ltd136

Probabi l irequencyProbabi l irequencyEXHUMATION IN LARGE HOT OROGEN 277underthrust to eclogite conditions as well (Konopa´ seket al., 2002).The orogenic rootThe Gföhl Unit and Lugian high-grade rocksThe Gfo¨ hl Unit (Fig. 2) is composed predominantly offelsic granulites (11–20 kbar, 800–1000 °C) and highgradegneisses (e.g. OÕBrien & Ro¨ tzler, 2003; Sˇtı´ pska´ &Powell, 2005b) that contain eclogite lenses of basalticcomposition, suggesting a crustal origin (18–19 kbar,800)900 °C, Fig. 4; e.g. Medaris et al., 1998; Sˇtípska´& Powell, 2005a and references therein). The eclogitesand granulites, as well as other lithologies of the Gfo¨ hlUnit, were overprinted at amphibolite facies conditions(4)12 kbar, 750)850 °C, Fig. 4; e.g. Petrakakis,1997; Hasalova´ et al., 2008b). The Gfo¨ hl Unit containsnumerous tectonically emplaced bodies of garnet andspinel peridotite (28–44 kbar, 900–1300 °C, Fig. 4),associated with lenses of eclogite (16–20 kbar, 800–950 °C) and clinopyroxenite, representing partial meltcrystallized in the upper mantle (Medaris et al., 1995).The Lugian eclogites and granulites (18–20 kbar, 800–900 °C, Fig. 4; e.g. Steltenpohl et al., 1993; Sˇtípska´et al., 2004) are considered to reflect UHP conditions(Kryza et al., 1996), but this has not been confirmed;they were re-equilibrated at variable P–T conditions(4)11 kbar, 700)800 °C, Fig. 4).Multigrain zircon and monazite fractions fromgranulites and high-grade gneisses of the Gfo¨ hl Unityielded conventional upper intercept U–Pb agesbetween 550 and 510 Ma, interpreted to date the formationof the protoliths (Fig. 3b; e.g. Van Breemenet al., 1982; Schulmann et al., 2005). In addition,xenocrystic cores and prismatic zircon from felsicgranulites and Gfo¨ hl gneiss yield Silurian to DevonianU–Pb ages also interpreted to date the formation of theprotoliths (Fig. 3b; Friedl et al., 2004). However, amajority of U–Pb zircon ages fall in the range360)340 Ma with a prominent peak at c. 340 Ma,which is interpreted to record the age of Variscanmetamorphism (Fig. 3d). Ultra-potassic syenites (durbachites)spatially related to Ky–Kfs granulites yieldages around 338 Ma and 325 Ma (Fig. 3e; Holubet al., 1997; Janousˇek & Holub, 2007).The Monotonous and Varied GroupsIn the Moldanubian domain, paragneisses of theMonotonous and Varied Groups generally recordmedium-pressure metamorphism (8)9 kbar, 610)660 °C) associated at higher temperatures with widespreadanatexis ( ~ 9 kbar, 700)800 °C), followed byre-equilibration in the sillimanite stability field atconditions around 4)6 kbar and 600)800 °C (Fig. 4;e.g. Petrakakis, 1997). The metamorphic conditionsof the Monotonous Group paragneisses aret y fBrunia and Moravo-Silesian0.7(a)503 Ma0.6580 Ma0.50.4684 Ma0.30.2Protolith0.1U–Pb0350 400 450 500 550 600 650 700 7504(b)3.532.521.510.5Orogenic middle crust503 MaLugian (60)Moldanubian (5)ProtolithU–Pb0350 400 450 500 550 600 650 700 750Orogenic lower crust0.7(c)0.6413 Ma0.5389 Ma0.4550 Ma0.3487 Ma430 Ma0.20.1Moldanubian (14)Lugian (9)ProtolithU–Pb0350 400 450 500 550 600 650 700 750t y f3.5(d)3 340 Ma2.520.510.5359 MaMoldanubian (22)Lugian (15)1.4(e)1.2MetamorphicMagma.0.2U–PbU–Pb00300 320 340 360 380 400 420 280 300 320 340 360 380 40010.80.60.4335 Ma324 Ma340 Ma353 MaFig. 3. Probability curves of age histograms to show existing zircon U–Pb ages. For the Brunia continent and the Moravo–SilesianZone (a), Moldanubian and Lugian domain (b, c & d) and intrusive rocks (e). BB1, Brunia basement south of the Elbe zone; BB2,Brunia basement north of the Elbe zone; MB1, Moravian domain basement rocks; MB2, Silesian domain basement rocks; MSZ,Moravian Micaschist Zone; ID, Mg-rich syenites (durbachites) of the orogenic lower crust; IMC, intrusive rocks of the middle crust;IT, tonalite–granodiorite intrusions; IUC, granites in upper crustal units. Numbers in parentheses stand for number of available ages.Ó 2007 Blackwell Publishing Ltd137

278 K. SCHULMANN ET AL.Fig. 4. P–T conditions for (a) peak metamorphism in the orogenic root and Brunia continent south of the Elbe Zone; (b) peakmetamorphism in the orogenic root and Brunia continent north of the Elbe Zone; (c) retrograde conditions in the orogenic rootand Brunia continent south of the Elbe Zone; and (d) retrograde conditions in the orogenic root and Brunia continent north of theElbe Zone. Source of data: Carswell & O¢Brien (1993), Steltenpohl et al. (1993), Medaris et al. (1995), Sˇtı´ pska´ & Schulmann (1995),Bro¨ cker & Klemd (1996), Kryza et al. (1996), Pitra & Guiraud (1996), Parry et al. (1997), Petrakakis (1997), Medaris et al. (1998),Klemd & Bro¨ cker (1999); Romanova´ &Sˇtı´ pska´ (2001); Konopa´ sek et al. (2002); OÕBrien & Ro¨ tzler (2003); Sˇtı´ pska´ et al. (2004); Lexaet al. (2005); Sˇtı´ pska´ & Powell (2005a,b); Racek et al. (2006); Sˇtı´ pska´ et al. (2006); Tajcˇmanová et al. (2006); also shown are steadystategeotherms calculated for various mantle heat flows (as indicated in the figure) and exponential radioactive heat production.commonly re-equilibrated in the stability field ofcordierite at conditions around 4.5–6 kbar and 600–720 °C (Petrakakis, 1997), inferred to be due to thethermal effect of Carboniferous intrusions emplacedat 330)310 Ma (Fig. 3e). Despite the overall medium-pressurecharacter, in the Moldanubian domainthe Varied and Monotonous Group rocks includerare lenses of eclogite that register conditions of 13–16 kbar at 600–680 °C (Fig. 4; Medaris et al., 1995).The peak metamorphism of the volcano-sedimentaryrock series in the Lugian domain varies from garnetto kyanite grade, yielding 7)9 kbar at 550)650 °C(Romanova´ & Sˇtı´ pska´ , 2001; Jastrzębski, 2005;Murtezi, 2006) with re-equilibration in the sillimanitestability field at sub-solidus conditions of 5–6 kbar ataround 600 °C.In the Monotonous and Varied Groups of theMoldanubian domain, scarce U)Pb zircon data fromorthogneisses and metavolcanic rocks yield Neoproterozoicand Cambrian protolith ages (Fig. 3c;Schulmann et al., 2005). U)Pb zircon data frommetavolcanic rocks and granites of the Lugian domainÓ 2007 Blackwell Publishing Ltd138

EXHUMATION IN LARGE HOT OROGEN 279yield predominantly Cambrian protolith ages (Fig. 3c;e.g. Kro¨ ner et al., 2000 and references therein).Neoproterozoic and Lower Palaeozoic low-grade sequencesA sequence of weakly metamorphosed Neoproterozoicand Silurian to Devonian sedimentary rocks is separatedby a crustal-scale detachment from the underlyingmedium- and high-grade rocks of the Moldanubianand Lugian domains (Pitra et al., 1994; Mazur et al.,2005). The NE–SW-trending contact of the uppercrustal sedimentary rocks and middle crustal metamorphicrocks is marked by granodiorite sills, whichyield U–Pb zircon ages of 350–340 Ma (Fig. 3e).Intrusion of the sills was accompanied by metamorphismof the hangingwall rocks at conditions around4 kbar and 550)590 °C (Fig. 4). Locally, granite plutons,which yield U–Pb zircon ages around 330 Ma,intrude these low-grade rocks (e.g. Schulmann et al.,2005).define these units as representing the orogenic middlecrust. The Monotonous and Varied Groups and lowgradeunits yield only Neoproterozoic and Ordovicianprotolith ages, they lack a Devonian thermal reworkingand they are interpreted as forming shallowercrustal levels during the Devonian and Carboniferous.Using the same reasoning, the weakly metamorphosedNeoproterozoic and Lower Palaeozoic rocks that showpeak pressures around 4 kbar, and no evidence ofDevonian thermal reworking are defined as the orogenicupper crust.During the Carboniferous history of building theorogenic root, the rocks occupying structurally deeperpositions were transported upwards, whereas shallowerunits were moved downwards. The result is thatthe orogenic lower and middle crusts now form subparallelbelts at the surface. Exceptionally, the Variedand Monotonous Groups reached lower orogeniccrustal conditions during their burial, and in this case,they become part of the orogenic lower crust.Definition of crustal levels within the Moldanubian–Lugianorogenic rootThe lithologies and metamorphic conditions of themain units of the Moldanubian and Lugian domainsare used to define a lithotectonic zonation that correspondsto different crustal levels that existed duringmaximum thickening of the orogenic root. In turn, thepresent map distribution of the different crustal levelsdefines the spatial pattern of crustal units that experienceddifferent vertical displacements. Interpretationof these different crustal levels associated with thickeningand the differences in vertical displacement, togetherwith the geometry and structural record withineach unit may allow us to unravel the mechanism(s) ofexhumation of these units.According to the lithologies and peak P–T conditionsdescribed above, the Gfo¨ hl Unit is a granulite–eclogite unit representing the lower part of a thickenedcrust that reached lithostatic pressures of 18–20 kbar,corresponding to a depth of 60–70 km, which wedefine as the orogenic lower crust. In addition, thegeochronological and petrological data show a systematiccorrelation of protolith age, metamorphic ageand peak pressures (Figs 3 & 4). Only the high-graderocks of the Gfo¨ hl Unit and the Lugian granulitesrecord Devonian protolith ages, Early Carboniferousmetamorphic ages and peak pressures of 18–20 kbar.For these reasons, Schulmann et al. (2005) suggestedthat the orogenic lower crust in the Bohemian Massifrepresents Neoproterozoic–Cambro-Ordovician continentalcrust that experienced a major thermalreworking during Devonian intra-continental (backarc)rifting.The Monotonous and Varied Groups yield peakpressure conditions not exceeding 12 kbar and areinterpreted to have reached maximum crustal depthsaround 40 km under an elevated thermal gradient. WeTECTONICSOFTHEEASTERNPARTOFTHEBOHEMIAN MASSIFStructural evolution of the Brunia continental margin: theMoravo–Silesian ZoneThe Brunia continent forms a pile of internallyimbricated basement- and cover-derived thrust sheets,termed the Moravo–Silesian Zone, which developedduring Carboniferous dextral–oblique thrusting of theMoldanubian–Lugian domain (Schulmann et al.,1991). As a result, the whole Moravo–Silesian Zoneexperienced intense non-coaxial NE-directed deformation(Schulmann et al., 1994; Fritz et al., 1996).Crustal-scale folds with west-plunging hinges refoldthe nappe pile during the final stages of NE-directeddeformation (Schulmann, 1990; Sˇtı´ pska´ et al., 2000).Finally, the Svratka dome and the Silesian domainwere folded by west-facing folds that also affected theCulm foreland basin, suggesting the age of latestshortening to be around 310 Ma (Hartley & Otava,2001).Orogenic structure of the crustal rootThe eastern branch of the Moldanubian domainlocated between Central Moldanubian pluton andMoravo–Silesian Zone (Fig. 1) is examined here. Inorder to characterize structural and petrological variationsalong-strike of the orogenic margin we introducethe southern, central and northern Moldanubiandomains.The large-scale structure of the Moldanubiandomain is characterized from west to east as follows:(i) first, there is a sequence of east-dipping schists andmarbles of the Monotonous and Varied groups, whichis intruded by the Central Moldanubian pluton; (ii) tothe east, this sequence is overlain by high-gradeÓ 2007 Blackwell Publishing Ltd139

280 K. SCHULMANN ET AL.gneisses, migmatites and granulites containingnumerous inclusions of peridotites and eclogites; (iii)further east, this orogenic lower crust is thrust overwest-dipping middle crustal rocks that are composedof sillimanite micaschists, orthogneiss or leptyniticamphibolite of Cambro-Ordovician age in the south(Figs 5c & 6c); and (iv) finally, the easternmostextremity of the Moldanubian sequence is characterizedby a lower crustal segment of mafic and felsicgranulites associated with migmatitic biotite paragneissand a few mantle slivers (Fig. 5c). All these unitstrend parallel to the continental margin represented byMoravian nappes of the Thaya and Svratka domes,respectively (Fig. 2). This sequence of rocks is generallyvalid for the northern and southern parts of theMoldanubian domain, but in contrast the central partis dominated by orogenic lower crust rocks intruded byMg–K-rich syenite intrusions (figs 4 & 5 in Schulmannet al., 2005).The geological structure of the Lugian domain(Fig. 7) is characterized by alternations of orthogneisseswith metasedimentary and metavolcanic rocksthat comprise the orogenic middle crust. In the centralpart of the orthogneiss belt there is a narrow NEtrendingbelt of garnet–omphacite granulite and migmatitethat represents the orogenic lower crust. ACarboniferous granodiorite sill marks the easternboundary of the Lugian domain, with an Ordovicianleptyno-amphibolite unit extending further to the east(Fig. 7c). Directly in the footwall of the Lugiandomain there is an eclogite unit derived from the lowercrust of the Brunia continent (Sˇtı´ pska´ et al., 2006).Imbricated Brunia basement and cover thrust sheets ofthe Moravo–Silesian domain represent the structurallydeepest unit of the section. Superposed orogenic fabricsobserved in the field and their petrological characteristicsin four representative areas are describedbelow, three from the Moldanubian domain and onefrom the Lugian domain, where the spatial relationshipsbetween the orogenic lower crust and the adjacentorogenic middle crust may be critically evaluated.Structural development of the Moldanubian domainThe most characteristic structural feature of the Moldanubiandomain is the widespread flat-lying foliation.In the southern Moldanubian domain (Figs 2 & 5), theflat fabric gently dips to the east in the western partand to the west in the eastern part. This geometry ledAustrian geologists to define a large-scale Gfo¨ hl napperesting on middle crustal rocks in synformal structures(Tollmann, 1982; Fuchs, 1986). However, recentstudies have shown that the sub-horizontal fabrics (S 3 )were superimposed on vertical fabrics (S 2 ), andalthough the superposition is of variable intensity, itoccurs in almost all the rock types of the lower andmiddle crust.For the Moldanubian domain, we review three areaswhere the relationship of the steep to the flat fabricshas been studied in detail (Kolenovska´ et al., 1999;Racek et al., 2006; Tajcˇmanova´ et al., 2006; P. Sˇtı´ pska´& K. Schulmann, unpublished data). First, the structuralsuccession is described on the macroscopic scale,where the S 2 fabric, defined mainly by steep lithologicalalternation and ⁄ or steep internal rock layering, wasreworked by multiple mechanisms into a flat-lying S 3fabric of variable intensity (Fig. 8). Later a detaileddiscussion is given of the metamorphic characterizationof the steep S 2 fabric and metamorphic conditionsof the flat D 3 reworking in individual lithologies.The first area is located in the northernmost terminationof the southern Moldanubian domain (Figs 2 &5; Racek et al., 2006). Here, exceptionally well-preservedS 2 fabrics occur as steep alternations of paragneisses,marbles, felsic volcanic rocks, amphibolitesand quartzites of the middle crust that are surroundedby migmatites and granulites of the orogenic lowercrust (Fig. 5, stereoplot E). The S 2 fabric was the resultof isoclinal steep folding and from transposition ofcompositional banding S 1 in amphibolites (Fig. 8d).The S 2 foliation was refolded by open to closerecumbent F 3 folds in the central part of the middlecrustal domain (Fig. 8f), and approaching the contactswith surrounding orogenic lower crust the degree offlat D 3 reworking increases in conjunction with developmentof S 3 axial planar schistosity (Fig. 8e). Themineral stretching lineation trends NE–SW over thewhole area, in accordance with the orientation of thehinges of the F 3 folds (Fig. 5, stereoplot F). The S 3foliation clearly changes its orientation and dip, followingthe form of the orogenic middle crust, andsimilarly, the L 3 lineation plunges either to the SSW orFig. 5. (a) Structural map of a critical area of the southern Moldanubian domain with S 2 and S 3 fabrics (after Racek et al., 2006). SeeFig. 2 for regional location. The map shows a large area of orogenic middle crust, called the Drosendorf window, that is entirelysurrounded by orogenic lower crustal rocks. Structural trends (thin lines – S3 fabrics, thick lines – S 2 fabrics) indicate extrapolationsof the major orientation of structural fabrics in the field. The density of trend lines indicates the homogeneity of fabric elements inthe field. (b) Stereograms: A – stereogram showing poles to S 3 fabrics in the area of orogenic lower crust north of the orogenicmiddle crustal domain; B – poles to S 3 fabrics from the orogenic lower crust and orogenic middle crust close to the northerntermination of the orogenic middle crust domain; C – poles to S 3 fabrics from the orogenic lower crust and orogenic middle crustalong the western border of the orogenic middle crust domain; D – poles to S 3 fabrics from the orogenic lower crust underlying theorogenic middle crust domain; E – stereogram showing poles to relicts of S 2 preserved within central part of the orogenic middlecrust domain; F – stereogram showing L 3 mineral and stretching lineations from the whole area. Equal area projection, lowerhemisphere, contoured at multiples of uniform distribution. (c) Interpretative cross-section shows main structural features describedin the text. The numbers on the cross-section and map show locations of samples used for P–T calculations shown in Fig. 9.Ó 2007 Blackwell Publishing Ltd140

EXHUMATION IN LARGE HOT OROGEN 281NNE according to the lenticular shape of the orogenicmiddle crustal (Fig. 5, stereoplots B, C & D).The second area is the southernmost part of the3000 km 2 region of orogenic lower crust of the centralMoldanubian domain (central part of Fig. 2, andnorthern part of Fig. 5; Kolenovska´ et al., 1999). Here,the orogenic lower crust is a vast domain typicallycomposed of felsic migmatitic orthogneisses andmigmatites containing large bodies of Ky–Kfsgranulite, eclogite and peridotite. The whole area is(a)(b)ABDPaPaMiAmFeMiMi(c)PaPaMiAmMiFeAmMiOrÓ 2007 Blackwell Publishing Ltd141

282 K. SCHULMANN ET AL.characterized by a flat migmatitic S 3 foliation (Fig. 5,stereoplot A) that is axial planar to rare isoclinal foldswith hinges generally trending NNE–SSW, parallel tostretching and mineral lineations. Numerous shearbands filled with leucosome indicate a dominant topto-the-NEshearing (Urban, 1992; Schulmann et al.,1994). The eclogites and granulites form lozengeshapedboudins surrounded by migmatitic fabric.These rocks locally contain relics of early NE-trendingsteep S 2 fabric, but because of their scarce occurrence,(a)(b)PaMiMiOrFeSy(c)PaPaOrFeÓ 2007 Blackwell Publishing Ltd142

EXHUMATION IN LARGE HOT OROGEN 283we cannot demonstrate the regional extent of the S 2fabric in this area (Fig. 8h).The third area is located in the northern Moldanubiandomain (Figs 2 & 6) and is characterized bythrusting of a NNW–SSE-elongated granulite bodyover the middle crust (Fig. 6a, c). The granulite showsan exceptionally well-preserved steep NNW–SSEtrendinggranulite-to-amphibolite facies S 2 foliation(Fig. 6, stereoplot A). D 3 shallow to moderately steep,south-dipping shear zones, which show top to theNNE shear sense, cut the S 2 fabric heterogeneously.The surrounding migmatites and migmatitic orthogneissesexhibit predominantly S 3 fabric, which dipsgently to moderately to the south (Fig. 6, stereoplotB). This foliation, formed essentially by migmatiticbanding, contains numerous close to isoclinal F 3 foldswith hinges trending NNE–SSW parallel to minerallineation. Within the adjacent middle crust, thedeformation is also polyphase, comprising an earlyfoliation that dips steeply to the SW, which was reworkedby a schistosity that dips gently in the samedirection (Fig. 6, stereoplot C).Structural development of the Lugian domainThe northern part of the Lugian domain is characterizedby a central granulite belt surrounded on thewestern and eastern sides by strongly migmatized orthogneisses(Figs 7c & 8b). The granulite and theadjacent orthogneisses reveal the NNE–SSW-trendingvertical S 2 foliation parallel to the trend of the orogeniclower crust (Fig. 7, stereoplot A). This S 2 foliationcontains rootless folds formed by layers of metabasitethat preserve evidence of an early S 1 foliation (Sˇtípska´et al., 2004).East of the granulite belt, the orthogneisses andmigmatite show reworking of the vertical S 2 fabric bymoderately west-dipping S 3 foliation (Fig. 8g). The dipof the S 3 fabric progressively decreases eastwards towardsthe underlying Ordovician lower crustal leptyno-amphiboliteunit (Fig. 7, stereoplot B); agranodiorite sill emplaced syntectonic with the D 3deformation marks the boundary between this unitand the Lugian orthogneiss, dating the D 3 structures atc. 340 Ma (Parry et al., 1997; Sˇtípska´ et al., 2001).Importantly, the structure of the leptyno-amphiboliteunit is discordant with respect to D 3 fabrics of thewestern Lugian orthogneisses; the leptyno-amphiboliteunit is inferred to be of Early Ordovician age, based onU–Pb zircon dates that yield an age of c. 510 Ma(Fig. 7, cross-section; Sˇtı´ pska´ et al., 2001; Lexa et al.,2005).To the NW from the granulite belt, the S 2 fabric isaffected by sub-horizontal, variably dipping amphibolitefacies S 3 fabric (Fig. 8a,b). This S 3 fabric hasintensely reworked the steep S 2 foliation of the adjacentorogenic middle crust to the point that it progressivelypasses into a several hundred metre widenormal-sense shear zone dipping gently to the NE(Fig. 8c).P–T–t history of the Moldanubian–Lugian rootThe relationship between the prograde andretrograde P–T paths and orogenic fabrics is a keypiece of information necessary to understand thethermo-mechanical processes that operated withinthe orogenic root. Microstructural, petrological andthermodynamic modelling studies during the lastfive years have revealed a systematic pattern of P–Tevolution related to the early steep fabrics in boththe orogenic lower and middle crust. In addition,combination of the P–T evolution with Sm–Nd and40 Ar ⁄ 39 Ar cooling ages reveals important differencesin the thermo-chronological evolution of theorogen.Microstructural and petrographic characterization of S 2 andS 3 fabricsIn the orogenic lower crust, the steep fabric is characterizedby compositional banding formed by maficand felsic granulite (Sˇtı´ pska´ et al., 2004). More commonlyS 2 is defined by a mineral fabric, marked byoriented kyanite, biotite and recrystallized ribbons ofquartz and feldspar in the case of the felsic granulite(Tajcˇmanova´ et al., 2006), and by the elongated shapeof garnet and aligned omphacite grains in the case ofthe mafic granulite (Sˇtı´ pska´ et al., 2004). In the orogeniclower crust composed of migmatitic orthogneisses,the S 2 fabric is defined by alternation of infinitemonomineralic recrystallized bands of quartz, plagioclaseand K-feldspar, in an assemblage with kyanite,biotite and garnet that results from a high-temperatureFig. 6. (a) Structural map of an area showing thrusting of the orogenic lower crust over the orogenic middle crustal in the NE part ofthe Moldanubian domain, with regional S 2 and S 3 fabrics as shown in Fig. 5 (after Tajčmanova´ et al., 2006). See Fig. 2 for regionallocation. The map shows a large body of orogenic lower crustal granulites that preserve relicts of the S 2 fabrics. The structuraltrends (thin layers – S 3 fabrics, thick layers – S 2 fabrics) indicate extrapolations of major orientations of structural fabrics in the field.The density of trend-lines indicates the homogeneity of fabric elements in the field. (b) Stereograms: A – stereogram shows poles to S 2fabrics in the orogenic lower crustal granulites; B – stereogram of poles to S 3 fabrics from the orogenic middle crust. Squares instereogram show integrated directions of L 3 lineations from the area. C – Stereogram of poles to S 3 fabrics from the orogeniclower crustal granulites and associated migmatites. Equal area projection, lower hemisphere, contoured at multiples of uniformdistribution. (c) Simplified cross-section shows well-preserved S 2 fabrics in competent granulites surrounded by cordierite migmatitescompletely reworked by S 3 fabric. The migmatites and granulites of the orogenic lower crust are thrust over orogenic middle crustto the east. The numbers on the cross-section and map show locations of samples used for P–T calculations in Fig. 9.Ó 2007 Blackwell Publishing Ltd143

284 K. SCHULMANN ET AL.(a)(b)OrMeOmLaGrMeFe(c)OrOrMeOmLaGrMeFeÓ 2007 Blackwell Publishing Ltd144

EXHUMATION IN LARGE HOT OROGEN 285and medium- to high-pressure recrystallization ofcoarse-grained granite (Hasalova´ et al., 2008a). Inareas inferred to be a part of the orogenic lower crust,based on the presence of HP relicts and the high degreeof anatexis, the S 2 fabric is defined by steep compositionallayering in amphibolites, felsic metavolcanicrocks, orthogneisses and paragneisses.In the orogenic middle crust the S 2 fabric is characterizedby a steep lithological alternation of paragneisseswith amphibolites, felsic metavolcanic rocks,quartzites, marbles and calcsilicate rocks (Romanova´&Sˇtípska´ , 2001; Racek et al., 2006). In the amphibolites,the steep fabric is characterized by alignedhornblende and recrystallized plagioclase ribbons inassociation with garnet that indicates medium-pressureconditions for formation of this fabric. Evidence of theprograde fabrics in the paragneisses of the Moldanubiandomain occurs only as rare oriented inclusions ofstaurolite and kyanite in garnet and remnants of theseminerals in the matrix (Racek et al., 2006). Even insome areas of the Moldanubian domain where themacroscopic fabric of the paragneisses remained steep,the higher-grade assemblages were commonly overprintedpassively by lower-grade assemblages with sillimaniteor cordierite, which we ascribe to the highreactivity of the paragneisses in the presence of melt.As the other lithologies do not show variation of theassemblage with changing P–T conditions, the mainguide in characterizing the steep fabric in the orogenicmiddle crust of the Moldanubian domain is the steeplithological alternation combined with peak P–T conditions.In the Lugian domain, in the orogenic middle crustthe characterization of the steep fabric is much betterconstrained because of the absence of anatexis in themetasedimentary rocks and the absence or limitedextent of anatexis in the orthogneisses. In the orthogneisses,the S 2 fabric is marked by steep alternation ofrecrystallized augen and ribbons of quartz and feldsparand by the orientation of biotite and muscovite. Steeplocalized ultramylonitic zones are developed locally.The prograde and peak P–T conditions of the steepfabrics in Lugian metapelites are constrained by thesuccessive growth of garnet, staurolite and kyanitein the micaschists (Romanova´ &Sˇtípska´ , 2001; Jastrzębski,2005; Murtezi, 2006).As described above, the flat fabric in the orogeniclower and middle crust commonly originates throughhorizontal folds affecting the steep fabric and is onlyrarely associated with sub-horizontal or gently dipping,thrust-related shear zones. Highly anisotropicrocks, such as banded orthogneisses, show strong axialplanar crenulation cleavage, which, in places, passesgradually into a sub-horizontal foliation with evidenceof the early vertical anisotropy occurring only as a fewrootless folds (Fig. 8h). The intensity and characteristicmineralogy of the flat reworking is variabledepending on the specific area, lithology and degree ofmetamorphism. However, some mineralogical featuresare common for the Ky–Kfs granulites, the orthogneissesand the paragneisses. These include thewidespread development of the assemblages garnet–sillimanite–K-feldspar, in the Moldanubian domain,and garnet–muscovite ± sillimanite, in the Lugiandomain, showing the overall moderate pressure characterof the S 3 reworking and the higher temperature inthe Moldanubian domain compared with the Lugiandomain. Because of the variability of the S 3 structuresalong the orogen, the detailed mineralogical characteristicsand P–T conditions for both the S 2 and S 3fabrics for the four areas identified above are describedbelow.P–T evolution of the Moldanubian domainRacek et al. (2006) examined petrological relationshipsbetween the steep and flat fabrics in the orogenic lowerand middle crust of the southern Moldanubian domain(Figs 5 & 9a). Modelling of garnet zoning in eclogitefrom the western part of the orogenic lower crust (no. 5in Figs 5 & 9a) indicates burial process to have takenplace from around 10 kbar and 700 °C to around 15–16 kbar and 800 °C. Similar peak conditions aredetermined for kyanite–sillimanite migmatite structurallyabove the eclogite (around 14 kbar at 750 °C;no. 6 in Figs 5 & 9a). However, these prograde andpeak conditions are associated with the steep S 2 fabriconly as the early structural relations of relict mineralsFig. 7. (a) Structural map of a critical area showing thin vertical belt of the orogenic lower crust surrounded by orogenic middlecrust in the eastern part of the Lugian domain with regional S 2 and S 3 fabrics as in Fig. 5 (after Sˇtı´ pska´ et al., 2004). See geological mapin Fig. 2 for regional location. The map shows a narrow body of the orogenic lower crustal granulites, preserving relicts of the S 2fabrics, and the surrounding fabric of the orogenic middle crust. The fabric in the mafic lower crust and the syntectonic tonalite is alsoshown. The structural trends (thin layers – S 3 fabrics, thick layers – S 2 fabrics) indicate extrapolations of major orientations ofstructural fabrics in the field. Density of trend lines indicates the homogeneity of fabric elements in the field. Stereograms:A – stereogram of poles to S 2 fabrics in the garnet–omphacite-bearing granulites; B – stereogram of poles to S 3 fabrics from theorogenic lower crust and surrounding orogenic middle crust rocks represented by thin dots and contoured. Thick dots in the stereogramsshow S 3 foliations typical for the Ordovician metagabbros and Carboniferous tonalite sill of the Lugian mafic complex.C – Stereogram of L 3 directions from the granulite belt and surrounding orogenic middle crust (thin dots and contoured data).Black squares show mineral and stretching lineations from mylonitized gabbros and granodiorite sill. Note strong discrepancy inorientations of S 3 foliations and L 3 lineations between granulite belt and adjacent Ordovician mafic complex. Equal area projection,lower hemisphere, contoured at multiples of uniform distribution. (c) Simplified cross-section shows the vertical central granulitebelt showing thrusting of the orogenic lower crust over the Ordovician leptyno-amphibolite complex. Note the syntectonic D 3intrusion of the 340 Ma granodiorite sill and discordant fabrics in the 510 Ma leptyno-amphibolite complex.Ó 2007 Blackwell Publishing Ltd145

286 K. SCHULMANN ET AL.Fig. 8. Field photographs of the S 2 –S 3 relationships. (a) Well preserved S 2 fabric in the garnetiferous amphibolite in the orogenicmiddle crust of the Lugian domain. (b) Strongly reworked S 2 fabric in the amphibolite in the orogenic middle crust of the Lugiandomain. (c) Complete transposition of early fabric in new S 3 in epidote amphibolite in the orogenic middle crust of the Lugian domain.(d) Excellent preservation of the steep S 2 fabric in garnetiferous amphibolite in the orogenic middle crust of the southern Moldanubiandomain. (e) S 2 folded by F 3 folds in the felsic granulite of the orogenic lower crust in the southern Moldanubian domain.(f) Asymmetrical folding of the S 2 fabric in the S 3 channel-flow fabric of the orogenic lower crust in southern Modanubiandomain. (g) Character of S 2 foliation in high-grade banded orthogneiss of the Lugian domain. (h) Microfolding and partial transpositionof high-grade S 2 fabric in the Gfo¨ hl gneiss of the southern Moldanubian domain.Ó 2007 Blackwell Publishing Ltd146

EXHUMATION IN LARGE HOT OROGEN 287were obliterated by strong S 3 reworking. High-pressureconditions are also reported from the Ky–Kfs granulitebody to the east (15 kbar, 800 °C, no. 2 in Figs 5 &9a), also part of the orogenic lower crust, where theretrieved P–T is clearly related to the S 2 fabric.The retrograde path of rocks from the orogeniclower crust situated far from the Brunia margin ischaracterized by an almost isothermal decompressionto about 7 kbar, as recorded by the S 3 assemblagegarnet-hornblende-plagioclase in eclogite and garnet–sillimanite–biotite in migmatite (no. 5, 6 in Fig. 9a).Unlike the central part of the Moldanubian, the Ky–Kfs granulite forming the orogenic lower crust at theboundary with the Brunia margin shows cooling withdecompression (no. 2 in Fig. 9a). Modelling of zoningin garnet that is syntectonic with the S 2 fabrics inparagneisses of the orogenic middle crust is consistentwith increase of P–T conditions to about 9 kbar and700 °C (nos 3 & 4 in Figs 5 & 9a). The retrograde P–Tpath of the sillimanite-bearing S 3 fabric in rocks of theorogenic middle crust is associated with decrease ofpressure coupled with increase of temperature, toabout 7 kbar and 700–750 °C (nos 3 & 4 in Fig. 9a).Kolenovska´ et al. (1999), Sˇtı´ pska´ & Powell (2005a)and P. Sˇtípska´ & K. Schulmann (unpublished data)examined the P–T evolution of rocks and relics of boththe orogenic lower and middle crust over a large areadominated by flat S 3 fabric in the central Moldanubiandomain (Fig. 2). Here, Ky–Kfs granulites register peakconditions around 18 kbar and 850 °C with no clearindication of the prograde path (no. 8 in Figs 5 & 9b).In contrast, a MORB-type eclogite provides evidenceof the prograde path from about 10 kbar and 750 °Cto 17–18 kbar in the form of inclusions of hornblendeand plagioclase in a prograde garnet (no. 7 in Fig. 9b).Both samples from the orogenic lower crust showalmost isothermal decompression and strong D 3reworking at 7)10 kbar and 750–850 °C (nos 7 & 8 inFig. 9b). Associated paragneisses from the orogeniclower crust contain early staurolite and kyanite inclusionsin garnet that preserves prograde zoning, indicatingburial to about 10 kbar and 700 °C (no. 9 inFigs 5 & 9b; P. Sˇtípska´ & K. Schulmann, unpublisheddata). The matrix was completely transposed by the D 3deformation, as marked by sillimanite and biotitedeveloped during decompression (no. 9 in Figs 5& 9b). In another paragneiss the flat sillimanite–biotite-bearingfabric was overgrown by garnet, indicatingheating at around 7)8 kbar and 700)850 °C (no. 10 inFig. 9b; P. Sˇtípska´ & K. Schulmann, unpublisheddata). The prograde conditions and peak pressuresfrom samples of this area could not be directly linkedto the steep fabric because of almost complete transpositionof rocks by the S 3 fabric. However, the progradeinclusions in garnet from the eclogite and fromthe paragneiss form straight, sub-vertical internalfabrics oriented at a high angle to the external S 3foliation, indicating that these garnet have probablygrown in the steep fabric.The steep S 2 fabric in the granulite body rimming theNE margin of the Moldanubian domain is marked bygarnet, kyanite and K-feldspar indicating conditions ofabout 18 kbar at 850 °C, whereas the crystallization ofbiotite in the same fabric occurred during decompressionaccompanied by slight cooling (no. 11 in Figs 6& 9c; Tajcˇmanova´ et al., 2006). Within the S 2 fabric,sillimanite–biotite intergrowths replacing garnet indicatesignificant decompression and cooling (no. 11 inFig. 9c). The S 3 fabric is associated with hercyniterimmingmetastable kyanite in the granulite and withthe development of cordierite, sillimanite and biotite inadjacent gneisses and migmatites, indicating conditionsof around 4 kbar and 700 °C (nos 12, 25 in Figs 6& 9c). Muscovite–biotite schists of the adjacent orogenicmiddle crust contain sillimanite that overgrowsrelicts of kyanite in the matrix. The peak P–T conditionsestimated by Pitra & Guiraud (1996) are 8–9 kbarat 610–660 °C, which was followed by near-isothermaldecompression, as recorded in the whole middle crustalcomplex, to 4–6 kbar (no. 13 in Figs 5 & 9c).P–T evolution of the Lugian domainEarly petrological studies of granulite from the orogeniclower crust suggested peak conditions of about28 kbar at 1000 °C (Bro¨ cker & Klemd, 1996; Kryzaet al., 1996; Klemd & Bro¨ cker, 1999), based on twofeldsparthermometry and Grt–Ky–Qtz–Pl barometry.However, Sˇtípska´ et al. (2004) questioned these resultsbecause of the possibility of non-equilibrium compositions,which may have been involved in the calculations.Instead, these authors proposed that peakpressures were around 18–20 kbar at 800)900 °C (no.14 in Figs 7 & 9d). Sˇtípska´ et al. (2004) further proposedthat these conditions are characteristic for thesteep S 2 foliation. These authors also constrainedconditions to around 10 kbar and 700 °C for theamphibolite facies retrogression within the S 3 fabric(no. 14 in Fig. 9d). Petrological studies and modellingof prograde garnet zoning from rocks of the Ky–St–Grt micaschists revealed a prograde path up to about10 kbar and 650 °C (no. 15 in Figs 7 & 9d; Romanova´&Sˇtı´ pska´ , 2001; Jastrzębski, 2005). A microstructuralstudy confirmed that the growth of prograde garnetwas syntectonic with the steep S 2 fabric. The mainreworking in the horizontal S 3 fabrics occurred in thefield of sillimanite stability at about 7 kbar and 650 °C(no. 15 in Fig. 9d).The Ordovician metamorphic fabric in the leptynoamphiboliteunit developed at about 10 kbar and800 °C (no. 18 in Figs 7 & 9d; Sˇtípska´ et al., 2001;Lexa et al., 2005). Carboniferous metamorphismrelated to the S 3 reworking of rocks at the westernmargin of this unit, close to the Carboniferous sillintruded at around 7 kbar (no. 17 in Figs 7 & 9d; Lexaet al., 2005), occurred under similar conditions ofaround 8 kbar and 750 °C (no. 16 in Figs 7 & 9d;Baratoux et al., 2005).Ó 2007 Blackwell Publishing Ltd147

288 K. SCHULMANN ET AL.(d)(c)PrRe(b)RePrPrPe(a)PePrReRe(e)(f)AAFig. 9. P–T diagrams showing prograde and retrograde P–T paths of samples studied in the structural context discussed in the text.(a) Southern Moldanubian area in Fig. 5: Source of data: 1 – kyanite micaschist (Sˇtı´ pska´ & Schulmann, 1995); 2 – Ky–Kfsgranulite (Racek et al., 2006); 3 – kyanite micaschist (Racek et al., 2006); 4 – staurolite-kyanite paragneiss (Racek et al., 2006);5 – retrogressed eclogite (Racek et al., 2006); 6 – sillimanite migmatite (Racek et al., 2006). (b) Central Moldanubian area in Fig. 5:7 – eclogite (Sˇtı´ pska´ & Powell, 2005a); 8 – Ky–Kfs granulite (P. Sˇtı´ pska´ & K. Schulmann, unpublished data), 9 – migmatitic paragneiss(P. Sˇtı´ pska´ & K. Schulmann, unpublished data); 10 – migmatitic paragneiss (P. Sˇtı´ pska´ & K. Schulmann, unpublished data).(c) Central Moldanubian area in Fig. 6: 11 – Ky–Kfs granulite (Tajcˇmanova´ et al., 2006); 12 – cordierite gneiss (Tajcˇmanová et al.,2006); 13 – metapelite (Pitra & Guiraud, 1996). (d) Lugian domain in Fig. 7: 14 – garnet–omphacite-bearing granulite (Sˇtı´ pskáet al., 2004); 15 – staurolite–kyanite micaschist (Romanova´ &Sˇtı´ pska´ , 2001); 16 – amphibolite (Lexa et al., 2005); 17 – granodiorite(Parry et al., 1997); 18 – Ordovician metapelitic granulite (Lexa et al., 2005). (e, f): Probability curves of age histograms dealingwith existing Moldanubian and Lugian Sm–Nd and 40 Ar ⁄ 39 Ar cooling ages. Bt – biotite 40 Ar ⁄ 39 Ar cooling ages; Grt–Cpx andGrt–whole rock, Sm–Nd ages; Hbl – hornblende and Mu – muscovite, Ar 40 ⁄ Ar 39 cooling ages.Ó 2007 Blackwell Publishing Ltd148

EXHUMATION IN LARGE HOT OROGEN 289Thermochronology of Moldanubian and Lugian domainsThe distribution of isotopic ages reflects a two-stagecooling, related to exhumation processes. An olderstage is marked by cooling ages from minerals withhigh blocking temperatures (Sm–Nd system in garnet,700)750 °C, Hensen & Zhou, 1995;40 Ar ⁄ 39 Ar inhornblende, around 480 °C, Harrison et al., 1985).Systematically, younger cooling ages are retrievedfrom minerals with low blocking temperatures( 40 Ar ⁄ 39 Ar in muscovite, around 350 °C, and in biotite,around 280 °C, Harrison et al., 1985). Peridotitesand associated eclogites yield Sm–Nd whole-rock–garnet ages between 350 and 325 Ma (Brueckner et al.,1991; Beard et al., 1992; Medaris et al., 1995). Thepeak for Sm–Nd ages for both types of rocks is around336 Ma (Fig. 9e). The 40 Ar ⁄ 39 Ar hornblende data arecompatible with the distribution of Sm–Nd ages,whereas the 40 Ar ⁄ 39 Ar muscovite data record systematicallyyounger ages ranging from 331 to 325 Ma(Fig. 9e; Matte et al., 1990; Dallmeyer et al., 1992;Fritz et al., 1996).The Variscan cooling history of the Lugian eclogitesis documented by Sm–Nd garnet–clinopyroxene–whole-rock ages of 340)330 Ma (Brueckner et al.,1991). Recently, Lange et al. (2005) provided Sm–Ndisochrons for the Lugian granulites that define agesbetween 357 and 337 Ma, which fit the existing40 Ar ⁄ 39 Ar hornblende and biotite cooling ages fromthis region (Schneider et al., 2006). The 40 Ar ⁄ 39 Armuscovite ages from all lithologies of the Lugiandomain are similar to Sm–Nd garnet,40 Ar ⁄ 39 Arhornblende and 40 Ar ⁄ 39 Ar biotite data from granulitesand eclogites, indicating a rapid and monocycliccooling history (Fig. 9f).Polyphase fabric and metamorphic evolution of theorogenic beltThe structural pattern in the Lugian domain reveals awell-preserved D 2 vertical fabric in the orogenic lowerand middle crust as well as coherency of these units,defined by sub-parallel alternations of continuous beltsof orogenic lower and middle crust on the geologicalmap. Here, the geometry of the geological units andmap patterns are fully controlled by the D 2 deformation.The D 3 deformation was generally weak andheterogeneous, being mostly concentrated close to thethrust of the Lugian domain over the Ordovician leptyno-amphiboliteunit to the east and close to the areaof flat fabric and the normal-sense shear zone boundingthe Lugian domain in the west (Figs 7 & 8).Normal-sense non-coaxial shearing developed continuouslyfrom the earlier pure shear deformation andprobably this was responsible for displacement ofschists of the orogenic middle crust to the NW andunroofing of the orogenic lower crust.In the orogenic lower and middle crust of the Lugiandomain, petrology indicates that retrograde conditionsof omphacite-bearing granulites and peak pressureconditions of adjacent micaschists are similar. In theorogenic middle crust, the prograde mineral growth isclearly related to the steep S 2 fabric (Romanova´ &Sˇtı´ pska´ , 2001), whereas in the granulites from theorogenic lower crust, development of the S 2 fabric isrelated to their retrogression (Sˇtípska´ et al., 2004). Theremarkable differences in prograde and retrogradeP–T paths of rocks from the orogenic middle andlower crust, respectively, related to steep S 2 foliationsmakes this region the best example of vertical materialand heat transfer during D 2 .Sˇtípska´ et al. (2001) further discussed the significanceof similar Carboniferous and Ordovicianmetamorphic conditions for the easterly lower crustalleptyno-amphibolite complex. They concluded thatthis complex cooled after Ordovician rifting andremained stable in the crust before the onset ofVariscan deformation. Heterogeneous Variscanamphibolite facies deformation was localized in thewestern margin of the Ordovician block and sharesthe same P–T conditions and kinematics withamphibolite facies retrograde fabric of the easternmostorogenic lower crust. Consequently, Sˇtı´ pska´et al. (2004) suggested that the transition from thesteep-to-the-west moderately dipping fabrics in theorogenic lower crust was linked to thrusting of theserocks over a rigid lower crustal block made up ofleptyno-amphibolite unit rocks at a depth equivalentto about 10 kbar (Fig. 5c).In contrast, the horizontal S 3 fabrics affectingrocks of the orogenic lower crust and the normal-senseshear zone developed in the west were linked to thedevelopment of retrograde and syntectonic mineralassemblages indicating an important decrease of temperatureand pressure typical for detachment zones(e.g. Vanderhaeghe & Teyssier, 2001). The Lugian rootrecords a single-stage cooling history characterized bytelescoping of ages for minerals with different blockingtemperatures.The structural development of the NE terminationof the Moldanubian domain shares a number offeatures with the Lugian domain. In this area, the S 2foliation was well preserved in felsic granulites inconjunction with a weak D 3 reworking, whichemphasizes a coherency of the lower and middleorogenic crust and indicates that the map patterndeveloped during D 2 vertical movements, similar tothat in the Lugian domain. The steep fabric in theorogenic lower crust records retrogression to the sillimanitestability field, suggesting that the orogeniclower crust was exhumed to middle crustal conditionsalong the S 2 fabric. The HP characteristics of thisfabric are not necessarily always preserved, preservationbeing dependent mainly on whether the crustavoids re-hydration during exhumation. However, incontrast with the Lugian domain, the heterogeneousD 3 deformation does not show any normal-sensecomponent of shear and is exclusively associated withÓ 2007 Blackwell Publishing Ltd149

290 K. SCHULMANN ET AL.top-to-the-NE-oriented thrusting in the stability fieldof cordierite at low pressures of 3–4 kbar.Structural and petrological studies of the southernand central Moldanubian domains show that similarP–T conditions around 7 kbar at 750 °C are characteristicfor all crustal levels during development of theflat S 3 fabric. Rocks from different depths were mixedand reworked together during the D 3 deformation, asevidenced by contrasting prograde P–T paths withdifferent pressure peaks. The increase of temperaturein the orogenic middle crust and the slight decrease oftemperature in the orogenic lower crust during retrogressionsuggest a mutual thermal equilibration duringdevelopment of the S 3 flat fabric.Structural observations show that all rocks werestrongly deformed during D 3 horizontal flow, althoughmore competent lithologies, mostly middle crust andsome granulites, retained their original steep S 2 fabric,whereas in the less competent, weak lower crust, thehorizontal top-to-the-NE D 3 ductile shearing dominates.During this process, the originally coherentorogenic middle and lower crust was disaggregated toform boudins and rootless folds in a pervasivelyflowing migmatitic matrix. We suggest that providingthe volume of orogenic middle crustal is high enough,the mineral assemblages from the D 2 vertical fabricsare preserved (Figs 6 & 9a). In contrast, the smallerboudins were completely re-equilibrated and evenheated in the flowing mass of hot orogenic lower crust(Figs 6 & 9b). The metamorphic conditions associatedwith the D 3 flow show clearly that pressure, temperatureand intensity of D 3 reworking decreases fromsouth to north across the whole continental margin.Based on dating of metamorphic zircon from granulites,the timing of HP metamorphism and S 2 fabricformation in the Moldanubian domain was estimatedby Schulmann et al. (2005) and Tajcˇmanova´ et al.(2006) to occur between 350 and 340 Ma. This agespan corresponds with cooling ages of minerals withhigh blocking temperatures in the northern Moldanubiandomain (Matte et al., 1990; Macintyre et al.,1992), and indicates that already during the D 2 stagean important cooling and exhumation of this area wastaking place, similar to the Lugian domain. Theyounger 330–325 Ma 40 Ar ⁄ 39 Ar cooling ages obtainedfor muscovite and biotite imply a second distinct periodof thermal reworking and isotopic resetting in theregion, correlated with the strong D 3 reworking in thesouth.Geophysical imagery of the subsurface shape of the Bruniacontinental promontoryIt is apparent that the thrust-related horizontal flow,which dominates the deformation of the Moldanubiandomain at around 325 Ma, does not exist in the Lugiandomain. This indicates a significant difference in thebulk exhumation processes between the two crustalsegments and the important involvement of the Bruniacontinent in the Moldanubian exhumation history. Inorder to discuss the relative contribution of the Bruniabasement in the tectonic evolution of the orogenic rootit is necessary to know the sub-surface extent of thebasement promontory underneath the orogenic rootrocks. Therefore, after the compilation of gravity data,a Bouguer anomaly map of the area (Fig. 10) wasproduced.At a large scale, the Bouguer anomaly map (Fig. 10)shows two main domains: (i) to the west, a domaincharacterized by low- and intermediate-gravity anomalies,inferred to be associated with rocks that have lowto intermediate densities; and (ii) to the east, a domainwith a succession of gravity highs inferred to be associatedwith significantly denser rocks. In the NE, thelow- to intermediate-gravity anomalies coincide withthe Lugian domain, whereas the gravity highs areassociated with the Brunia continent. The steep horizontalgradient in gravity reflects a steep boundarybetween the Lugian domain and the Brunia continent.In the SW, the Moldanubian domain west of theNNE–SSW-striking Central Moldanubian pluton ischaracterized by low- and intermediate-gravity anomaliessimilar to the Lugian domain. Similar to thenorth, we infer that the gravity highs over the Bruniacontinent in the SE represent mostly dense rocks.However, these gravity highs continue to the west, intothe eastern part of Moldanubian domain as far as theCentral Moldanubian pluton. Therefore, the observationmade in the north, where the gravity boundarycoincides with the geological boundary between theLugian domain and Brunia, is not valid in the south,where the geophysical boundary is located about 50–70 km west of the mapped geological boundary.Within the eastern Moldanubian domain, the rocksat outcrop are similar west and east of the SouthBohemian pluton, and rocks with high densities thatcould explain the gravity high in the eastern part of theMoldanubian domain are not represented at outcrop.Consequently, the dense rocks must be located at agreater depth. The gentle gradient of the gravity highsin the west indicates that the dense Brunia continentalpromontory dips towards the west, beneath the easternMoldanubian rocks. This interpretation is confirmedby preliminary 3D modelling, which shows that theboundary between the dense Brunia basement and thelow- to intermediate-density rocks of the Moldanubiandomain has a gentle dip down to a depth of some 2–3 km near the Central Moldanubian pluton, where itdips more steeply beneath the pluton (60–70°).Implication of Bouguer anomaly patternThis analysis of the Bouguer anomaly map in Fig. 10reveals striking differences between the Moldanubianand Lugian domains, which may be interpreted interms of the presence of Brunia basement under a thinlayer of Moldanubian rocks. This interpretation isconsistent with the suggestion that the S 3 fabric in theÓ 2007 Blackwell Publishing Ltd150

EXHUMATION IN LARGE HOT OROGEN 291OrdovicianBruniaMoldanubianBruniaLugian mafic lower SilesianCulm basincrust(a) A′ (b) B′CMPOMCOLCOLCCretaceous(c) basin C′CMPMg-K syeniteOMCThaya domeA-A′Southern Central NorthernBruniaMoldanubianMg-K syeniteOMCOLCOLCOMCUpper crustFig. 10. Interpretative cross-sections located in Fig. 2, showing distribution of the orogenic lower, middle and upper crust of theorogenic root and the Brunia continent. (a) East to west cross-section of southern Moldanubian showing synformal structure. TheS 3 fabric associated with channel flow is dominant in the orogenic lower crust, whereas in the orogenic middle crust the steep S 2fabric is dominant. Dismembered orogenic middle crust forms large-scale boudins surrounded by migmatitic S 3 fabric. (b) Crosssectionof Lugian–Silesian domain, to show relationships between the D 2 extrusion of the orogenic lower crust that was thrust overthe Ordovician leptyno-amphibolite complex and the S 3 fabric associated with ductile thinning in the Lugian section. (c) North tosouth cross-section of the Moldanubian domain. Figure shows strong D 3 deformation of the rear part of the system with importantmixing of orogenic lower crust and orogenic middle crust in the channel flow fabric and weak D 3 deformation in the northernpart of the section.Moldanubian domain may result from underthrustingof the Brunia continental promontory. In contrast, theabsence of a gravity high over the Lugian domain,coupled with the structural and geochronological criteriadiscussed above, indicate that the Brunia continentwas not underthrust beneath the orogenic root ofthe Lugian domain. In other words, the horizontalfabrics in the Lugian domain cannot be explained bythe mechanical interaction between the Brunia continentand the Lugian root system.DISCUSSIONTectonic significance of steep orogenic fabricsThe zone of vertically foliated Lugian domain granulitesthat are surrounded by migmatized and highlysheared orthogneisses that also preserve a steep metamorphicS 2 fabric represents a key area for understandingthe early stage of exhumation of the lowerorogenic crust. Thrusting or normal faulting exhumationmechanisms cannot explain the steep S 2 fabric inthe granulite and orthogneisses, which we have interpretedinstead as recording a vertical channel alongwhich upward extrusion of the orogenic lower crustoccurred.Another important observation is that the S 2 foliationresults from macroscopic refolding of an earlysub-horizontal fabric at the scale of both the mesoscopicsub-horizontal compositional anisotropy andthe orogenic crust (Sˇtípska´ et al., 2004; Racek et al.,2006). In addition, petrological studies show that theS 2 fabric in rocks of the orogenic middle crust adjacentto the granulites is associated with the progrademetamorphic evolution to a metamorphic peak ataround 9 kbar, whereas in the orogenic lower crustthe S 2 fabric is related with retrogression from highpressures around 18–20 kbar to pressures around7–10 kbar.The structural pattern and opposite metamorphicevolution developed in the kinematically similar S 2fabrics indicates vertical material transfer compatiblewith large-scale folding. We suggest that folding of thecrustal layers was responsible for the exhumation ofthe orogenic lower crust in cores of large antiforms andburial of the orogenic middle and upper crust in synformalregions. Furthermore, we suggest that thealternations of vertical belts of orogenic lower andmiddle crust that parallel the continental marginoriginated during this D 2 event. At the end of thisprocess most of the orogenic middle and lower crusthad already cooled below the blocking temperaturesfor the 40 Ar ⁄ 39 Ar system in hornblende and partly alsothe 40 Ar ⁄ 39 Ar system in muscovite.In the Polish part of the Lugian domain large-scalefolding of crustal layers was recognized by Don (1964)and Dumicz (1979). Moreover, a theoretical backgroundfor exhumation of lower crust by folding hasbeen provided by (Burg & Podladchikov, 1999) andapplied to the exhumation of lower crustal rocks forÓ 2007 Blackwell Publishing Ltd151

292 K. SCHULMANN ET AL.example in the area of the Namche Barwa syntaxis(Burg et al., 1997).Based on shape analysis of macroscopic folds, Franeˇket al. (2006) proposed that the folding of lowercrustal layers occurred at granulite facies conditions bypassive amplification mechanisms. Consequently, theseauthors suggested that the folding started by thedevelopment of cuspate structures at the boundarybetween the lower and middle crust because of a lowviscositycontrast at this interface (cf. Kisters et al.,1996). Growth of the upward-pointing cusps wasreplaced by vertical extrusion as the whole domainbecame sufficiently shortened and the horizontalanisotropy was replaced by a vertical anisotropy. Atthat stage, the difference in the integrated vertical bulkviscosity between a weak lower crustal cusp and anadjacent stronger middle crustal lobe increased. Theseviscosity differences lead to strain partitioning andvertical extrusion of the weak orogenic lower crust butslower exhumation of adjacent middle crust (Jezˇeket al., 1998).To make the mechanism of vertical extrusion possiblea rigid floor is required (Schulmann et al., 2003),which is represented by a strong sub-root mantle in themodel proposed here. The strength of the sub-crustalmantle was modelled for a given thermal gradient byThompson et al. (2001) and Schulmann et al. (2002),who showed that at least 40 km of rigid mantle lithospherewas likely to have existed at the onset ofexhumation of the thermally weakened lower crust.The occurrence of slivers of mantle peridotite, whichhave sharp boundaries and discordant internal fabricswith respect to the flow fabric of the surroundinggranulites, within the orogenic lower crust providesfurther evidence for the existence of a strong mantleand implies an important strength contrast between themantle and the orogenic lower crust (Medaris et al.,2006). Although the granulite extrusion structuresdescribed in this study resemble extrusions of lowercrustdriven by density inversion (Martinez et al.,2001), they were controlled mostly by lateral shorteningforces in this case, as shown in Fig. 11a.Horizontal flow of orogenic lower crustA transition from vertical S 2 fabrics to horizontal S 3fabrics occurs in almost all examples of lower crustalvertical structures in the Moldanubian and Lugiandomains. Therefore, the main questions that arise are:is there a causal relationship between vertical extrusionand the development of horizontal flow? and what arethe driving forces for the deformation?Ductile thinning and collapse of vertical fabric in the LugiandomainA satisfactory tectonic model to explain the structural,petrological and geochronological data from theLugian domain is one involving folding of a layeredFig. 11. Bouguer anomaly map of the eastern margin of theBohemian Massif (provided by the Czech Geological Survey).Thick lines superimposed on the Bouguer anomaly map arelimits of geological unit boundaries and orogenic crustal levelsfrom Fig. 2. Br, Brunia continent; MDE, eastern branch of theMoldanubian domain; CMP, the Central Moldanubian pluton;MDW, western branch of the Moldanubian domain; LD, theLugian domain.orogenic crust followed by the asymmetrical northeastwardextrusion of the orogenic lower crust causedby the indentation of an Ordovician mafic lowercrustal block at c. 340 Ma. Extrusion of the orogeniclower crust was associated with the development of thehorizontal D 3 fabric accompanied with detachment ofthe western units and complete reworking of the orogenicmiddle crust in adjacent synforms. Consequently,the horizontal fabric in the Lugian domain cannothave been created in response to intracrustalflow because of lateral variations in lithostaticpressure ⁄ gravitational potential energy of thickenedcrust ⁄ lithosphere (Milnes & Koyi, 2000; Vanderhaeghe& Teyssier, 2001).A more likely explanation is a model in which thevertically moving material experiences a reversal in theprincipal strain-rate directions (Feehan & Brandon,1999), which was expressed as a switch from the verticalfabric at depth to the sub-horizontal fabric atshallow crustal levels (Ring & Brandon, 1999). However,such a model requires the presence of a thickcontinental accretionary wedge, as proposed by Platt(1986), which has a mixed flow field involving verticalthickening at depth and vertical thinning near thesurface.Previously Schulmann & Gayer (2000) interpretedthe Lugian domain as an obliquely convergent wedgedeveloped above the Saxothuringian subduction zone.In this model, the Ordovician mafic lower crustal blockÓ 2007 Blackwell Publishing Ltd152

EXHUMATION IN LARGE HOT OROGEN 293represents a relict of the rigid backstop and the normal-senseshear zone in the west reflects the collapse ofvertically extruded material (Fig. 11b). Importantly,the U–Pb zircon data show Ordovician protolith agesfor both the felsic orogenic root (orogenic lower andmiddle crust) and the Ordovician mafic lower crustalblock to the east. These data, together with the coincidenceof the gravity and geological boundariesbetween the Lugian domain and the Brunia continent(Fig. 10), indicate that the transition from vertical tohorizontal fabric results from an intra-Lugianmechanical interaction and that the Brunia continentwas not involved in this process (Fig. 12).Heterogeneous channel flow and hot fold nappe fabricsin the Moldanubian domainThere are a number of differences associated with thedevelopment of the flat fabric in the Moldanubiandomain compared with the development of the flatfabric in the Lugian domain. In the southern part ofthe Moldanubian domain, the horizontal S 3 fabricdeveloped at decreasing pressure from south to north(Figs 4 & 5d) in conjunction with decreasing intensityof the D 3 reworking. This is associated with an increasein temperature and decrease in pressure in the orogenicmiddle crust concomitant with a decrease in temperatureand pressure of the orogenic lower crust. Inaddition, migmatized gneisses of the orogenic lowercrust commonly surround blocks and boudins of theorogenic middle crust that preserve evidence of theearly HP metamorphism. The D 3 deformation showsconstant NNE–SSW-oriented stretching and top-tothe-NNEthrust-related shear movement, which isconsistent with oblique transpressive deformation inthe adjacent Moravo–Silesian Zone. The commongeometry and oblique thrust kinematics of the D 3deformation are the chief features of the mechanicalinteraction between the Moldanubian domain, Moravo–SilesianZone and the Brunia continent. Thegeophysical observations confirm the hypothesis oflarge-scale displacement of the Brunia basementunderneath the Moldanubian domain, which interpretationlocates the sub-surface margin of the Bruniacontinental promontory far to the west from thepresent Moldanubian domain and Moravo–SilesianZone boundary.This highly non-coaxial deformation correlates withyoung cooling ages (330–325 Ma) for muscovite andbiotite 40 Ar ⁄ 39 Ar systems compared with older (350–340 Ma) hornblende 40 Ar ⁄ 39 Ar data, Sm–Nd coolingages from HP metamorphic rocks and metamorphicages from zircon, preserved mainly in the north.Muscovite and biotite from the adjacent Moravo–Silesian Zone nappes and the deformed Brunia continentshow similar young 40 Ar ⁄ 39 Ar cooling ages,which correspond to those of detrital muscovite fractionsfrom the Culm foreland basin. The detritalmuscovite shows convergence of the stratigraphic ageof sedimentation and the 40 Ar ⁄ 39 Ar cooling ages at330–325 Ma (F. Neubauer, pers. comm.). Pebbles ofgranulite and Mg-rich syenite (durbachite) in theforeland conglomerates have yielded similar informationfrom U–Pb dating of zircon, where two groups ofages occur at c. 340 and c. 325 Ma, the latter beingconsistent with the stratigraphic age of 330–320 Ma(Kotkova´ et al., 2007).We suggest that the D 3 deformation, which involvedsubhorizontal top-to-the-NE thrusting of the Moldanubiandomain over the Brunia continent, occurred(a)(c)Fig. 12. Sequence of block diagrams toshow the principal exhumation mechanismsand progressive tectonic evolution of theeastern margin of the Bohemia Massif. (a)Early stage of crustal folding and extrusionassociated with vertical material and heattransfer. (b) Second stage (c. 340 Ma) ofdevelopment of a flat fabric due to collapseof the vertical anisotropy – probably due toductile thinning mechanisms that accompanydetachments of the upper crust. (c)Weak orogenic root deformed by the Bruniacontinental promontory and the developmentof large hot nappes during D 3 at 330–325 Ma. The section shows a northwardattenuation of the crust together with anorthward-facing topographic slope withcoeval surface erosion in the North. Thelatter is consistent with the geometry of theBrunia continent.(b)Ó 2007 Blackwell Publishing Ltd153

294 K. SCHULMANN ET AL.at 330–325 Ma and was associated with exhumation ofHP metamorphic rocks to the surface, as demonstratedby sedimentation of granulite pebbles and metamorphicmuscovite in the foreland basin. The earlier agesclustering around 340 Ma correspond to D 2 event, i.e.to the vertical material transfer associated with crustalscalefolding, evidence of which is partially preservedin flat-lying migmatites.The structure of the Moldanubian domain is consistentwith the Ôstage 3Õ of the channel-flow model – thehot fold nappe (Beaumont et al., 2006; Culshaw et al.,2006). In these channel-flow models, stage 3 coincideswith the arrival of a lower crustal block that forces weakmiddle and lower crust into large-scale gently inclinedfold nappes rooted in the thickened Moho.The hot fold nappe model can be applied to theMoldanubian root system, keeping in mind that themodel represents the result of a continuous 2D historywhereas the Moldanubian root system comprises anexhumation history in two stages that are kinematicallyindependent. Nevertheless, the hot fold nappemodel simulates well the relatively weakly deformedNE part of the Moldanubian system with well-preservedD 2 fabric, which reaches shallow crustal levelsand ultimately the surface (Figs 8 & 11). To the south,the extruded nappes were derived from more internaland hotter parts of the orogen. Increasingly largevolumes of weak lower crust were gradually transportedsouthwards over the Brunia continent leadingto an increase in the amount of lower crustal materialat the surface.The channel flow is heterogeneous, as predicted byBeaumont et al. (2001, 2006), because of the previoushistory, which generates important crustal strengthvariations. The implication of this model is that theheterogeneous crust makes the geometry and compositionof the channel flow similarly heterogeneous. Inthe Moldanubian case, the channel transportsdetached blocks and boudins of distinctly differentcompositions such as HP granulites and mid-crustalsegments that still preserve relicts of the D 2 fabric. Inthis southern part of the channel all isotopic systemsare re-equilibrated and the cooling ages are younger incomparison with the non-uniformly reset isotopicsystems in the north.CONCLUSIONSThis study has shown that exhumation of the orogeniclower crust in the eastern sector of the Variscan orogenicbelt is characterized by two independent stages.(1) The first stage is best preserved in the Lugiandomain, where it is characterized by an intra-crustalfolding that is responsible for vertical material transferassociated with exhumation of the deep orogenic lowercrust to shallower crustal levels corresponding topressures around 10 kbar at about 350 to 340 Ma. Thefolding and vertical extrusion events are followed by avertical shortening leading to development of subhorizontalfabrics at medium to low pressures. Theearly horizontal shortening was probably triggered bythe existence of a rigid Ordovician block, part of whichis preserved at the eastern boundary of the Lugiandomain. We suggest that this early exhumation eventwas related kinematically to Saxothuringian continentalsubduction to the east, creating a convergentcontinental accretionary wedge – the Lugian domain.Mechanical interactions with the large Brunia continentduring the exhumation process remain unconstrained.(2) The Moldanubian domain has steep fabrics insimilar orientations to those preserved in the Lugiandomain. Vertical material transfer during the LowerCarboniferous led to the development of alternationsof steeply inclined domains of lower and middle orogeniccrust, similar to Lugian domain. However, thistectonic event was followed by a NNE-directed subhorizontalshearing at about 330–325 Ma resultingfrom subsurface indentation by the Brunia continentalpromontory into the Moldanubian domain. The arrivalof the Brunia continent is responsible for the progressiveemplacement of hot fold nappes andheterogeneous channel flow in the rear (western) partof the system generating mixtures of middle and lowercrustal units with preserved early exhumation fabrics.ACKNOWLEDGEMENTSThis study was made possible thanks to the ANRproject ÔLFO in orogensÕ funding as well as to financialsupport of CNRS (UMRs 7516 and 7517) and theCzech Science Foundation (GACR 205 ⁄ 05 ⁄ 2187 andGACR 205 ⁄ 04 ⁄ 2065). Ondrej Lexa is indebted to theUniversite´ Louis Pasteur for covering his salary. Weare grateful to J. Platt, D. Grujic and an anonymousreviewer for constructive reviews. The editorial workof M. Brown is highly appreciated.REFERENCESAllemand, P. & Lardeaux, J. M., 1997. Strain partitioning andmetamorphism in a deformable orogenic wedge: Applicationto the Alpine belt. Tectonophysics, 280, 157–169.Andersen, T. B., Jamtveit, B., Dewey, J. F. & Swensson, E.,1991. 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Author's personal copyC. R. Geoscience 341 (2009) 266–286TectonicsAn Andean type Palaeozoic convergence in the Bohemian MassifKarel Schulmann a, *, Jiří Konopásek b,c , Vojtĕch Janoušek b , Ondrej Lexa c ,Jean-Marc Lardeaux d , Jean-Bernard Edel a , Pavla Štípská a , Stanislav Ulrich ea EOST, UMR 7517, université Louis-Pasteur, 1, rue Blessig, 67084 Strasbourg, Franceb Czech Geological Survey, Klárov 3, 118 21 Praha 1, Czech Republicc Institute of Petrology and Structural geology, Charles University, Albertov 6, 128 43 Prague, Czech Republicd Géosciences Azur, UMR CNRS 6526, université de Nice Sophia-Antipolis, parc Valrose, 06108 Nice cedex 02, Francee Geophysical Institute, Czech Academy of Sciences, Boční II/1401, 141 31 Praha 4, Czech RepublicReceived 24 December 2008; accepted after revision 24 December 2008Available online 25 February 2009Written on invitation of the Editorial BoardAbstractThe geological inventory of the Variscan Bohemian Massif can be summarized as a result of Early Devonian subduction of theSaxothuringian ocean of unknown size underneath the eastern continental plate represented by the present-day Teplá-Barrandian andMoldanubian domains. During mid-Devonian, the Saxothuringian passive margin sequences and relics of Ordovician oceanic crusthave been obducted over the Saxothuringian basement in conjunction with extrusion of the Teplá-Barrandian middle crust along the socalledTeplá suture zone. This event was connected with the development of the magmatic arc further east, together with a fore-arc basinon the Teplá-Barrandian crust. The back-arc region – the future Moldanubian zone – was affected by lithospheric thinning whichmarginally affected also the eastern Brunia continental crust. The subduction stage was followed by a collisional event caused by thearrival of the Saxothuringian continental crust that was associated with crustal thickening and the development of the orogenic rootsystem in the magmatic arc and back-arc region of the orogen. The thickening was associated with depression of the Moho and the fluxof the Saxothuringian felsic crust into the root area. Originally subhorizontal anisotropy in the root zone was subsequently folded bycrustal-scale cusp folds in front of the Brunia backstop. During the Visean, the Brunia continent indented the thickened crustal root,resulting in the root’s massive shortening causing vertical extrusion of the orogenic lower crust, which changed to a horizontal viscouschannel flow of extruded lower crustal material in the mid- to supra-crustal levels. Hot orogenic lower crustal rocks were extruded: (1)in a narrow channel parallel to the former Teplá suture surface; (2) in the central part of the root zone in the form of large scale antiformalstructure; and (3) in form of hot fold nappe over the Brunia promontory, where it produced Barrovian metamorphism and subsequentimbrications of its upper part. The extruded deeper parts of the orogenic root reached the surface, which soon thereafter resulted in thesedimentation of lower-crustal rocks pebbles in the thick foreland Culm basin on the stable part of the Brunia continent. Finally, duringthe Westfalian, the foreland Culm wedge was involved into imbricated nappe stack together with basement and orogenic channel flownappes. To cite this article: K. Schulmann et al., C. R. Geoscience 341 (2009).# 2009 Published by Elsevier Masson SAS on behalf of Académie des sciences.RésuméConvergence paléozoïque de type Andin dans le Massif de Bohême. Le Massif varisque de Bohême est le résultat de lasubduction, au Dévonien supérieur, de l’océan Saxothuringien sous la plaque continentale représentée à l’est par les zones actuelles* Corresponding author.E-mail address: schulmann.karel@gmail.com (K. Schulmann).1631-0713/$ – see front matter # 2009 Published by Elsevier Masson SAS on behalf of Académie des sciences.doi:10.1016/j.crte.2008.12.006159

Author's personal copyK. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286 267de Teplá-Barrandien et de Moldanubien. L’ampleur de cet océan demeure inconnue. Pendant le Dévonien moyen, les sériessédimentaires de la marge passive saxothuringienne et les reliques de la croûte océanique ordovicienne ont été obductées sur le soclesaxothuringien alors que, dans le même temps, la croûte moyenne Teplá-barrandienne était extrudée le long de la suture de Teplá.Plus à l’est, cet événement était associé au développement d’un arc magmatique et à la formation d’un bassin avant-arc sur le socleTeplá-Barrandien. Le domaine arrière-arc – la future zone moldanubienne – subissait un amincissement lithosphérique qui affectaitaussi en partie la croûte continentale de Brunia. La subduction s’est achevée avec la collision de la croûte continentalesaxothuringienne et s’est traduite par un épaississement crustal et la formation d’une racine orogénique dans les zones del’arc magmatique et de l’arrière-arc. L’épaississement était associé à une dépression du Moho et au flux de croûte felsiquesaxothuringienne dans la racine orogénique. L’anisotropie sub-horizontale de la racine a donné suite à un plissement serré, d’échellecrustale, au front du butoir continental de Brunia. Pendant le Viséen, le continent Brunia a indenté la racine crustale épaissie enprovoquant un important raccourcissement de la racine, l’extrusion verticale de la croûte inférieure de l’orogène passant à un fluxvisqueux en chenal horizontal, de matériaux de la croûte inférieure à des niveaux médio- et supra-crustaux. Les roches de la croûteinférieure orogénique ont été extrudées (1) dans d’étroits chenaux parallèles à la surface de l’ancienne suture de Teplá, (2) au centrede la zone de racine sous la forme d’un large antiforme et (3) en nappe plissée chaude au-dessus du promontoire de Brunia, tout enproduisant un métamorphisme barrovien, ainsi que l’imbrication des niveaux supérieurs. Les roches extrudées les plus profondes dela racine ont atteint la surface, puis ont été reprises, peu de temps après, sous la forme de galets dans la sédimentation de type Culmdans l’épais bassin d’avant-pays qui recouvre la partie stable du continent Brunia. Enfin, pendant le Westphalien, le prismed’accrétion du Culm, ainsi que le socle et les nappes de chenaux crustaux ont été repris dans une suite de nappes imbriquées. Pourciter cet article : K. Schulmann et al., C. R. Geoscience 341 (2009).# 2009 Publié par Elsevier Masson SAS pour l’Académie des sciences.Keywords: Bohemian Massif; Saxothuringian oceanic subduction; Building of Variscan orogenic root system; Channel flowMots clés : Massif de Bohême ; Subduction océanique saxothuringienne ; Formation de la racine orogénique varisque ; Flux en chenal1. IntroductionStarting with [16] and followed by [33] and [93,94],the eastern Variscan belt is interpreted as the result ofDevonian to Carboniferous continent–continent collision,resembling the (sub-)recent Himalayan-Tibetantype collisional system. However, there is a wealth ofdata (see for example [93,94]) suggesting that thisorogenic belt could have resulted from an Andean typeconvergence, i.e. as a typical upper plate orogen locatedabove a long-lasting Devonian–Carboniferous subductionsystem.The aim of this article is to show that all the currentcriteria defining an Andean type of convergent marginare present and surprisingly well preserved in theVariscan Bohemian Massif. These criteria are inparticular [66,71,122,157]: the development of Franciscan type blueschist-faciesmetamorphism in the lower plate; arc type magmatism marked by calc-alkaline topotassium-rich (shoshonitic) series in the distance of150–200 km from the trench; back-arc basin developed on continental upper platecrust and replaced by thick continental root; deep granulite-facies metamorphism associated withsupposed underplating of the crust by mafic magmasat the bottom of the root; continental lithosphere thrust underneath the thickenedroot system.Based on these criteria, the architecture of theeastern Variscan belt is interpreted as the result of alarge-scale and long-lasting subduction process associatedwith crustal tectonics, metamorphism, magmaticand sedimentary additions that developed over thewidth of at least 500 km, in present-day coordinates,and time scale of 80 Ma.2. Present-day architecture of the BohemianMassif and location of Palaeozoic suturesThe Bohemian Massif occurs at the eastern extremityof the European Variscan belt, representing one of itslargest exposures (Fig. 1). From west to east, the EasternVariscan belt forming the Bohemian Massif can besubdivided into four major units: the Saxothuringian Neoproterozoic basement with itsPalaeozoic cover corresponding to the continentalcrust of the Armorican plate [87,105,147]; the Teplá-Barrandian Unit consisting of Neoproterozoicbasement and its Early Palaeozoic coverinterpreted as an independent crustal block (theBohemia Terrane of South Armorica sensu [34]); the Moldanubian high- to medium-grade metamorphicdomain intruded by numerous Carboniferous160

Author's personal copy268K. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286Fig. 1. Simplified geological map of the Bohemian Massif (modified after [34]). CBPC: Central Bohemian Plutonic Complex; CMP: CentralMoldanubian Pluton. The lower left insert shows position of the Bohemian Massif in the frame of the European Variscides (modified [22]).RH: Rhenohercynian zone; ST: Saxothuringian Zone; M: Moldanubian Zone; B: Brunia Continent; L: Lugian domain.Fig. 1. Carte géologique simplifiée du Massif de Bohême (modifiée d’après [34]). CBPC : Complexe plutonique de Bohême centrale ; CMP : Plutoncentral moldanubien. L’encart du bas à gauche montre la position du Massif de Bohême dans le cadre des Variscides européennes (modifié d’après[22]). RH : Zone rhénohercynienne ; ST : Zone saxothuringienne ; M : Zone moldanubienne ; B : Brunia ; L : Zone lugienne.granitic plutons, altogether forming the high-gradecore of the orogeny; the eastern Brunia Neo-Proterozoic basement withEarly to Late Palaeozoic cover [35,140].The palaeontological record of Lower Palaeozoic(Cambrian and Ordovician) sediments of the Saxothuringianand Teplá-Barrandian domains shows affinitiesto the Gondwana faunas implying that these blockswere derived from the northern Gondwana margin[20,27,28,68]. In addition, there is a range of isotopicand U–Pb zircon data suggesting a Gondwananprovenance of all units composing the BohemianMassif [96,106].2.1. Saxothuringian domainThis domain is represented by Neoproterozoic parautochthonousrocks (580–550 Ma) formed by migmatitesand paragneisses intruded by Cambro-Ordoviciancalc-alkaline porphyritic granitoids converted to augenorthogneiss during the Variscan orogeny [38,39,76,89].These rocks are unconformably covered by Cambrianand Ordovician rift sequences (Fig. 1) overlain by LateOrdovician to Famennian pelagic sediments andFamennian to Visean flysh of the Thuringian facies[85,125,144]. The par-autochthon is overthrust byallochthonous units containing deep-water equivalentsof the Ordovician to Devonian rocks of the paraautochthon,and by proximal flysh sediments (theBavarian facies).The allochthonous units occur in Münchberg,Wildenfels and Frankenberg klippens and exhibit pileof thrust sheets marked by decreasing pressure andmetamorphic age from the top to the bottom [34]. In thehangingwall occur thrust sheets with Mid-Ocean RidgeBasalt (MORB)-type metabasites of Ordovician protolithage eclogitized (Fig. 2; P = 25 kbar, T = 650 8C)161

Author's personal copyK. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286 269[34] during the Devonian 395 Ma [102]. Structurallydeeper occur sheets marked by medium pressure (MP)assemblages and Late Devonian (365 Ma) zircon andhornblende cooling ages. This rock pile represents bothdistal and proximal Late Ordovician to Devonianpassive margin rocks tectonically inverted during theDevonian convergence [44,134]. In the Sudetic part(Figs. 1 and 2) of the Bohemian Massif, the Ordovicianrift sequences are well developed and marked by thepresence of deep marine sediments and MORB-typevolcanics followed by Silurian and Devonian sedimentarysequences [95]. The Ordovician oceanic rocks areenhanced by a blueschist-facies metamorphic overprintof Late Devonian age [33,34].However, a later Carboniferous underthrusting ofthe Saxothuringian/Armorican continental rocksunderneath the easterly Teplá-Barrandian block wasidentified suggesting a continuous convergence in thisarea, which was responsible for eclogitization of boththe oceanic and continental crust resulting fromcontinental underthrusting [90] at 340 Ma [71].The continental underthrusting reached even thepressure conditions of the diamond stability [150].This event is responsible for the global reworking of theSaxothuringian terrane at HP conditions, imbricationof subducted continental crust and exhumation of deeprocks in form of crustal scale nappes [92]. Thus,thestructure of the Saxothurinigian domain is defined bypar-autochthonous domain (Fig. 2; P =13–15 kbar,T =580–630 8C) [69,70,72] and eclogite-bearing‘‘lower’’ crustal nappe (P =20–26 kbar, T =630–700 8C) [70]. The highest tectonic unit, ‘‘the uppernappe’’ reached the high pressure (HP) granulite faciespeak conditions (P =15–20 kbar, T =8008C [69,116])at 340 Ma [79,154]. The exhumation of granulite faciesunit occurred at 340 Ma as shown by 40 Ar/ 39 Ar coolingages [80,83].2.2. The Saxothuringian–Teplá-BarrandianboundaryThe boundary between the two crustal domains ischaracterized by the presence of units with highproportion of ultramafic and mafic rocks (Fig. 1;Mariánské Lázně Complex and Erbendorf–VohenstraussZone). The former complex is marked by a presence ofserpentinites at the bottom overlain by a thick sequenceof amphibolites, eclogites and metagabbros. The U–Pbzircon protolith ages discriminate the mafic rocks intotwo main groups with Cambrian (540 Ma [145]) andOrdovician (496 Ma [5]) ages. On the other hand, themetamorphic and cooling ages (Sm–Nd garnet–pyroxeneand Ar–Ar hornblende) are Devonian, ranging between410 and 370 Ma [4,15]). Similarly, the newly determinedU–Pb ages for metamorphic zircon in mafic rocks, as wellas monazite from orthogneiss, and titanite in leucosomeall cluster around 380 Ma and possibly date theexhumation of the whole Mariánské Lázně Complex[163]. The Devonian metamorphic evolution started witheclogite-facies metamorphism (Fig. 2; P =16–18 kbar,T = 640–715 8C) [145] and continued at c. 380 Ma bygranulite-facies re-equilibration [99]. The Erbendorf–Vohenstrauss Zone further west shows similar lithologicaland petrological zonation as the Mariánské LázněComplex and it is thus commonly interpreted as a part ofthe same tectonic unit [101,145].2.3. Teplá-Barrandian domainStratigraphically, the Teplá-Barrandian Unit (Fig. 1)consists of Neoproterozoic basement with the lower arcrelatedvolcano-sedimentary sequence (the Kralupy–Zbraslav Group), followed by siliceous black shales anda flyshoid sequence (shales, greywackes and conglomerates,[35]). The Neoproterozoic basement is unconformablyoverlain by a thick sequence (1500–2000 m)of Lower Cambrian conglomerates, greywackes, andsandstones [19,82] followed by shales and a rift-relatedvolcanic sequence in the Upper Cambrian. Thedevelopment of the Lower Palaeozoic Prague Basin[17] is marked by Early Ordovician (Tremadocian)transgression followed by mid-Ordovician riftingassociated with volcanic activity, and with sedimentationof Silurian graptolite shales. The sedimentationcontinued mainly with carbonates, namely the UpperSilurian to Devonian calcareous flyshoid sequence. TheEarly Silurian was associated with important volcanicactivity, accompanied by basaltic and ultramaficintrusions [51]. The sedimentation terminated in mid-Devonian with distal turbidites of the Srbsko Formation(Givetian) containing, among others, detrital zircons ofDevonian (390 Ma) age [108].The whole sequence is folded by steep foldspresumably of Late Devonian age as shown by Culmfacies sediments unconformably deposited on foldedOrdovician strata [135]. The deformation affected alsothe underlying Neoproterozoic basement, havingintensity and age increasing progressively to the west.In the same direction rises also the metamorphic gradereaching amphibolite-facies conditions close to theTeplá-Barrandian/Saxothuringian boundary [10]. In thisarea is developed a typical Barrovian metamorphiczonation ranging from biotite–garnet zone (Fig. 2a;P =3–4 kbar, T = 450–520 8C) [146,156,160,161] in162

Author's personal copy270K. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286Fig. 2. Pressure–temperature (P–T) diagram showing the location of the principal metamorphic facies in the P–T space (after Brown, [6]). BS:blueschist facies; AEE: amphibole-epidote eclogite facies; ALE: amphibole-lawsonite eclogite facies; LE: lawsonite eclogite facies; AE: amphiboleeclogite facies; GS: greenschist facies; A: amphibolite facies; E-HPG: medium-temperature eclogite facies – high-pressure granulite metamorphism;G: granulite facies; UHTM: the ultra-high-temperature metamorphic part of the granulite facies. a: Devonian HP–LT metamorphism (390–380 Ma) from the Münchberg and Mariánské Láznĕ units contrasted to Devonian (380 Ma) MP metamorphism of the western part of the Teplá-Barrandian; b: Carboniferous (340 Ma) HP–LT metamorphism from the Saxothuringian par-autochthon and lower nappe contrasted to the HP–HTmetamorphism of the granulite bearing upper nappe (modified after [71]); c: Carboniferous (340 Ma) HP–HT metamorphism of the orogenic lowercrustal belt from the Moldanubian domain, HT–MP and LP metamorphism related to the exhumation of the orogenic lower crust contrasted to theMP–MT peak metamorphism of the orogenic middle crust followed by heating during exhumation (modified after [123]; d: Carboniferous HP–MTmetamorphism of the Moravian eclogites contrasted to the MP–MT Barrovian metamorphism of the Moravian Zone (modified after [77]).Fig. 2. Diagramme pression–temperature (P–T) montrant la localisation des principaux faciès métamorphiques dans l’espace P–T (d’après [6]). BS :faciès schiste bleu ; AEE : faciès éclogitique à amphibole-épitote ; ALE : faciès éclogitique à amphibole-lawsonite ; LE : faciès éclogitique àlawsonite ; E-HPG : métamorphisme à éclogite de moyenne température et à granulite de haute pression ; G : faciès granulite ; UHTM : partiemétamorphique d’ultrahaute température du faciès granulite. a : métamorphisme dévonien HP–LT (390–380 Ma) des unités de Münchberg etMariánské Láznĕ contrastant avec le métamorphisme MP (380 Ma) de la partie occidentale de l’unité de Teplá-Barrandien ; b : métamorphismecarbonifère HP–LT (340 Ma) du parautochtone Saxothuringien et de la nappe inférieure, contrastant avec le métamorphisme HP–HT de la granulite163

Author's personal copyK. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286 271the east up to kyanite zone (Figs. 1 and 2a; P =5–8 kbar,T = 530–620 8C) [156] in the west marked by a‘‘normal’’ gradient (increase of pressure and temperatureto the structural footwall). The 40 Ar/ 39 Ar muscoviteand hornblende dating yielded exclusively MiddleDevonian cooling ages [155].2.4. The Teplá-Barrandian–Moldanubian domainsboundary – the Central Bohemian PlutonicComplexThe igneous activity in the area of the CentralBohemian Plutonic Complex (Fig. 1) started withintrusions of calc-alkaline Devonian (protolith370 Ma) tonalites to granodiorites, latter transformedinto highly sheared orthogneisses [74]. The firstunmetamorphosed plutonic rocks were Late Devonian(354 Ma) calc-alkaline tonalites, granodiorites,trondhjemites, quartz diorites and gabbros of theSázava suite [62,64,158]. Their Sr–Nd isotopic ratiosand trace-element signature indicate that the source ofthe basic magmas was a slightly depleted mantle abovea subduction zone. In addition, a significant role formixing with acidic magmas is to be assumed [60,64].Further south/southeast occur voluminous Early Carboniferous(349–346 Ma [18,55,56]) high-K calcalkalineplutonic bodies of the Blatná suite (mainlygranodiorites with minor quartz monzonite and monzogabbrobodies). The intermediate rock types resultedfrom mixing of slightly enriched mantle-derived andcrustal magmas [61]. Finally, further east occur syndeformationalbodies or post-tectonic elliptical intrusionsof (ultra-)potassic rocks of mid-Carboniferous(343–337 Ma) ages [53,55,63,157]. Both the Sázavaand Blatná suites contain numerous xenoliths, screensand roof pendants of the Barrandian-like Palaeozoic andNeo-Proterozoic sequences, indicating a major role forstopping as an emplacement mechanism [159].All these features indicate that the Central BohemianPlutonic Complex corresponds to a relatively shallowsection (

Author's personal copy272K. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286conditions to the magmatic emplacement of granuliteprotoliths at moderate depth of c. 7–13 kbar, followed bycooling and prograde metamorphic overprint anddecompression occuring at temperatures lower than850 8C [128,129]. The granulite-facies overprint isprobably Visean in age as shown by a number of zirconages [1,84,122,131,153] but recent studies indicatepossible Devonian age of 370 Ma [2].Based on existing pressure-temperature (P–T)estimates two NW–SE trending belts of HP rocks(granulites, eclogites and peridotites) are distinguished,one located close to the Barrandian–Moldanubianboundary (the western belt of Finger et al. [29]) andthe other rimming the eastern margin of the BohemianMassif. These belts alternate with MP units, representedby the Varied and Monotonous groups, which also formNW–SE trending wide belts.The deformation history in the Moldanubian Zonereveals early vertical NNE–SSW trending fabrics,associated with crystallization of HP mineral assemblages[31,113,131,142,151]. These are reworked byflat deformation fabrics that are associated with MP tolow-pressure (LP) and high-temperature (HT) mineralassemblages [50,133,142,143,151]. The flat fabricsshow intense NE–SW trending mineral lineation that iscommonly associated with generalized ductile flowtowards northeast and this kind of deformation is typicalof the whole eastern margin of the Bohemian Massif.The early steep fabrics are dated at 350 to 340 Ma [122],while the ages of the flat ones cluster generally around335 Ma [123]. In the southwestern part of theMoldanubian domain, younger set of steep NW–SEmetamorphic fabrics reworks the flat foliation, havingbeen associated with LP metamorphic conditions ataround 325–315 Ma [29,138].The HP granulites are spatially, structurally andtemporally associated with (ultra-)potassic melasyenitesto melagranites, which can be divided into twogroups differing in modal mineralogy and textures: coarsely porphyritic K-feldspar melasyenites tomelagranites of the so-called durbachite group/serieswith a ‘‘wet’’ mineral assemblage Mg-biotite plusactinolitic hornblende (e.g., Čertovo břemeno andTřebíč intrusions); even-grained melasyenites–melagranites (Tábor andJihlava intrusions), containing a variously retrogressed,originally almost ‘‘dry’’, assemblage oftwo pyroxenes plus Mg-biotite [29,53,59,149].For the most basic ultra-potassic rocks, the highcontents of Cr and Ni with high melting point toderivation from an olivine-rich source (i.e., Earth’smantle). On the other hand, elevated concentrations of U,Th, LREE and LILE, pronounced depletion in HFSE aswell as high K/Na and Rb/Sr ratios seem to contradict themantle origin. This dual geochemical character andcrustal-like Sr–Nd isotopic compositions require meltingof anomalous lithospheric mantle sources, metasomatisedand contaminated by mature crustal material ([59]and references therein), and interaction of these maficmelts with crustally-derived leucogranitic magmas.The Moldanubian metamorphic units were penetratedby numerous and voluminous anatectic plutons looselygrouped into the Moldanubian (or South Bohemian)Plutonic Complex [26,45,46,54]. These are mostlyfelsic–intermediate, two-mica granitic to granodioriticintrusions with either S- or transitional I/S type character[26,88,148]. High-resolution conventional U–Pb zirconand monazite dating showed that the bulk of theMoldanubian Plutonic Complex (c. 80%) was emplacedat 331–323 Ma. The fine-grained granodiorites associatedwith minor diorites followed in a second, lessimportant event at 319–315 Ma [47]. The post-tectonicintrusions of the Moldanubian Plutonic Complexpostdated shortly the thermal peak of the regionalmetamorphism, but were significantly younger than theemplacement of both the Central Bohemian PlutonicComplex and (ultra-)potassic plutonic rocks scatteredthroughout the Moldanubian domain.Granitic rocks in the southwestern sector of theBohemian Massif (Bavarian Forest) have largelyindependent position. Here, large-scale crustal anatexiswas connected with a significant reheating (LP–HTregional metamorphism) and a tectonic remobilisation ofthe crust (Bavarian Phase sensu [29]). The intrusionsnortheast of the Bavarian Pfahl Zone are dated between328 and 321 Ma, whereas ages between 324 and 321 Maare obtained southwest of this zone. It apparentlyrepresents an important terrane boundary and the graniticintrusions sampled distinct basement units [124].2.6. The Moldanubian–Brunia continentaltransition zoneThis boundary was defined by Suess [136,137] as azone of severe deformation and metamorphism ofcontinental rocks (the so-called Moravo-Silesian Zone,Fig. 1). In his seminal work, Suess [136] defined the deepthrusting of the Moldanubian Zone over more externaland shallower Moravo-Silesian Zone, the latter emergingthrough Moldanubian nappe in a form of several tectonicwindows. The contact between these units is marked by aparticular unit, the Moravian ‘‘micaschist zone’’, which165

Author's personal copyK. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286 273is composed of kyanite-bearing micaschists that wereinterpreted by Suess [136] as a result of deep crustalretrogression (‘‘diaftorism’’) of the Moldanubiangneisses. Modern studies by Konopásek et al. [73] andŠtípská et al. [132] have shown that this zonecontains boudins of eclogites (Fig. 2; P = 16 kbar,T = 650 8C), HP granulites and peridotites embeddedin the metapelites. This zone, which contains elements ofboth Moravian and Moldanubian parentage, is thusregarded as a first order tectonic boundary.The underlying Moravo-Silesian Zone (Fig. 1) ischaracterized by two nappes composed of orthogneissesat the bottom and a metapelite sequence at the top.These nappes are thrust over the Neoproterozoicbasement which is often imbricated with its Pragianto Famennian cover [9,43]. A number of isotopic studies(e.g. [28]) show that the orthogneisses of the MoravianZone are derived from the underlying Brunia continent[20]. This 50 km wide and 300 km long zone of intensedeformation is marked by an inverted metamorphicsequence ranging from chlorite to kyanite-sillimanitezones and P–T conditions ranging from 5–10 kbar and550–650 8C (Fig. 2) [77,127]. The prograde metamorphismis interpreted as a result of continentalunderthrusting associated with intense top-to-the-NNEoriented shearing, development of sheath folds andgneissification of Brunia-derived granite protoliths[118,121]. The subsequent deformation is connectedwith recumbent folding and imbrication of Neoproterozoicgneisses with Devonian cover [119,120]. This laterphase is interpreted as late imbrications of metamorphosedBrunia crust in a kinematic continuum with earlyunderthrusting event. The cooling after the Barrovianmetamorphism is constrained at 340–325 Ma byhornblende Ar–Ar data [14,40] and monazite inclusionsin garnets yielding 340–330 Ma CHIME ages [78].2.7. The Brunia continentThe Brunia continental foreland (Fig. 1) originallycalled the Bruno-Vistulicum by Dudek [20] consists ofNeoproterozoic granitoids, migmatites and schists.They reveal the existence of a 680-Ma-old crust,intruded by 550-Ma-old granites [28,40]. This basementis unconformably overlain by Cambrian strata [7],Ordovician and shallow marine Lower Devonianquartzites and conglomerates followed by Givetiancarbonate platform sedimentation [3,68]. Since theEarly Carboniferous (350 Ma), foreland sedimentaryenvironment developed being accompanied by extensionalfaulting indicating the flexural subsidence of theBrunia margin [49,96]. From 345 until 300 Ma, a7.5 km thick Variscan flysch (Culm facies) wasdeposited onto the Brunia foreland. This sedimentationwas coeval with the onset of massive exhumation in theneighbouring Moldanubian domain and the continuousloading of the Brunia continent [49]. Low-grade sourcerocks associated with clastic muscovites dated at 345–330 Ma [117] gradually pass to a high-grade metamorphicsource material marked by pyrope-rich mineralfraction and granulite pebbles dated at 340–330 Ma[13,49,81,152]. Since 330 Ma, began also lateralshortening of the flysch basin marked by a deformationfront progressively migrating from eastward in conjunctionwith decreasing intensity of deformation[30,57]. Deformation of this flysch terminated atc. 300 Ma as indicated by Ar–Ar cooling ages of themetamorphosed Culm facies in the west [91] anddeformation of the Variscan molasse further east [11].3. Geodynamic evolution of the BohemianMassifThe spatial and temporal distribution of geologicalunits, magmatic fronts and metamorphic zones can beinterpreted in an evolutionary scheme very similar to thatcurrently reported from the Andean type orogeny. Thefollowing account provides a succession of tectonicevents that can be interpreted in terms of south-eastward(in the present-day coordinates) oceanic subductionunderneath an active continental margin, obduction ofthe Saxothuringian oceanic domain, formation of afore-arc region, growth of magmatic arc and developmentof a large-scale back-arc system on the continentallithosphere. The early oceanic subduction event couldhave been followed by a continental underthrusting of theArmorican plate leading to increased friction between theupper and lower plates, gradual flattening of thesubduction zone marked by eastward migration of arcand subsequent crustal thickening. The latter event couldhave been responsible for the development of a thickcontinental root due to thickening of the upper platerepresented by the Teplá-Barrandian and Moldanubiancrustal material. The final evolution is marked by thecontinental indentation of eastern Brunia continent into aweak orogenic root, exhumation of the Moldanubianorogenic lower crust, collapse of the Teplá-Barrandian lidand Moldanubian thrusting over the Brunia platform.3.1. Early Devonian oceanic subductionunderneath the active continental marginThe contact between the Saxothuringian-Armoricanplate and the overriding Teplá-Barrandian continent is166

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Author's personal copyK. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286 275marked by relics of Ordovician MORB eclogites andmetabasites locally associated with Ordovician sedimentsmetamorphosed under blueschist–eclogite faciesconditions indicating a Mid-Devonian oceanic subduction[34] and an existence of Lower PalaeozoicSaxothuringian oceanic domain (Fig. 3a). The exactage and size of the Saxothuringian ocean is not known,because the differences of Ordovician and LowerDevonian faunas between Saxothuringian and Barrandiancontinental domains are not confirmed, precludingan existence of large oceanic barrier of biogeographicalsignificance [10,52,106]. Therefore, the oceanic domainmust have been narrow and potentially short lived inorder not to be recorded by palaeobiogeographic faunalevidence. The only existing argument for larger oceanicseparation are from palaeomagnetic data of Tait et al.,[141] suggesting that the Saxothuringian and Tepla-Barrandian blocks rotated independently during Silurianand Devonian times.The hornblende and mica cooling ages of theVohenstraus–Erbendorf Zone, Mariánské Lázně complexand the Sovie Gory granulites cluster around380 Ma, which indicates early collision betweenthe Teplá-Barrandian and Saxothuringian domains(Figs. 3middle and 4B). This is also confirmed byFamennian flysh sediments of the Saxothuringian basinthat contain detrital zircons dated at 380 Ma [114]which is a direct evidence of mutual contact betweenthe Teplá-Barrandian and Saxothuringian domainsduring Famennian but most likely already during mid-Devonian. The Barrovian metamorphic zonation developedat the western flank of the Teplá-Barrandian domain(Figs. 3middle and 4B) is potentially related todeformation of the overriding Teplá-Barrandian continentalmargin and is consistent with a model ofextrusion of lower part of the Teplá-Barrandian crustduring Devonian shortening event [160,161]. The foldingof the eastern very low grade shales and volcanics of theNeoproterozoic sequences is interpreted as a supracrustalresponse to the same deformation event affecting highgrade rocks in the west [161]. Thus, the exhumation ofHP rocks in the Mariánské Lázně Complex and thedeformation of the Teplá-Barrandian basement arecoupled and related to convergent processes at theSaxothuringian plate boundary (Figs. 3middle and 4B).The strong argument for eastward subduction of theSaxothuringian ocean underneath the eastern Teplá-Barrandian continent is a Devonian–Lower Carboniferousmagmatic arc that is firmly founded on thecontinental crust. This arc, the Central BohemianPlutonic Complex, separated the Teplá-Barrandiandomain from the future Moldanubian domain and inthis model the former unit represented a fore-arc regionduring Devonian (Figs. 3middle and 4B). Thus, theposition of Devonian HP rocks thrust over theSaxothuringian basement, existence of the MariánskéLázně Complex – the oceanic fragment at sutureposition, and location of calc-alkaline magmatic rocksfurther east confirm a polarity of the oceanic subductionunderneath the eastern fore-arc and magmatic arcsystem during Late Devonian (Figs. 3middle and 4B).The distance between the arc and the trench arearepresented by the suture indicates that the dip ofsubduction zone was probably moderate (30–408,Fig. 4B). The temporal evolution of magma geochemistryfrom calc-alkaline to more potassic/shoshoniticaffinities (from 370 to 336 Ma) is compatible withflattening of the subduction zone and increased meltingof continental material during Early Carboniferous. TheBarrandian (Prague) Basin is interpreted as a part of thefore-arc basin marked by 380-Ma-old detrital zirconsfound in Mid-Devonian flyshoid sequences [135]suggesting early erosion of the eastern magmatic arc.The most problematic question in the Devoniansubduction model for the Bohemian massif is the role ofthe future Moldanubian domain as well as the positionof the Brunia continent. The amphibolites derived fromSiluro-Devonian tholeiitic basalts associated withcarbonates, widespread in Lower Austria and SouthBohemia, are interpreted as volcanic products of a relicFig. 3. Time slices maps showing the chronological and tectonic significance of the various units of the Bohemian Massif. Top: Ordovician timeslice showing the location of Ordovician MORB and deep marine sediments, and that of the Proterozoic units containing the Ordovician magmatism;middle: Devonian time slice showing the location of Devonian high pressure (HP) rocks, the passive margin sequences, the active marginmetamorphism, the fore-arc region, the magmatic arc, the location of Devonian infrastructure and sediments; bottom: Carboniferous time sliceshows the location of orogenic lower and middle crust belts in the Moldanubian domain, the metamorphism of the Moravian Zone, the MoravianMicaschist Zone, the HP units in Saxothuringian domain and the various types of magmatic rocks.Fig. 3. Cartes selon des tranches de temps montrant la signification chronologique et tectonique des diverses unités du Massif de Bohême. Enhaut : tranche de temps ordovicienne montrant la localisation du MORB ordovicien et des sédiments marins profonds et celle des unitésprotérozoïques incluant le magmatisme ordovicien ; au milieu : tranche de temps dévonienne montrant la localisation des roches de haute pression(HP) dévoniennes, les séquences de marge active, le domaine avant-arc, l’arc magmatique et la localisation de l’infrastructure et des sédimentsdévoniens ; en bas : tranche de temps carbonifère montrant la localisation des chaînes orogéniques crustales inférieure et moyenne du Domainemoldanubien, le métamorphisme de la zone moravienne, de la zone des micaschistes moraviens, des unités HP du domaine saxothuriningien et desdifférents types de roches magmatiques.168

Author's personal copy276K. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286Fig. 4. Conceptual model of the geodynamic evolution of the Bohemian Massif shown as schematic cross sections through the orogen. A. Silurian–Early Devonian subduction setting with the beginning of arc and back-arc formation on upper plate. B. Onset of the Mid-Devonian continentalunderthrusting of Saxothuringian lithosphere, deformation of active margin, formation of magmatic arc and persistence of back-arc region on theupper plate. C. Crustal thickening processes marked by the influx of Saxothuringian crust and progressive individualization of the Brunia activemargin. D. Brunia continent indentation associated with the channel flow process, imbrications of underthrust Brunia and the formation of theMoravian Zone.Fig. 4. Modèle conceptuel de l’évolution géodynamique du Massif de Bohême représenté par des coupes schématiques au travers de l’orogène.A. Subduction Silurien–Dévonien inférieur se mettant en place avec le début de la formation de l’arc et de l’arrière-arc sur la plaque supérieure.B. Au Dévonien moyen, mise en place du sous-charriage continental et de la lithosphère saxothuringienne, déformation de la marge active, formationde l’arc magmatique et persistance de la zone d’arrière-arc sur la plaque supérieure. C. Processus d’épaississement crustal marqué par la venue decroûte saxothuringienne et individualisation progressive de la marge active de Brunia. D. Indentation de Brunia, associée à un processus de flux dechenal, imbrications du sous-charriage de Brunia et formation de la zone moravienne.169

Author's personal copyK. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286 277of large-scale back-arc system [24,67]. In addition, thefelsic metavolcanics and amphibolite layers in theVaried Group are regarded as the continuity of back-arcbimodal volcanism till Givetian [37]. However, thesupposed depositional Devonian age of the Variedgroup is questioned thanks to low Sr isotopic ratioswhich indicate shallow marine environment duringLate Proterozoic rather than Palaeozoic [32]. Schulmannet al. [122] interpreted the common occurrence ofLate Silurian–Early Devonian zircons in the high gradeamphibolites of the Moldanubian domain as a result ofimportant magmatic reworking of the lower part of thecontinental crust associated with mafic magmaticadditions during lithospheric thinning of this domain(Fig. 3middle). In their model, the Monotonous Grouprepresents relic of Proterozoic middle crust that is lessaffected by thermal and magmatic reworking while theupper crust recorded sedimentary and volcanic evolutionrelated to the Silurian–Devonian extensionalevent. A back-arc environment is further supportedby bimodal volcanic activity in narrow Devonian basinsdeveloped on the north-eastern margin of the Bruniacontinent [107] suggesting only minor thinning ofcontinental crust at the easternmost termination of theback-arc system (Figs. 3middle and 4B). In this conceptthe rest of the Brunia platform represents a stablecontinental domain not affected by the back-arcspreading.The Devonian Saxothuringian oceanic subuductionand onset of continental underthrusting of the Saxothuringian–Armoricanlower plate underneath theupper plate is thus recorded in all units forming thepresent-day Bohemian Massif (Fig. 5). The followingcontemporaneous processes were identified: eclogitization of the Saxothuringian crust and itsexhumation along the Teplá suture; sedimentation of Mid-Devonian distal turbidites ofthe Srbsko Formation followed by inversion of thefore-arc basin in the Teplá-Barrandian domain, originof the magmatic arc in the area of the CentralBohemian Plutonic Complex and exhumation of thewestern metamorphosed margin of the Teplá unitsuggesting convergent processes in the upper plate; formation of the back-arc system on the continentallithosphere as shown by isotopic data in theMoldanubian domain and marginally in the Devonianbasins affecting western margin of the Bruniacontinent (here, the formation of small oceanic basinis not excluded). All existing data point out to thesubduction process that operated at least 45 millionyears from 400 to 355 Ma (Fig. 5).3.2. The Early Carboniferous crustal thickening ofthe upper plateThis event is recognized in all units except the Teplá-Barrandian supracrustal unit. The western margin of theBohemian Massif is characterized by the arrival of theSaxothuringian continental crust and its subductionunderneath the eastern Teplá-Barrandian–Moldanubiandomain (Fig. 4B). The main thrust boundary migratedfurther west, so that the continental crust was thrustunderneath the fossil Devonian suture and former forearcregion [71]. At the same time the deformationregime changed in the far field back-arc region (futureMoldanubian domain), which recorded the progressivethickening of the whole previously thinned andthermally softened domain as indicated by severalrecent petrological studies (e.g. [77,128]). Recentstructural studies have shown that the earliest preservedfabrics have been sub-horizontal [31,112,123,131],which may indicate that the lower crustal materialwas originally flowing horizontally from the area of thecontinental subduction channel towards the region ofeastern backstop in a manner proposed by [104] for theclosure of the Rhenohercynian Basin and formation ofthe Mid-German Crystalline Rise.Indeed, the influx of lower crustal material transportedby south-east dipping Saxothuringian continentalsubduction zone underneath the fore arc (the Teplá-Barrandian domain) and further below the former backarcdomain is regarded to be at the origin of the future‘‘Gföhl Unit’’. This hypothesis is in line with the wholerockgeochemical and Sr–Nd isotopic composition aswell as the zircon inheritance patterns in the MoldanubianHP–HT granulites [59,65]. Importantly, thecrustal material involved in the subduction and extrudedover the sub-arc and sub-back-arc mantle lithospheremay have been turned into voluminous HP granulitesknown from many regions of the Bohemian Massif[59,100]. Such a processing of the lower platecontinental lithosphere in a subduction zone and itsextrusion above mantle wedge was successfullymodelled by Gerya and Stockert [48]. Alternatively,the back-arc domain with high thermal budget inheritedfrom Devonian stretching may have been thickened andthe partially molten lower crust may have beentransported downwards and transformed into the HPgranulites as suggested by Štípská and Powell [129] orSchulmann et al.[122]. However, this model fails toexplain the occurrence of the Gföhl gneiss and felsicgranulites in the lower crustal position.The onset of thickening of the root is not recordedin the Teplá domain (Fig. 4C), which behaved as a170

Author's personal copy278K. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286Fig. 5. Summary of geochronology data showing the time scales of the three major events forming the Bohemain Massif: oceanic subduction andcontinental underthrusting of Saxothuringian domain, building of thick orogenic root system and collapse of orogen due to indentation of Brunia.Fig. 5. Résumé des données chronologiques montrant les échelles de temps des trois évènements majeurs de la formation du Massif de Bohême :subduction océanique et sous-charriage du domaine saxothuringien, édification de l’épais système d’enracinement orogénique et « collapse » del’orogène, en raison de l’indentation de Brunia.supra-structural unit at this time, but it is shown bydeformation of the Lower Palaeozoic rocks of thePrague basin and adjacent Late Proterozoic rocks. Here,the steep fabric is well dated by syntectonic calcalkalineplutons at about 355–345 Ma [64,122,158]. Incontrast, the eastern sector of the orogen records onsetof loading of the Brunia platform (Fig. 4C) duringTournaisian manifested by destruction of the Givetiancarbonate platform and sedimentation of coarse basalclastics [49].The timing of crustal thickening is relatively poorlyconstrained compared to the subduction and laterexhumation processes (Fig. 5). However, the Tournaisianand Early Visean massive clastic sedimentation onboth Saxothuringian and Brunia plates suggests load ofboth continental lithospheric plates and high topographyin between them. This corroborates the peakmetamorphic ages in the granulites and the Moldanubianeclogites as well as the Visean age of compressionof the magmatic arc. All geochronological and othergeological information point to a crustal thickeningperiod that was very short and did not last more than20 million years, from 355 to 335 Ma, with a peakaround 340 Ma (Fig. 5).3.3. Late Visean exhumation of orogenic lowercrust of the upper plateThe exhumation of the Variscan lower crust duringEarly Carboniferous is exemplified by the three NE–SW171

Author's personal copyK. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286 279trending belts of granulites, eclogites and peridotites(Fig. 3c) intimately associated with the (ultra-)potassicmagmatites [29,59]. The first granulite belt is representedby narrow strip of felsic granulites that occur at theSaxothuringian–Teplá-Barrandian boundary and it isinterpreted as an extrusion of the orogenic lower crustalong a large scale crustal shear zone at 340 Ma[71,163]. The coeval 340 Ma age of the Saxothuringianeclogites and of the felsic granulites [80,150] ledKonopásek and Schulmann [71] to propose a model ofsimultaneous exhumation of nappes derived from theSaxothuringian crust and viscous extrusion of orogeniclower crust from underneath the Teplá-Barrandian supracrustalunit (Fig. 4D) [36]. The second belt, recognizedeast of the Central Bohemian Plutonic Complex, i.e.,‘‘the magmatic arc’’, was exhumed along huge westdipping detachment zone [111,157,158], which was alsoresponsible for collapse of the upper part of the magmaticarc system and the downthrow of the whole Barrandiansection [115,157]. Such a huge vertical material transfer(Fig. 4D) could have been responsible for verticalexchange of the lower crustal and upper crustal materialin a range of 50 km with final throw of 15 km [110,162].The cooling ages from the lower crustal domain show thatthe granulites have crossed the 300 8C isotherm duringCarboniferous (330–310 Ma) [75,138] suggesting thetime at which the lower crustal bulge reached shallowposition in the upper plate.The third lower crustal belt rims the eastern marginof the Bohemian Massif, i.e. the boundary with theBrunia continent (Fig. 4D). Here the granulite fabricis also vertical and interpreted in terms of massivevertical exchanges with orogenic middle crust [123].The zone of the lower crustal bulge is interpreted as anenormous zone of vertical extrusion surrounded by amiddle crust coevally transported downwards in formof a crustal scale synform. The vertical materialtransfer along the Moldanubian lower crustal beltswas a matter of research for several Czech authorsduring the last five years [31,112,123,131,142]. Themodel of vertical extrusion is based on the concept ofbuckling of the lower and mid-crustal interfacefollowed by growth of crustal scale antiforms. Thisprocess is thought to be triggered by rheological andthermal instabilities in the arc region, while to the eastit is forced by rigid backstop, preserved only locally[130].However, the most important feature of the easternVariscan front is the development of horizontal fabricsin the Moldanubian root zone parallel to the Bruniacontinental margin. The intense deformation of theBrunia continent leading to the formation of theMoravo-Silesian imbricated nappe system was associatedwith the development of tectonically invertedBarrovian metamorphism (Fig. 4D) and formation ofcrustal mélange in its upper part (the MoravianMicaschist Zone). These phenomena, as well as mixingof HP rocks and migmatites in the overlyingMoldanubian nappe have been recently interpreted interms of indentation of the Brunia continent into the hotand thick continental root [123]. This lower crustalindentation and flow of hot lower crustal rocks insupracrustal levels are consistent with a model ofcontinental channel flow driven by the arrival of acrustal plunger, a model which is advocated for twodecades for the deformation of the Eastern Cordillera inthe Andes (e.g. [87]). Finally, the continuous load of theBrunia platform related to deep indentation process ledto the development and eastward propagation of theforeland basin. In our model, as the hot Moldanubianrocks advance over the Brunia platform, an imbricatedfootwall nappe system is generated and thrust over theprogressively buried foreland basin rocks.The time scale of exhumation of the orogenic lowercrust is well constrained due to a set of well dateddiachronous processes that start with the exhumation ofthe West-Bohemian granulites, growth of western andeastern orogenic lower crustal antiforms and a shallowchannel flow of partially molten lower crust associatedwith the inversion of the eastern foreland basin. All thatis linked with a major thermal event in the mantle asshown by the ages of a syn-extrusion high potassicmagmatism, the exhumation of large number of mantlefragments and the overall melting of the Moldanubiancrust. The time scale of all these processes wassurprisingly short, i.e., about of 20 Ma, and ranges from335 to 315 Ma (Fig. 5).4. Palaeographic constraints for Andean typeorogeny in the Bohemian MassifFinger and Steyrer [24] attempted to explain thetectonic processes at the Moldanubian–Moravianboundary with a plate tectonic model involving thesubduction of a Silurian–Devonian oceanic domainwestward beneath the Moldanubian zone. This conceptsupposes large rotation of the Brunia platform forming apart of an Old Red continent (Avalonia) around the coreof the Bohemian Massif during Visean times [140]. Themodel of westward subduction is also based on thepresence of Silurian MORB-type amphibolites [25]within the Gföhl Unit, which are interpreted as relics ofoceanic crust of the Rheic Ocean. In addition, theoccurrences of eclogites located along the eastern172

Author's personal copy280K. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286Variscan front [73,132] can serve as an additionalargument for this subduction model.However, the palaeomagnetic investigations carriedout on granitoids of the Central Bohemian Pluton and itsextension to the east (the Nasavrky Plutonic Complex[58]) demonstrate that the Teplá-Barrandian, theMoldanubian as well as Brunia domains had a commongeodynamic behaviour since the Late Visean (335–330 Ma according to Edel et al. [22]). Therefore, theinterpretation of palaeomagnetic directions from Devoniansediments of the Brunovistulian platform[86,98,139] in terms of an oroclinal bending of theRhenohercynian domain along the Moldanubian corecannot be valid for several reasons: the consistency of these ‘‘Devonian’’ palaeomagneticdirections with Middle-Late Carboniferous directions(Cp directions at 330–325 Ma and B directions at320–315 Ma, Fig. 6) in the Moldanubian zonesuggests that the Devonian sediments overlying theBrunia continent were re-magnetized in Late Variscantime. Consequently, no relative rotation between theBrunia continent and the Moldanubian core can beconsidered prior to 330 Ma; magnetic overprinting of the Devonian sequences onthe Brunia continent is supported by importantdeformation and burial of Culm and Devoniansediments during Late Carboniferous deformation ofthe Brunia margin [30,120,126]. Hence, it is unlikelythat primary magnetizations could have survived sucha severe structural and thermal reworking of theDevonian sediments. We suggest that magnetic overprintingwas the result of the Carboniferous tectonothermalprocesses associated with the underthrusting ofthe Brunia platform together with Culm accretionarywedge [11] underneath the hot Moldanubian root zoneand with later uplift and erosion of the Moldanubian–Moravian nappe sequence.Fig. 6. a: paleomagnetic north directions obtained in Devonian (meta)sediments from the Teplá-Barrandian and Brunia zones (K: [86];N:[98];T:[139]; E:[21]) and from Early Carboniferous granitoids from Central Bohemian Pluton and Nasavrky Pluton [22]. Cn and Co directions representmagnetizations acquired at the end of the NW–SE shortening and the uplift of the Moldanubian root; Cp and B directions correspond to overprintsacquired during NNE–SSW compression and clockwise rotation of the Variscides. Note the similarity of ‘‘Devonian’’ directions with Carboniferousdirections; b: published mean virtual geomagnetic poles (VGPs) and associated apparent polar wander curve (grey dashed line) from Early Paleozoicrocks of the Bohemian Massif and mean poles from western Europe Variscides (stars with ages of magnetization) and associated apparent polarwandering curve (dark grey line) [22].Fig. 6. a : directions du nord paléomagmatique obtenues dans les (méta)sédiments dévoniens des zones Teplá-barrandienne et Brunia (K :[86] ;N :[98] ; T :[139] ; E :[21]) et dans les granitoïdes du Carbonifère inférieur et du Pluton Nasavrky [22]. Les directions Cn et Co représentent lesmagnétisations acquises à la fin du raccourcissement NW–SE et du soulèvement de la racine moldanubienne ; les directions Cp et B correspondent àdes réaimantations acquises durant la compression NNE–SSW et la rotation dans le sens des aiguilles d’une montre des Variscides. À noter, lasimilarité des directions « dévoniennes » et carbonifères ; b : pôles géomagnétiques virtuels (VGP) publiés, associés à la courbe de dérive apparentedes pôles (ligne en tiretés) pour les roches du Paléozoïque inférieur du Massif de Bohême et pôles moyens pour les Variscides d’Europe occidentale(étoiles avec âges de la magnétisation), associés à la courbe de dérive des pôles (ligne continue gris foncé) [22].173

Author's personal copyK. Schulmann et al. / C. R. Geoscience 341 (2009) 266–286 281The Siluro-Devonian basin in the area of theMoldanubian zone existed as a back-arc basin abovethe Saxothuringian subduction zone [122] but thequestion of development of oceanic crust in this arearemains a matter of discussions. Therefore, in contrastto the generally accepted rotation of a Brunia continentindependent of a stationary Moldanubian domainassociated with a closure of a large (Rheic) oceanicdomain we propose a model of common geodynamichistory of the Bohemian Massif during Devonian andEarly Carboniferous (Fig. 5). In our model thepalaeomagnetic, crustal scale geophysical and structuraldata are in favour of an early counterclockwiserotation of composite blocks (Saxothuringian, Barrandian,Moldanubian and Brunia altogether) accommodatedby a large scale NW–SE trending dextral wrenchzones (such as the Elbe, Pfahl, Franconian and Pays deBray faults) during Early Visean [22]. These movementscould have resulted from northwest drift ofGondwana continental masses and continued after330 Ma by clockwise rotation of the whole Variscan beltaround the Euler pole that was continuously translatedwestward parallel to the Teysseire–Tornquist zone(southern margin of the Baltica).AcknowledgementsThe French National Science Foundation project‘‘ANR LFO in orogens’’, internal research funds ofCNRS UMR 7615 and grant MSM0021620855 of theMinistry of Education of the Czech Republic areacknowledged for financial support and salary of OndrejLexa. Jiří Konopásek appreciates the financial supportof the Grant Agency of the Charles University (projectNo. B-GEO-270/2006), as well as the support by theMinistry of Education, Youth and Sports of the CzechRepublic through the Scientific Centre ‘‘AdvancedRemedial Technologies and Processes’’ (identificationcode 1M0554).References[1] M. Aftalion, D. Bowes, S. 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J. metamorphic Geol., 2011, 29, 79–102 doi:10.1111/j.1525-1314.2010.00906.xHeat sources and trigger mechanisms of exhumation of HPgranulites in Variscan orogenic rootO. LEXA, 1 K. SCHULMANN, 2 V. JANOUŠEK, 1,3 P. ŠTÍPSKÁ, 2 A. GUY 2,4 AND M. RACEK 1,31 Institute of Petrology and Structural Geology, Faculty of Science, Charles University in Prague, Albertov 6, 128 43 Prague 2,Czech Republic (lexa@natur.cuni.cz)2 Institute de Physique de Globe, UMR 7516, École et Observatoire de Science de la Terre, Université de Strasbourg, 1 RueBlessig, Strasbourg 67084, France3 Czech Geological Survey, Klárov 3, 118 21 Prague 1, Czech Republic4 Department of Geophysics, Faculty of Mathematics and Physics, Charles University in Prague, V Holešovičkách 3, 180 00Prague 8, Czech RepublicABSTRACTThe structure of the Moldanubian domain is marked by felsic granulites of Ordovician protolith ageforming the cores of domes that are separated from mid-crustal Neoproterozoic and Palaeozoicmetasedimentary rocks that occur in synclines by a late Ordovician to Silurian metabasic unit. Reflectionand refraction seismic sections combined with gravity inversion modelling suggest the presence of a lowdensity layer at the bottom of the crust (interpreted as felsic granulite) overlain by a denser layer(interpreted as amphibolite) with layers of intermediate density at the top (interpreted as metasedimentaryrocks). It is proposed that the granulite domes surrounded by middle crustal rocks reflect transposedhorizontal layering originally similar to that preserved in the deep crust and imaged by the geophysicalsurveys. This geological and geophysical structure is considered to be a result of Vise´an gravityredistribution initiated by radioactive heating of felsic crust tectonically emplaced at the bottom of aPalaeozoic orogenic root. The radioactive layer with heat production of 4 lW m )3 correspondsgeochemically and isotopically to Ordovician felsic metaigneous rocks of the Saxothuringian domain thathave been emplaced at Moho depth under thickened crust during late Devonian–early Carboniferouscontinental subduction. Part of the continental crust continued to be subducted and produced fluids ⁄low-volume melts which directly contaminated and enriched the local lithospheric mantle by lithophileelements, most notably Cs, Rb, Li, Pb, U, Th and K. Thermal incubation of 10–15 Myr was sufficient toheat and convert the underplated felsic layer into granulites via dehydration melting and meltsegregation. The process of melt loss was responsible for the removal of radioactive elements and forswitching off the heat at the beginning of the exhumation process. At the same time, the metasomatizedunderlying mantle was heated producing characteristic ultrapotassic magmas. Gravitational instabilitywas then induced by the density contrast between the light granulites and the overlaying denser maficlower crustal layer and a viscosity drop related to thermal weakening and partial melting of the latter.Key words: crustal structure; exhumation of lower crust; heat sources; radioactive heating.Mineral abbreviations: bi, biotite; mu, muscovite; g, garnet; ky, kyanite; sill, sillimanite; cd, cordierite; pl,plagioclase; ksp, K-feldspar; liq, granitic liquid; o, omphacitic clinopyroxene; ilm, ilmenite; ru, rutile.INTRODUCTIONThe Variscan Bohemian Massif, Czech Republic, ischaracterized by the occurrence of large high-pressure(HP) granulite bodies within the central high-gradepart of the orogen. These granulite bodies are mainlycomposed of alkali feldspar-garnet-kyanite granulites,eclogites and fragments of mantle peridotites surroundedby mid-crustal rocks. The exhumation andemplacement of the whole assemblage is commonlyinterpreted in terms of two contrasting models. In thefirst model, granulite massifs are inferred to be allochthonous,representing klippen of far travelled nappesrooted in the central part of the Bohemian Massif (e.g.Franke, 2000), whereas in the second model, granulitescorrespond to eroded windows of the orogenic infrastructurethat have been vertically extruded to midcrustallevels from lower crustal depths (e.g. Sˇtı´pska´et al., 2004; Schulmann et al., 2005). The latter modelis analogous to that proposed for the Saxonian granuliteMassif by Behr (1978) and Weber (1984).The problem of exhumation of granulite lower crusthas been discussed by a number of authors in relationto Archean and Mesoproterozoic orogenic belts (e.g.Perchuk, 1989). In these terranes, the emplacement ofhot granulite lower crust has been interpreted in termsof gravity overturn driven by an inverted densityprofile (e.g. Roering et al., 1992), whereby heavy maficÓ 2010 Blackwell Publishing Ltd 79181

80 O. LEXA ET AL.rocks (greenstone belts) rest upon light intermediate tofelsic TTG rocks [Tonalite-Trondhjemite-Granodioritegneisses]. Perchuk (1989) proposed that such crustalgravity redistributions are triggered by heat fluxresulting from large-scale plume tectonics.Gerya et al. (2001) proposed a conceptually similarmodel of crustal-scale gravity redistribution to explainthe emplacement of felsic orogenic crust into supracrustallevels of Palaeozoic orogens. In contrast toArchean and Mesoproterozoic orogenic belts, the heatsource is inferred to be bulk radioactive heat productionwithin a granodioritic lower crustal layer locatedat the bottom of the crustal pile (Gerya et al., 2002,2004). These authors suggested that the gravitationalinstability of doubly stacked lithologically heterogeneouscrust is related to an initial density contrast ofdissimilar intercalated layers enhanced by hightemperaturephase transformations.In this article, it is argued that the granulite bodiesabundant in the central parts of the Bohemian Massifcould have originated through a similar kind ofgravity redistribution enhanced by lateral forcing.Using gravity inversion modelling, we demonstratethat relics of an original stratification with densityinversion are still preserved in the Variscan crust ofthe Bohemian Massif, with a high density mafic middlecrustal layer resting upon 10 km of felsic crust(Guy et al., 2010). It is also shown that the internalstructure within each of the granulite massifs forms apattern consistent with deceleration of verticallyemplaced diapiric-like bodies and horizontal spreadingof low-viscosity, partially molten rocks at supracrustallevels. Geochemical and geochronological data areused to argue that the felsic lower crust (FLC) is anallochthonous body emplaced underneath the preexistingmafic lower crust during late Devonian–earlyCarboniferous continental subduction. The time-scalesof thermal processes driving the gravity redistributionare estimated for two end-member scenarios: (i) amodel of radioactive heat production solely in theFLC and ⁄ or mantle; and (ii) a large-scale thermalanomaly in the mantle caused by deep mantle processessuch as tectospheric root delamination (Deweyet al., 1993) or slab break off (Chemenda et al., 2000)as proposed for the Bohemian Massif by Janousˇek &Holub (2007). Finally, we offer a dynamic numericalmodel which is consistent with the available geologicaland geophysical data and allows assessment of exhumationmetamorphic paths of the granulite rocks inthe core of the Bohemian Massif.GEOLOGICAL SETTINGTraditionally the Bohemian Massif has been subdividedinto four main tectonic domains, which are, fromthe west to the east (Fig. 1): the Saxothuringian, theTepla´-Barrandian, the Moldanubian and the Bruniadomains. The Saxothuringian domain is regarded as aNeoproterozoic continental block accreted in theDevonian to, and partly subducted under, the moreeasterly Tepla´-Barrandian (suprastructure or orogeniclid) and Moldanubian (infrastructure or deep orogenicroot) continental domains of the Variscan orogen(Schulmann et al., 2009). The oceanic Tepla´ suture, leftover after the closure of Saxothuringian ocean, is representedby Devonian to Carboniferous HP and ultrahigh-pressure rocks, gabbros and mantle fragmentslocated between the Saxothuringian and the Tepla´-Barrandian domains (Mlcˇoch & Konopa´sek, 2010 andreferences therein). The whole system is bounded fromthe east by a Neoproterozoic Brunia promontory,which indented the Moldanubian orogenic root duringthe early Carboniferous (Schulmann et al., 2008).Moldanubian domainThe Moldanubian domain is composed mostly of theorogenic middle crustal unit subdivided into two lithostratigraphicgroups: the Monotonous Group of probableNeoproterozoic age (Friedl et al., 2004), and theVaried Group, at least partly of early Palaeozoic affinity(e.g. Janousˇek et al., 2008). Their mutual contact isdefined by uniformly deformed bodies of granitic gneissof Neoproterozoic to Mesoproterozoic age (Wendtet al., 1993; Friedl et al., 2004). The Varied Group iscomposed of metasedimentary rocks intercalated withamphibolites, quartzites, marbles and calc silicates. TheMonotonous Group consists of paragneisses intercalatedwith orthogneiss bodies. The lower part of thelatter group is characterized by a thick sequence ofamphibolites and metagabbros (Racek et al., 2006)locally containing eclogites (OÕBrien & Vra´na, 1995).The orogenic lower crust is represented by the Gfo¨ hlUnit (Fuchs, 1976), which comprises felsic and intermediateHP granulites accompanied by A type eclogites,garnet pyroxenites and peridotites (Medaris et al.,1995), amphibolites accompanied by Mid-Ocean RidgeBasalt-type eclogites (Sˇtı´pska´ & Powell, 2005a) andanatectic Gfo¨ hl orthogneisses.Today the structure of the Moldanubian domain ischaracterized by alternation of three, NE–SW trendingorogenic lower crustal and middle crustal belts (Fig. 1;Behr, 1978; Finger et al., 2007). The first occurs mostlyin south Bohemia, the second belt follows the easternboundary between the Moldanubian domain andBrunia continent and the third is a less extensive belt ofgranulite rocks tracing the contact between the Tepla´-Barrandian and the Saxothuringian domains The lastlarge occurrence of granulites is represented by theNE–SW trending domal structure of the Saxoniangranulite Massif (e.g. Franke, 2000).Recent studies of the principal granulite bodiesreveal a systematic lithotectonic pattern marked bygranulite lenses rimmed by amphibolite–gabbro belts(AGB) of the local ÔBegleitÕ (= accompanying) seriesand Gfo¨ hl orthogneiss (Fig. 2). The boundary betweenthe AGB and the Monotonous Group is not clearlydefined, and hence these basic rocks are sometimesÓ 2010 Blackwell Publishing Ltd182

HEAT SOURCES AND EXHUMATION MECHANISMS 81Fig. 1. Simplified geological map of the Bohemian Massif (modified after Franke, 2000). CBPC: Central Bohemian Plutonic Complex;CMP: Central Moldanubian Pluton. The lower left insert shows the position of the Bohemian Massif in the European Variscides(modified after Edel et al., 2003). RH: Rhenohercynian zone; ST: Saxothuringian Zone; M: Moldanubian Zone; B: Brunia Continent;L: Lugian domain.attributed to the lowermost part of the MonotonousGroup (Racek et al., 2006) and sometimes to theGfo¨ hl Unit (Fuchs, 1976). In addition, the generalpattern is complicated by the presence of large accumulationsof amphibolites within the Varied andMonotonous groups without clear tectonostratigraphicaffinity. The amphibolites and gabbros inLower Austria reach several kilometres in thickness,but locally narrow to several hundred metres (e.g.Tajcˇmanova´ et al., 2010) and in some places they arecut out completely (Franeˇk et al., 2006). Because ofthe regional significance of these rocks, they are interpretedby some authors as a relict of the Silurian (c.430 Ma) oceanic crust (Finger & von Quadt, 1995) oras the result of late Ordovician igneous activity relatedto thinning of continental crust (Schulmann et al.,2009). In contrast, the Gfo¨ hl orthogneiss yields dominantlyNeoproterozoic protolith U–Pb zircon ages(e.g. 550 ± 1 Ma; Schulmann et al., 2005) accompaniedby early Ordovician ages (e.g. 488 ± 6 Ma;Friedl et al., 2004). The Moldanubian felsic–intermediategranulites reveal a more complex spectrum ofprotolith U–Pb zircon ages clustering at c. 360, c. 400and 470–450 Ma (Kro¨ ner et al., 2000; Friedl et al.,2004; Janousˇek et al., 2004b).CURRENT DEEP STRUCTURE OF THE FORMERVARISCAN ROOTA new model of the structure and composition ofVariscan crust in the Bohemian Massif was recentlyproposed by Guy et al. (2010) based on 3D gravitymodelling, geological data and seismic refraction andreflection sections (Tomek et al., 1997; Hrubcova´et al., 2005; Ru˚žek et al., 2007). All results suggest thatthe deep structure of the Bohemian Massif crust, whichwas consolidated during the Variscan orogeny, reflectstectonic processes related to Palaeozoic subductionand collision at the subcrustal lithosphere level as well(Babusˇka et al., 2010).According to this model, the crust is characterizedby a succession of positive and negative anomalies of60–80 km wavelength for a nearly constant Mohodepths. The central part of the Bohemian Massif displaysa large negative Bouguer anomaly correspondingto the Palaeozoic crustal root represented by theMoldanubian domain (Fig. 3). The adjacent NeoproterozoicBrunia microcontinent displays an importantgravity high caused by mafic and intermediate medium-grademetamorphic and magmatic rocks. However,the strong gradient marking the deep crustalÓ 2010 Blackwell Publishing Ltd183

82 O. LEXA ET AL.(a)(a´)(b)(b´)(a)(b´)(b)(a´)Fig. 2. Geological maps of Blansky´ les and St. Leonhard granulite areas with schematic structural profiles. Two P–T space insetsshow the metamorphic evolution of lower crustal and middle crustal rocks (Petrakakis, 1997; Pitra et al., 1999; Scheuvens, 2002;Sˇtı´pska´ & Powell, 2005b; Racek et al., 2006, 2008). Protolith ages (Kro¨ ner et al., 2000; Friedl et al., 2004; Verner et al., 2008) ofmajor rock types are shown on the map.boundary between the root domain and the Bruniamicrocontinent is shifted 50–70 km westwards relativeto their contact on the surface suggesting that the highdensity basement rocks are covered by a thin sheet oflight granulites and migmatites in this area (Schulmannet al., 2008). North-west of the Moldanubian domainthere is an important gravity high corresponding to theNeoproterozoic basement of the Tepla´-BarrandianUnit. This is limited to the north by southeast dippingreflectors of the Tepla´ suture, which is characterized byhigh density eclogites and ultramafic rocks. The footwallof the suture corresponds to low density felsiccrust of the Saxothuringian basement.The seismic reflection and refraction sections andgravity modelling suggest a complex lithologicalstructure of the Moldanubian domain marked by a lowdensity, 5–10 km thick lower crustal layer locatedabove the Moho, a 5–10 km thick dense mafic layer, a10-km thick mid-crustal layer of intermediate densityand a locally developed 2–5 km thick low density layerat the top (Fig. 3). The low density lower crust correlateswell with low-P velocities in the range 6.0–6.4 km s )1 in the CEL09 section (Ru˚žek et al., 2007).Guy et al. (2010) proposed that the low density layerlocated above the Moho corresponds to felsic rocks,which are interpreted as deep crustal equivalents ofsurface outcrops of the Gfo¨ hl Unit. These authorsinterpreted the high density thick layer located abovelight granulites as an equivalent of the AGB (Fig. 3).The intermediate density layer forming recent uppercrust of the Moldanubian domain is interpreted asMonotonous and Varied group rocks, whereas the lowdensity rocks on the surface are directly correlated withexposures of the Gfo¨ hl Unit. It is suggested that thislayered structure of the Variscan crust reflects that ofthe original thick root, which was thinned by lateVariscan and Permian extensional processes (Burget al., 1994).Exhumation model connecting deep crustal geophysics andsurface geologyGuy et al. (2010) and Franěk et al. (2011a) proposed amodel connecting the deep structure of the Variscancrust with surface distribution of lower and mid-crustalrocks. The layered structure of the orogenic rootreported by geophysics thus represents a relict ofCarboniferous distribution of horizontally layeredcrust prior to exhumation. In addition, it is supposedthat the observed vertically layered distribution oforogenic lower crust surrounded by middle crustalunits reflects steepening of the deep crustal horizontallayering. The model connecting deep crustal layeringwith surface geology is based on several recent studiessuggesting that the granulites were exhumed alongsteep channels from lower crustal depth by a ductileextrusion mechanism (Sˇtı´pska´ et al., 2004; Schulmannet al., 2005; Franeˇk et al., 2006, 2011a; Tajcˇmanova´et al., 2006).In this view, the granulites represented, beforeexhumation, the structurally deepest orogenic lowercrust located at a depth of 60–70 km (18–20 kbar,Ó 2010 Blackwell Publishing Ltd184

HEAT SOURCES AND EXHUMATION MECHANISMS 83ABABFig. 3. Bouguer anomaly map of the Bohemian Massif (combined data provided by the Czech Geological Survey and Guy et al.(2010)) and gravimetric model for section A–B (profile no. 4 in Guy et al., 2010). The main lithological unit boundaries are representedby superimposed thick black lines. The section A–B emphasizes the present density structure of the Moldanubian Domain with relictsof felsic rocks in the lower crust.800–900 °C; Sˇtı´pska´ & Powell, 2005b; Tajcˇmanova´et al., 2006). Structurally above were mafic rocks of theAGB, which formed during early Palaeozoic magmaticunderplating of the Proterozoic crust of the MonotonousGroup (Schulmann et al., 2005, figs 16 & 17), andthe structurally highest Varied Group corresponding tosupracrustal early Palaeozoic sequences (Schulmannet al., 2009). According to this hypothesis, the extrusionÓ 2010 Blackwell Publishing Ltd185

84 O. LEXA ET AL.mechanism lead to inversion of this crustal profile sothat the granulites presently form cores of crustalanticlines, surrounded by AGB, variably thinned duringextrusion, and by the Gfo¨ hl gneiss followed bymetamorphic sequences of the Monotonous and Variedgroups. In many places, the whole structure is eveninverted because of the lateral spreading of verticallyextruded lower crustal material (Fig. 2). This model iscorroborated by detailed structural and petrologicaldata, which show that the granulites and surroundingrocks exhibit common vertical fabrics which are associatedwith HP–high temperature (HT) mineralassemblages. The flat fabric, which locally may have abowl shape, is superimposed on vertical extrusionfabrics and developed during subsurface spreading ofthe extruded partially molten lower crust (7–4 kbar,700–650 °C; Racek et al., 2006; Tajcˇmanova´ et al.,2006; Hasalova´ et al., 2008b).CHRONOLOGY OF THE VARISCANCALC-ALKALINE TO POTASSIC MAGMATISM –CONSTRAINTS ON TIME-SCALES OF THEOCEANIC AND CONTINENTAL SUBDUCTIONThe contact of the Tepla´-Barrandian and Moldanubianunits is marked by voluminous granitic plutons ofVariscan age with (normal or potassic) calc-alkalinechemistry and large ion lithophile element (LILE) ⁄ -high field strength element (HFSE) enrichmentresembling magmatic associations of active continentalmargins. The oldest vestige of this igneous activityinitiated by a late Devonian–early CarboniferousAndean-type subduction is preserved in the Cˇista´ andSˇteˇnovice plutons (Venera et al., 2000; Zˇa´k et al.,2010) as well as orthogneisses in the roof of the largeCentral Bohemian Plutonic Complex, CBPC (protolithc. 370–360 Ma; Kosˇler et al., 1993). A newlyidentified member of the subduction-related associationis the mafic Lisˇov low-pressure granulite unit insouthern Bohemia, the protolith of which was emplacedat c. 360 Ma into middle crustal levels (15–20 km) of the same igneous arc (Janousˇek et al.,2006).After a significant time gap, subduction-relatedmembers of the CBPC were emplaced, including thenormal calc-alkaline gabbros, quartz diorites and tonalitesof the Sa´zava suite (354.1 ± 3.5 Ma; Janousˇeket al., 2004a) and the high-K calc-alkaline, mainlygranodioritic Blatna´ suite (347 + 4 ⁄ )3 Ma; Do¨ rr &Zulauf, 2010; 346.4 ± 1.1 Ma; Janousˇek et al., 2010)with associated monzonitic bodies. Finally, furthereast, commonly in association with HP granulite massifsand high-grade Gfo¨ hl orthogneisses, syn-deformationalor post-tectonic intrusions of earlyCarboniferous (343–336 Ma) (ultra-) potassic rockswere emplaced (Holub, 1997; Holub et al., 1997;Janousˇek et al., 2003; Verner et al., 2008; Kotkova´et al., 2010; Kusiak et al., 2010). The time lag betweenthe end of normal calc-alkaline magmatism (c. 354 Ma)and the onset of intrusion of K-rich magmas (c.346 Ma) can be related to the transition from oceanplate subduction to continental underthrusting.PETROLOGY AND GEOCHEMISTRY OFMOLDANUBIAN GRANULITES AND POTASSICMAGMATIC ROCKSA peculiar feature of the Moldanubian domain in theBohemian Massif is an intimate spatial and temporalassociation between felsic HP, kyanite–garnet granulitesand large ultrapotassic plutonic bodies (Janousˇek& Holub, 2007). Whereas the granulites tend to bedepleted in the radioactive elements U, Th and K, the(ultra-) potassic rocks are characterized by strongenrichment in these elements that shows clearly in theradiometric map (Fig. 4). Thus, understanding of theK, U and Th depletion ⁄ enrichment in individual rocktypes at various stages of the Viséan HP metamorphismand igneous activity seems to be a key to decipheringthe thermal history of the Moldanubianorogenic crust.Moldanubian HP granulitesHigh-pressure granulites represent a voluminous andubiquitous component of the Gfo¨ hl Assemblage inboth Austria and the Czech Republic. The most typicalare felsic types consisting essentially of garnet, quartzand hypersolvus feldspar, and commonly containingkyanite. Rutile, zircon, apatite, ilmenite ± monaziteare the common accessories (OÕBrien & Ro¨ tzler, 2003).The felsic Moldanubian granulites were considered asformer rhyolites ⁄ granites that acquired their highgrademetamorphic character during the Variscancollision and which suffered only limited HP melting(Fiala et al., 1987a; Vellmer, 1992; Janousˇek et al.,2004b). Indeed, some granulites in rare domains thatescaped later mylonitization ⁄ recrystallization aremigmatitic in appearance (Franeˇk et al., 2006). On theother hand, other authors (Vra´na & Jakesˇ, 1982; Jakesˇ,1997; Kotkova´ & Harley, 1999) suggested that thefelsic granulites represent HP granitic liquids thatformed and separated from their source during theVariscan metamorphic cycle. There seem to be severalarguments against such a model (Janousˇek et al.,2004b): (i) concentrations of Zr are far too lowcompared to calculated saturation levels at ‡900 °C,coupled with significant, mainly Ordovician zirconinheritance; (ii) consistently low (750 °C) zircon andmonazite saturation temperatures, whereas pre-Variscaninheritance is by no means rare, documenting thatsaturation was reached (Janousˇek, 2006); (iii) preservationof Ordovician–Silurian whole rock and thin slabRb–Sr ages corresponding to the principal inheritedcomponent in granulite zircon; and (iv) high heavy rareearth element + Y contents, ruling out the presence oflarge amounts of garnet in the residue and thus indicatinga rather low-P melting.Ó 2010 Blackwell Publishing Ltd186

HEAT SOURCES AND EXHUMATION MECHANISMS 85Fig. 4. Radiometric map of the south-eastern Bohemian Massif with isolines of natural air absorbed dose rate (nGy h )1 ). Themain bodies of ultrapotassic rocks (durbachite series and melasyenitoids sensu Holub, 1997) that show very high radioactivity areidentified by name. The Variscan granites (crosses) and HP granulites (dots) are outlined. Source: Czech Geological Survey MapServer, http://www.geology.cz.The most peculiar feature of the felsic Moldanubiangranulites is the lack of LILE depletion, except for Cs,U and Th (Fiala et al., 1987a,b; Janousˇek et al.,2004b). Thus, unlike in many other granulite terranesworldwide (Rudnick & Presper, 1990; Rudnick & Gao,2003), prograde metamorphism appears to have beenlargely isochemical. As shown by Janousˇek et al.(2004b), the composition of felsic Moldanubian granulitesmatches well with felsic Ordovician–Silurianmetaigneous rocks from the Saxothuringian domain,for instance orthogneisses and meta-rhyolites from theFichtelgebirge. The similarities include whole-rockgeochemistry (excluding the most mobile elements Cs,Rb, Th, U, Pb and Li – Fig. 5a), Sr–Nd isotopiccompositions and protolith ages of c. 480–455 Ma(Siebel et al., 1997; Wiegand, 1997), forming animportant maximum within the spectrum of inheritedages in the granulites.A good candidate for complementary small-volume,HP–HT (>900 °C) melt that managed to separatefrom the calc-alkaline granulites are the rare hyperpotassicgranulites from Plesˇovice, Blansky´ les Massif(Vra´na, 1989; Janousˇek et al., 2007). These rocks,which are composed essentially of K-feldspar, garnet,zircon and apatite, show rather extreme geochemicalcompositions (e.g. strong enrichments in Cs, Ba, Rb,U, Th, K, P and Zr; Fig. 5b). They yielded U–Pb zirconages of 338 ± 1 Ma (Aftalion et al., 1989) and337.13 ± 0.37 Ma (Sla´ma et al., 2008), which is closeto the established best estimate of the HP metamorphicclimax in the granulite massifs (c. 340 Ma – see above).Moreover, rare coeval glimmerite veins in peridotitefragments enclosed by granulite bodies show highconcentrations of LILE, U and Th, accompanied bylow Rb ⁄ Cs and K ⁄ Rb ratios as well as low HFSEcontents (Fig. 5c); the Sr and Nd isotopic compositionsoverlap with the granulites (Becker et al., 1999).The glimmerites were interpreted as having crystallizedfrom an ultrapotassic, F-rich aqueous-carbonic fluid,bearing a direct witness for the HP–HT devolatilizationof granulite massifs.Viséan (ultra-) potassic magmatism in the MoldanubiandomainIn the Moldanubian domain of the Bohemian Massif,relatively large volumes of (ultra-) potassic plutonicrocks constitute several plutons and stocks spatiallyassociated to granitoids of the Central Bohemian PlutonicComplex, the Moldanubian Plutonic Complex,Ó 2010 Blackwell Publishing Ltd187

86 O. LEXA ET AL.(a)(b)(c)and high-grade metamorphic rocks of the Gfo¨ hlAssemblage, most notably felsic HP granulites. Thisprominent group represents the late-syntectonic durbachiteseries (e.g. Cˇertovo brˇemeno, Trˇebı´cˇ and Knı´žecı´stolec intrusions – Zˇa´k et al., 2005; Verner et al., 2008)(Fig. 4) of quartz melasyenites to melagranites withhydrous ferromagnesian minerals, Mg-rich biotite andactinolitic amphibole. These are mainly coarsely porphyriticrocks that contain abundant K-feldspar phenocrystsand ultrapotassic mafic microgranularenclaves. The spatially associated, less deformed, oreven post-tectonic, biotite–two-pyroxene melasyenitesto melagranites (Ta´bor and Jihlava intrusions) arecharacterized by a ÔdrierÕ ferromagnesian mineralassemblage of orthopyroxene, clinopyroxene andMg-biotite and lack the porphyritic texture (Ža´k et al.,2005).The petrogenesis of the Moldanubian (ultra-)potassic igneous rocks has been a matter of debate aseven the basic, Mg and K-rich members or primitivelamprophyric dykes have a mixed geochemical character.While their high contents of Cr and Ni with highmg# point to derivation from an olivine-rich source(mantle peridotite), the elevated concentrations of U,Th, light rare earth element (LREE) and LILE,depletion in Ti, Nb and Ta (Fig. 5d) and highK 2 O ⁄ Na 2 O and Rb ⁄ Sr ratios apparently contradict amantle origin (Holub, 1997; Janousˇek & Holub, 2007).These features led several authors to invoke partialmelting of anomalous (LILE- and LREE-enriched)lithospheric mantle domains (e.g. Janousˇek et al.,1995; Holub, 1997; Wenzel et al., 1997, 2000; Janousˇek& Holub, 2007), followed by mixing with lower crustalleucogranitic melts (Holub, 1997; Gerdes et al., 2000).In any case, the crustal-like Sr–Nd isotopic signaturescannot be reconciled solely by crustal assimilation⁄ contamination during the ascent of any primitive,mantle-derived magmas and require contamination bythe subducted continental crust directly in the source(d)Fig. 5. Multi-element variation diagrams. (a) Box and percentileplots for felsic (SiO 2 >70 wt%) Moldanubian granulites,normalized to an average of felsic metaigneous rocks from theFichtelgebirge (Siebel et al., 1997; Wiegand, 1997). See Janousˇeket al. (2004b and references therein) for the data set. The distributionof each of the normalized trace-element contents isplotted as irregular polygons, the width of which at any givenheight is proportional to the empirical cumulative distribution.As in box plots, the median, 25th and 75th percentiles aremarked with horizontal lines across the box. Compositions ofthe hyperpotassic Plesˇovice granulites (data from Janousˇeket al., 2007) (b) and Lower Austrian glimmerite veins (Beckeret al., 1999) (c) normalized by an average of felsic (SiO 2>70 wt%) HP Moldanubian granulites, except Lisˇov (takenfrom Janousˇek et al., 2004b). (d) Multi-element variationdiagram for selected ultrapotassic rocks (durbachite series andsyenitoids) from the Moldanubian Zone of the BohemianMassif (Janousˇek & Holub, 2007) normalized by N-MORB andthen adjusted to Yb N = 1 to minimize the effects of fractionalcrystallization (Pearce & Stern, 2006).Ó 2010 Blackwell Publishing Ltd188

HEAT SOURCES AND EXHUMATION MECHANISMS 87(Janousˇek & Holub, 2007). As noted by the sameauthors, multi-element plots for the Moldanubian felsicgranulites (Fig. 5a) and ultrapotassic rocks(Fig. 5d) are largely mutually complementary. Themost striking are the cases of Cs, Rb, Th, U, Pb and Li,which are impoverished in the felsic granulites butstrongly enriched in the ultrapotassic magmatic rocks.NUMERICAL MODELLING CONSTRAINTS ONHEAT SOURCES AND EMPLACEMENTMECHANISMS OF THE OROGENIC LOWERCRUSTTo constrain the heat source driving the tectonic processesof vertical extrusion three contrasting, but notmutually exclusive, hypotheses will be tested. Heatcould be generated by: (i) in situ decay of radioactiveelements contained in the FLC (U, Th and K); (ii)radioactive elements present in the metasomatized orcrustally contaminated mantle; and (iii) a large-scalethermal anomaly generated in the mantle as a result ofslab break off or mantle delamination.Fig. 6. Model geometry, initial lithology distribution, boundaryconditions and location of tracked samples used for numericalsimulations.The annual thermal productions were calculatedfollowing the method of Kramers et al. (2001), usingdecay constants and specific heat production datasummarized by van Schmus (1995; table 8). The pastannual heat production (lW kg )1 ) can be obtainedfrom the elemental concentrations of K, U and Thusing the equation:HðlW kg 1 623:45 10 2:638 10Þ¼Ke 0:554t þ The0:0495t324:03777 10 9:396852 10þ Ue 0:985t þe 0:1551t ;ð1Þwhere t represents age in Ga and K, U, Th are concentrationsin ppm.Model setupThe numerical model studies outlined below describe thetransient thermal evolution of a thickened orogenicdomain (Fig. 6) characterized by the presence of tectonicallyaccreted felsic rocks, including granulites,within orogenic lower crust. This FLC directly underliesa mafic layer which was added to Neoproterozoic crustduring early Palaeozoic crustal stretching and magmaticunder-plating. The presence of low density FLC below adense mafic layer introduces significant gravitationalinstability within the lower crust (Gerya et al., 2001)which could trigger crustal diapirism (Ramberg, 1981;Perchuk, 1989) and perturbate the thermal field. We usethermal and dynamic numerical models to examine therole of high radioactive heat production located in theFLC as a main candidate triggering the gravitationalinstability as a result of pronounced progressive generationof heat and subsequent change in density becauseof thermal expansion.The model is set up to allow the definition of differentmaterial domains with different thermal and mechanicalproperties (Table 2) on high-resolution Lagrangianmarkers initially arranged in a rectangular grid (Gerya& Yuen, 2003). Properties are mapped to a Eulerianstaggered grid where governing equations are solved fortemperature change (DT) and velocity. In each time,step markers are advected according to the updatedvelocity field and all temperature-dependent variablesTable 1. Calculated radioactive heat production values for Fichtelgebirge metaigneous rocks, felsic Moldanubian granulites andMoldanubian peridotite (present and at 340 Ma).Rock type Age (Ma) Density (kg m )3 ) Concentrations (ppm) A (lW m )3 )K Th U40 K232 Th235 U238 U Present PastFichtelgebirge 340 2700 38642.550 13.00 9.00 7.3421 13.00 0.064 8.935 3.670 3.920Moldanubian granulites 340 2750 38601.045 2.10 1.00 7.3342 2.10 0.007 0.993 0.788 0.885Hornı´ Bory peridotite 340 3200 940.537 0.11 0.57 0.1787 0.11 0.004 0.568 0.176 0.214Data sources for averaged whole-rock compositions: Fichtebeirge metaigneous rocks: Siebel et al. (1997), Wiegand (1997).Felsic Moldanubian granulites (SiO 2 > 70 wt%): Janousˇek et al. (2004b and references therein).Hornı´ Bory peridotite: Ackerman et al. (2009).Ó 2010 Blackwell Publishing Ltd189

88 O. LEXA ET AL.Fig. 7. Results of 1D static thermal models to compare temperature evolution controlled by different radioactive heat production offelsic lower crust (FLC; curves labelled 1–6 according to heat production) and radioactive heat production within the lithosphericmantle (dotted line) with thermal evolution caused by lithospheric delamination at 90 km depth (dashed line).such as thermal diffusivity, specific heat, viscosity anddensity are recalculated according to temperaturedependence for each marker (see Appendix).The geometry of the model, the distribution ofmaterial layers and the boundary conditions of thenumerical simulations are depicted in Fig. 6. The initiallayered geometry introduces an artificial perturbationin the interface between felsic and mafic lowercrust to allow immediate relaxation of gravitationalinstability in the central part of the computationaldomain. The initial temperature distribution is calculatedas the steady-state solution of Eqn A.3, withheat sources located only in the upper crust. Thelower crustal radioactive heat production is accountedonly for transient development. One of theimportant features of the model is the ability toprogressively eliminate sources of radioactive heatproduction according to the temperature achieved. Asargued above using the available geochemical data, ata certain stage of the evolution of the lower crust, mostof the radioactive elements, in particular U and Th,were stripped from the felsic granulites into partialmelt or ÔfluidÕ and transported, together with K-richmagmas, into the middle–upper crust. In the models,two values are used to cut-off heat production tosimulate radioactive element evacuation via fluid or viamelt.Heat sourcesThe effect of such behaviour on thermal evolution wasfirst examined in terms of a 1D static thermal model(Fig. 7). A sharp change in heating rate in Fig. 7ccorresponds to the time when heat production isswitched off in most of the lower crust. To compare thescale and magnitude of temperature change as a resultof processes like delamination or heat productionwithin lithospheric mantle, the plot is overlaid withresults of two additional numerical simulations. Thedotted line represents the results of a 1D model tosimulate production of heat within a >60 km thicklithospheric mantle with a radioactive heat productionof 0.2 lW m )3 calculated on the basis of the wholerockgeochemical data of Ackerman et al. (2009) forthe Hornı´ Bory garnet peridotite (Table 1). It is evidentthat even such an enriched mantle cannot providesufficient heat to be responsible for the significantincrease of temperature within the lower crust (Fig. 7c).Similarly, we argue that the process of lithospheredelamination (simulated by instantaneous replacementof lithospheric mantle below 90 km by asthenospherein the model) cannot provide the necessary heat input.Results related to the mantle delamination process areshown by the dashed line in Fig. 7.A series of numerical experiments was set up to studythe influence of radioactive heat production locatedwithin the FLC (Fig. 7). Our calculations show thatthe temperature required for partial melting of micabearingfelsic crust located at a depth of 70 km (850–900 °C) are reached after 20 Myr for radioactive heatproduction of 2 lW m )3 and in 7 Myr for radioactiveheat production of 4 lW m )3 . At 60 km depth, whichis the assumed upper limit of the felsic layer, themelting temperature is reached in >50 Myr and inÓ 2010 Blackwell Publishing Ltd190

HEAT SOURCES AND EXHUMATION MECHANISMS 89Table 2. Mechanical properties (density, thermal expansivity and coefficients of temperature-dependent viscosity) of individuallithologies used for numerical simulations.Material Description Reference density Thermal expansion coefficient Temperature-dependent viscosity range andcoefficientsRadioactive heat productionq0 (kg m )3 ) a Effect. viscosity (Pa s )1 ) C1 C2 Hr (lW m )3 )UC Upper crust 2700 0 10 22 Viscosity const. 2 · 10 )6MCs Middle crust (schists) 2800 2 · 10 )5 1.5 · 10 20 to 2.5 · 10 19 10 16 6000 0MCg Middle crust (gneisses) 2800 2 · 10 )5 2.5 · 10 20 to 3.5 · 10 19 10 17 6000 0MLC Mafic lower crust 2950 0 2 · 10 21 to 5 · 10 20 10 18 7000 0FLC Felsic lower crust 2750 3 · 10 )5 1.5 · 10 19 to 6 · 10 18 10 17 5000 2 · 10 )6 to 8 · 10 )6M Lithospheric mantle 3300 0 5 · 10 21 Viscosity const. 0 (2 · 10 )7 )c. 20 Myr for the two radioactive heat productionvalues, respectively. Keeping in mind the time-scalesof magmatic and metamorphic events related toPalaeotethys subduction discussed above, the time of5–15 Myr available between the arrival of continentalcrust into the subduction zone (354–346 Ma) and themetamorphic climax (c. 340 Ma) corresponds to thethermal incubation time estimated for radiogenic heatproduction of 4 lW m )3 .Results of 2D modellingThe distribution of densities and viscosities for allmodels generally resulted in the development of adiapiric structure located at the introduced perturbation.The P–T of selected samples (samples 1–3 arewithin lower crust, sample 4 in mafic layer and sample5 in middle crust; for locations, see Fig. 6), trackedduring the model evolution are plotted on Figs 8–10.Similar to the static 1D models, the simulations show aclear relation between the initial temperature increase(20–150 °C) and radioactive heat production withinlower crust.Two types of evolution have been calculated: (i) adiapiric structure formed because of mantle heatsource (heat production 0.2 lW m )3 ); and (ii) a diapiricstructure caused by radioactive heat productionwithin the FLC for radioactive heat productivitiesranging from 1 to 6 lW m )3 . The results of thenumerical simulations for the case of mantle heatproduction are shown in Fig. 8. For the case ofradioactive heat production within the FLC, it wasterminated at 900 and 1000 °C. These results are presentedtogether with a simulation in which the radioactiveheat production was not switched off (Fig. 9)allowing a direct comparison of differences in P–Tevolution (Fig. 10).Several major conclusions can be drawn from theseresults. The diapir reflecting the mantle heat source(Fig. 8) exhibits a typical bell shape during the first30 Myr and most importantly shows contrasting P–Tevolution for samples located in different units andinitial depths. Whereas the upper part of the mantle(Fig. 7d) is almost isobarically heated, the sampleslocated in the felsic crust, mafic lower crustal layersand the middle crust reveal relatively slow exhumationand moderate heating associated with the diapirgrowth.In contrast, experiments assuming high radiogenicheat production show typical bollard type diapirs.After 10–20 Myr imposed gravitational instability andvariable radioactive heat production within lowercrust, the viscosity of the overlying mafic layer is significantlyreduced allowing relatively fast exchange ofmaterial and development of the central diapir. Thereare differences in rates of vertical exchange betweenindividual simulations as increased heat productioncause increase of buoyancy forces and decrease ofviscosity of the diapir surroundings (samples 4 and 5).The growth of bollard type diapirs is associated witheither isothermal decompression or important coolinglinked to diapir growth for samples located in deepfelsic lower crustal layer. These are indeed the P–Tevolutions retrieved from for Bohemian Massif granulites(Sˇtı´pska´ et al., 2004; Racek et al., 2006;Tajcˇmanova´ et al., 2006). The samples located originallyhigher in the column and at the middle crustallevels reveal important heating associated with exhumation,which is also in accord with recent petrologicalstudies (Racek et al., 2006; Sˇtı´pska´ et al., 2008). Theother important consequence is that rocks fromany original position show convergence of P–T conditionsafter exhumation to mid-crustal depths. Thetime-scales of heating (10–20 Myr) and exhumation(5–10 Myr) calculated in this model are also in agreementwith petrological and geochronological datareported by Sˇtı´pska´ et al. (2004) and Tajcˇmanova´ et al.(2006, 2010). It should be noted that time-scales of ourmodels are directly controlled by the rheology of thematerials, which in our simulations is significantlysimplified (Eqn A.3, Table 2).Closer inspection of Fig. 10 confirms that the best fitof modern petrological and geochronological data withcalculated P–T paths is with an initial radioactive heatproduction of 4 lW m )3 . Here, the maximum temperaturesattained at the bottom of thickened crust are950 °C, while samples located in the central part ofthe felsic crustal column reach a maximum 800 °C atthe thermal peak. These values may reconcile modernTHERMOCALCmodelling data of Sˇtı´pska´ & Powell(2005b), Tajcˇmanova´ et al. (2006) and Racek et al.(2008) who reported peak temperatures 800 °C withÓ 2010 Blackwell Publishing Ltd191

90 O. LEXA ET AL.Fig. 8. Results of numerical simulation of radioactive heat production within the lithospheric mantle showing distribution oflithologies and temperature field developed during 50 Myr. Light shade of the mantle colour marks the asthenosphere (adiabaticgeotherm and no heat production).those of OÕBrien (2000) and Cooke & OÕBrien (2001)who reached significantly higher peak temperatureestimates. Higher heat production would produce significantlyhigher peak temperatures of nearly 1000 °Cmaintained even at pressures as low as 12–13 kbar,which is in contradiction with modern petrologicalstudies (see Schulmann et al., 2008 for review).Other important information comes from the temperaturedistribution in the core of the diapiric structure.The diagrams calculated for elevated radioactiveheat production show areas where the temperaturecondition of radioactive heat production switch off(i.e. partial melting associated with release of radioactiveelements) was attained during the evolution (thepinkish colour inside the diapirs in Fig. 9).As shown above, the metaigneous rocks fromFichtelgebirge are thought to be the best candidates forprecursors of the Moldanubian felsic granulites.Therefore, these rocks are interpreted as material thatcould have formed the felsic lower crustal layer, whichwas subsequently heated, partially melted andtransformed to typical felsic Moldanubian granulites.Ó 2010 Blackwell Publishing Ltd192

HEAT SOURCES AND EXHUMATION MECHANISMS 91Fig. 9. Results of numerical simulations showing distribution of lithologies and temperature field developed after 20 Myr. Light shadeof the lower crustal colour marks places where removal of radioactive heat production occurred while light shade of the mantle colourmarks asthenosphere (adiabatic geotherm and no heat production). First column shows results for low radioactive heat production andlithology and temperature fields are identical for all switch-off conditions as they are not reached during 20 Myr. The second columngives results for no switch-off condition, whereas the third and fourth columns shows results of simulations with 900 and 1000 °Cswitch off for the lower crustal layer.The metaigneous rocks from Fichtelgebirge yield anaverage radioactive heat production of 3.9 lW m )3obtained from the average elemental concentrations ofK, U and Th (Siebel et al., 1997; Wiegand, 1997) usingEqn 1 (Table 1). However, the radioactive heat productionfor the average felsic granulites (SiO 2 > 70,Table 1) is extremely low (0.9 lW m )3 ), suggestingthat the radioactive elements were mostly lost duringÓ 2010 Blackwell Publishing Ltd193

92 O. LEXA ET AL.Fig. 10. P–T evolution of six tracked samples for models with different radioactive heat production values for the felsic lowercrust (FLC). Samples 1, 2 and 3 were located in felsic lower crust, sample 4 within the mafic layer (MLC) and samples 5 and 6 arelocated in the middle crust. (a) No radioactive elements removed, (b) 900 °C and (c) 1000 °C threshold for radioactive elementremoval from the lower crust. Circles mark 10 Myr time step.the partial melting connected with the granulite faciesmetamorphism.In conclusion, the thermal structure which fits bestthe petrological and geochronological data obtained sofar from the Moldanubian granulite massifs is thatcalculated with an initial radioactive heat productionof 4 lW m )3 located in the felsic layer. The P–T–tevolutions of the AGB (mafic lower crust) and of theoverlying mid-crustal rocks are also in good agreementwith the model results, despite the fact that the simulatedevolution represents only the initial part of acomplex polyphase exhumation.DISCUSSIONIn this article, we discuss a particular structural patternin the Variscan orogenic root in which orogenic lowercrust composed of felsic granulites of OrdovicianÓ 2010 Blackwell Publishing Ltd194

HEAT SOURCES AND EXHUMATION MECHANISMS 93protolith age forming cores of domes that are separatedfrom mid-crustal Neoproterozoic and Palaeozoicmetasedimentary rocks in synclines by a late Ordovician–Silurianmetabasic layer. We argue that the originof these structures was related to diapiric materialexchange within the orogenic lower crust. The keyelement in deciphering the Viséan development is theoccurrence of FLC underneath dense mafic crust asdepicted by geology and geophysics. To understandthe formation of this peculiar lithological sandwich thepossible emplacement models for the FLC at the bottomof the root need to be discussed. Subsequently, weshall address the particular role of this felsic layer forthe origin of the felsic granulite–(ultra-) potassicmagma association. Finally, a model of gravity inversionis discussed together with thermal consequencesfor P–T–t paths of rocks in different positions withrespect to the diapiric structure.Model of relamination of felsic crust to early Palaeozoicmafic lower crustThe relamination of FLC has been proposed as analternative to the model of Chemenda et al. (2000) toexplain the structure of the Tibetan Plateau. TheTibetan Plateau is characterized by an exceptionallylarge gravity low indicating dominantly a felsic rootunderplated by Indian felsic crust, the density of whichcorresponds to felsic granulite at a pressure of 20 kbar(Hetényi et al., 2007). Indeed, Le Pichon et al. (1997)argued that the high topography of the Tibetan Plateauis due to presence of low density granulites atdepth. A similar gravity low is typical of the AltiplanoPlateau in central Andes, having been interpreted as aresult of underthrusting of the Brazilian crust underneaththe Andean root (Oncken et al., 2006). Thegravity anomaly associated with the Moldanubiandomain resembles remarkably the Tibetan and Altiplanoplateaux density structure, which in this case isreduced by subsequent isostatic reequilibration (Burget al., 1994). The common denominator of all thesegeophysical observations is the occurrence of felsiccrust at lower crustal depths. This is explained either ascontinental crust underthrusting thin lithosphericmantle (Chemenda et al., 2000) or directly by influx offelsic crust into the orogenic root at Moho depth liftingthe original lower crust and depressing the mantlelithosphere (Behr, 1978; Plesch & Oncken, 1999;Avouac, 2008).In the western Bohemian Massif, the influx of ductilelower crust at granulite ⁄ eclogite facies conditions wasproposed by OÕBrien (2000). In Fig. 11a,b, the influx ofSaxothuringian crust into the root domain is visualizedin the form of a 10-km wide channel splitting the earlyPalaeozoic mafic lower crust from lithospheric mantle,whereas the other part of the continental crust iscontinuously subducted, contaminating the localmantle and sampling the mantle lithosphere. The termrelamination (Hacker et al., 2007) is accepted as beingsuitable for the addition of low density crust underneaththe dense root, in contrast to the term delaminationto describe loss of heavy root material and itsreplacement by the asthenosphere.As discussed before, oceanic subduction had to haveceased by 354–346 Ma to be replaced by continentalunderthrusting. Incidentally, a Sm–Nd age of354 ± 6 Ma was determined by dating of calcium-richcores of garnet from the South Bohemian granulitesindicating the onset of eclogitization of continentalcrust at this time (Prince et al., 2000). If true, such ascenario allows a 5–15 Myr period to c. 340 Ma, thegranulite facies metamorphic climax, for the subductedcontinental crust to thermally incubate and elevate theorogenic geotherm (England & Thompson, 1984).Based on P–T estimates, crustal thickening had toproduce a 70-km thick crust at this time (Fig. 11c).Melting of this continental crust started at c. 345 Ma,as indicated by high-K calc-alkaline magmatism of theBlatna´ suite. The maximum melt production at boththe base of this crust and in the underlying mantlelithosphere occurred during, or soon after, the HPmetamorphic climax at c. 340 Ma, as indicated byintrusions of ultrapotassic syenites at mid- to highcrustallevels.The origin of the felsic granulite–ultrapotassic plutonic rockassociation and the heat sourceRoberts & Finger (1997) proposed that heating ofrelatively refractory felsic metaigneous rocks, the likelysource of the felsic granulites, to temperatures as highas 1000 °C would result in production of 5–15 vol.%of partial melt. This notion was confirmed by thermodynamicmodelling of Janousˇek et al. (2004b), whosuggested that the initial melting, limited by micaavailability, would not exceed 10 vol.% for any of theprospective granulite decompression paths and rapidincrease of the melt fraction would occur only at1100 °C. However, it is important to correctly assessthe melt production for the most typical felsic granulites.The most likely protolith for the felsic granulitesis considered to be granite or orthogneiss of theÔFichtelgebirge chemistryÕ that must have lost somemelt during metamorphic evolution in order to preservethe high-pressure assemblage. Therefore, toestimate the degree of melt loss, pseudosections werecalculated using THERMOCALC for a granulite with theoldest known preserved fabric from the Blansky´ lesMassif (sample H296-S1 from Franeˇk et al., 2011b; fordetails, see Appendix).The sample modelled contains relict porphyroclastsof perthitic feldspar with inclusions of kyanite, garnetwith 24 mol.% of grossular, rutile and biotite, whichconstrains the peak P–T conditions to 16–18 kbarand 860 °C (Fig. 12a; Franeˇk et al., 2011b). It isnecessary to add 7 mol.% of melt into the rockcomposition to obtain assemblages containing biotiteand muscovite without garnet or alumosilicate atÓ 2010 Blackwell Publishing Ltd195

94 O. LEXA ET AL.(a)(b)(c)(d)Fig. 11. Model proposed for the tectonic evolution of the orogenic root domain in the Bohemian Massif. (a) Model of continentalunderthrusting with development of the Barrandian forearc region, CBPC magmatic arc and backarc region represents the futureMoldanubian domain. The position of the Palaeotethys suture is indicated as the future Maria´nské La´zneˇ Complex. (b) Relaminationmodel with part of the allochthonous Saxothuringian crust injected between the Moho and the continental lithosphere. The other partis subducted thereby producing metasomatism of the overlying mantle. (c) Crustal thickening of the former backarc domain(Schulmann et al., 2005, 2009). (d) Vertical extrusion and gravity redistribution of relaminated Saxothuringian crust at the end ofVariscan orogeny.Ó 2010 Blackwell Publishing Ltd196

HEAT SOURCES AND EXHUMATION MECHANISMS 95(a) (b)Fig. 12. (a) Pseudosection for granulite sample H296-S1 (Blanský les Massif) to show the P–T path from the peak assemblage of g–ky–pl–ksp–q–liq–ru at 16 kbar and 860 °C bydecompression to 8 kbar and 820 °C (modified from Franeˇk et al., 2011b). Molar percentage of melt is very low and allows the preservation of the peak assemblage ondecompression and cooling. (b) Pseudosection for sample H296-S1 with 7 mol.% of melt added allows the stability of the mu–bi–pl–ksp–q assemblage at subsolidus conditions.Such a reconstructed granite whole-rock and mineral composition predicts 8 mol.% of melt at peak P–T conditions and 12 mol.% of melt at 8 kbar and 860 °C. Some melt mustbe lost to preserve the peak g–ky–pl–ksp–q–liq–ru assemblage.Ó 2010 Blackwell Publishing Ltd197

96 O. LEXA ET AL.subsolidus conditions. Such a reconstructed granitemineral composition calculated at 650 °C and 9 kbar is8 mol.% muscovite, 6.5 mol.% biotite, 39 mol.%quartz, 30 mol.% plagioclase and 16 mol.% K-feldspar(Fig. 12b). On heating, the melt proportion forthe reconstructed granite at 16 kbar and 860 °C is8 mol.%, and on isothermal decompression to 8 kbarand 860 °C it increases to 13 mol.%. The calculationspredict 14 mol.% of melt at 16 kbar and 1000 °C thatincreases to 37 mol.% at 8 kbar and 1000 °C, confirmingthe earlier estimates by Roberts & Finger(1997) and Janousˇek et al.(2004b).The modelling demonstrates that heating of graniticcrust with a restricted amount of muscovite and biotitecan generate a garnet-bearing residue and up to7 mol.% of melt at temperatures 900 °C. This temperaturewas used as a threshold for melt loss relatedremoval of radioactive elements and the switching offof the heat production in lower crustal rocks in thenumerical experiments.The detailed microstructural studies of Franěk et al.(2006, 2011b) showed that the origin of the granulitemicrostructure is related to significant deformation,dynamic recrystallization and viscous flow during thevertical extrusion stage. Before the vertical extrusionstage, the granulite facies rocks acquired a metamorphicfabric in a rather static environment. The deformationoccurred via diffusion-assisted grain boundarysliding, which is an efficient mechanism to extract meltfrom a continuously deforming source by a combinationof dynamic dilation and compaction (Za´vadaet al., 2007). The mechanism of dynamically developedporosity and melt extraction was discussed by Hasalova´et al. (2008a), who showed that a small amount ofmelt may be efficiently extracted from the source farbelow the rheologically critical melt percentagethreshold (>20–25 vol.% melt; Vigneresse et al.,1996). Therefore, we consider 900 °C to be a reasonabletemperature when melt can be extracted from theparent rock via the grain boundary sliding mechanismleaving behind a garnet-bearing mylonitic rock – theMoldanubian granulite. However, this processrequired significant deformation related to the verticalextrusion event to extract the melt. It is the pervasiveearly vertical fabric which is related to the efficient lossof melt (even at very small melt fraction), resulting insignificant depletion of the parental rocks in U, Th, Cs,Li ± Rb, but leaving the rest of the geochemical signatureunaffected.A consequence of melting and melt extraction is theremoval of a significant part of the radioactive elementbudget from the system. This is particularly true for Uand Th hosted by accessories such as zircon or monazite,which would resist fluid loss in course of theprogressive heating. On the other hand, they shoulddissolve readily even at low degrees of melting, as thepartial melt was likely to be rather corrosive, being hotand rich in alkalis with fluorine (Finger & Cooke,2004; Janousˇek et al., 2007). In granitic rocks, therelevant accessories are mostly enclosed in biotite (Bea,1996) that was undergoing melting. Moreover, as arguedby Watson et al. (1989) on theoretical grounds,the larger accessories are likely to be progressivelyconcentrated at grain boundaries in the course of highgrademetamorphism. Finally, there is a directevidence for the presence of low-degree, high-T, traceelement-richmelt in the felsic HP granulites, as thenewly grown metamorphic zircon and rutile are rich inU, Th and LREE or Zr and Nb, respectively (Finger &Cooke, 2004). Thus, the accessories hosting U and Thin the protolith to the HP Moldanubian granulitesseem to have been largely accessible to, and dissolvedin, the low-degree HP melt.The 1D thermal modelling has demonstrated that,for radioactive heat production of 4 lW m )3 , thetemperature threshold of 900 °C would be reached indeeply buried crust after 7 Myr (at 70 km) to 15 Myr(at 60 km). A radioactive heat production of4 lW m )3 is similar to the average radioactive heatproduction at c. 340 Ma calculated for the Fichtelgebirgemetaigneous crust, which is considered tocorrespond to the most appropriate protolith of felsicgranulites (Janousˇek & Holub, 2007). Moreover, themodelling shows that heat necessary for crustal meltingindeed could have been produced internally, within thetime frame allowed by the available geochronologicaldata (5–15 Myr).The heat source located at lower crustal depthswould also lead eventually to partial melting of theunderlying metasomatized and hydrated subcrustalmantle lithosphere. Melting of such an anomalous andfertile mantle source could have produced, soon afterthe HP metamorphic peak, ultrapotassic rocks withmixed crustal–mantle signatures. The effective removalof U and Th from the partially molten felsic metaigneousrocks (future felsic granulites) by melt extractionwould have grave consequences for the thermal evolutionof the whole system. First, the depletion ofradioactive elements from the granulite means that themelting process must have rapidly switched off. Second,the extracted melts could have partly mixed withenriched mantle-derived ultrapotassic magmas, whichinvaded the overlying partially molten crust, contributingto their further enrichment by U and Th.However, this mixing was probably volumetricallyrather insignificant and essentially different fromanother, large-scale hybridization event with S-typeleucogranitic magmas assumed for the durbachiteseries in Bohemia (Holub, 1997) and the Rastenbergsuite in Lower Austria (Gerdes et al., 2000). Thehybrid magmas finally intruded syn-tectonically orpost-tectonically at mid- to high-crustal levels in closespatial and temporal association with the HP granuliteand orthogneiss complexes so typical of the Variscanorogenic crust in the Moldanubian domain of theBohemian Massif (Schulmann et al., 2005; Zˇa´k et al.,2005; Tajcˇmanova´ et al., 2006; Janousˇek & Holub,2007).Ó 2010 Blackwell Publishing Ltd198

HEAT SOURCES AND EXHUMATION MECHANISMS 97Gravity overturns and 340 Ma crustal redistribution in theBohemian MassifGerya et al. (2001, 2002) in their pioneering work proposeda model of gravity redistribution of granodioriticrocks and overlying mafic crust using radioactive heatproduction localized in granodioritic crust. This progressiveheating of the FLC leads to its viscosity dropand density decrease triggering diapiric process. However,the time-scales required for gravity overturnsproposed by these authors are not compatible with thosereported for the Variscan orogeny (discussed before; seealso Schulmann et al., 2009). Therefore, the temperatureincrease must have been significantly more radicaland density contrasts higher than those predicted byGerya et al. (2002, 2004). This is consistent with thegravity inversion modelling results of Guy et al. (2010)who have shown that the density contrast betweenresidual FLC and an overlying thick gabbroic layer(Fig. 3) may be greater than that estimated by Geryaet al. (2001). Moreover, an additional source of gravitypotential may be seen in the presence of eclogites in thethickened Ordovician–Silurian AGB layer (50–60 km),overlying the felsic crust at a depth of 60–70 km (Sˇtípska´& Powell, 2005a,b).The scenario presented here (Fig. 11d) resembles theexplanation of the formation of migmatitic domes in hotorogenic belts driven by gravitational collapse ofthickened continental crust (Rey et al., 2001; Vanderhaeghe& Teyssier, 2001; Vanderhaeghe, 2009). All theseconceptual models suggest partial melting of the lowercrust to be a trigger mechanism for development ofgravitational instability in both fossil and modern orogenicbelts. The main difference of our model is inquantification of the gravitational instability, the rate ofthe heating and also the rate of the development of thediapiric structures. The initial position of less dense buthighly radioactive felsic material below more densemafic lower crust seems to be the necessary prerequisitefor driving the tectonic evolution of the Variscan orogenyin the Bohemian Massif. Importantly, in ournumerical models, the development of granulite domesis caused solely by gravitational instability. However,the structural data suggest that domes are initiated byfolding of the lower and mid-crustal interfaces (e.g.Sˇtípska´ et al., 2004; Schulmann et al., 2005) and thelinear distribution of elongate granulite belts indicatesthat lateral shortening was an important component(Burg et al., 2004) related to the growth of granulitedomes (Fig. 11d). Therefore, the combination of gravityand laterally forced extrusion of orogenic crust may leadto a development of gravity overturns more rapidlycompared with the numerical models presented herein.Our work shows that only a model of internalheating to drive the tectonic and gravity redistributionof orogenic lower crust can satisfactorily explain thetemporarily restricted orogenic event at c. 340 Ma,which is responsible for most of the Variscan tectonothermalevents reported so far in the BohemianMassif. Thus, the timing of growth of these laterallyforced diapirs seems to be connected with orogeniccollapse and the major plate reorganization of thewhole Variscan belt (Edel et al., 2003). The co-existenceof c. 340 Ma ultrapotassic plutons and extrudedfelsic HP granulites thus defines a key thermomechanicalevent, which probably signified a rheologicalcollapse of the whole Variscan belt in Europe (e.g.Rossi et al., 2009; Rubatto et al., 2010).CONCLUSIONSThe exhumation of c. 340 Ma felsic granulites in theBohemian Massif is interpreted in terms of tectonicallytriggered gravity redistribution of felsic orogenic lowercrust and high density mafic crust. The model showsthat radioactive heat production of 4 lW m )3 forlower crustal rocks, which is corroborated by calculatedvalues from likely protolith rocks, and the calculatedP–T–t evolution satisfy the thermal andgeochronological evolution of the Bohemian Massifgranulites. This radioactive heat production is typicalof Ordovician felsic igneous rocks in the Fichtelgebirge(Saxothuringian domain), which are believed to havebeen relaminated at the bottom of thickened continentalcrust during the early Vise´an continentalunderthrusting. The mutually complementary geochemicalcharacteristics of granulites and ultrapotassicmagmas highlight the shared thermomechanical historygoverned by radioactive heating of the lowercrustal layer, culminating in partial melting of both thefelsic lower crustal layer itself and its underlying fertile(metasomatized ⁄ contaminated) mantle lithosphere.Gravity-driven redistribution tectonics initiated byinternal heating is interpreted to be the principal agentcontrolling the rheological collapse of the Variscanorogenic crust at c. 340 Ma.ACKNOWLEDGEMENTSWe acknowledge grant MSM0021620855 from theMinistry of Education of the Czech Republic andinternal research funds from CNRS UMR 7615 forsalary and research support of Ondrej Lexa, and theFrench National Science Foundation ANR projectÔLFO in orogensÕ for additional research support. Wethank C. Clark and T. Gerya for their thoroughreviews and highly appreciate the comments andsuggestions, which significantly contributed toimproving the quality of the publication. M. Brown isacknowledged for careful editorial work.APPENDIXGoverning equations used for numerical modelling ofgravity overturnsTo simulate the thermal evolution, we use a numericalmodel based on the solution of the coupled equationsÓ 2010 Blackwell Publishing Ltd199

98 O. LEXA ET AL.of momentum (Eqn A.1) and energy (Eqn A.3), subjectto an incompressibility constraint (Eqn A.2) known asthe Boussinesq approximation (Hansen & Yuen, 2000):@r ij¼ @P qg i ; ðA:1Þ@x i @x i@v i@x i¼ 0;@T@t þ v @Ti ¼ 1 @ @TjqC p þ A ;@x i qC p @x i @x i qC pðA:2ÞðA:3Þwhere x denotes the coordinates in m, v velocity inms )1 , t time in s; r ij is stress tensor, P is pressure in Pa,g is gravitational acceleration (9.81 m s )2 ), T is temperaturein K, j denotes thermal diffusivity and C pdenotes the specific heat capacity. The constitutiverelationship between stress and strain is governed bythe transport coefficient g representing viscosity:r ij ¼ 2g_e ij ;ðA:4Þwhere _e ij is the strain-rate tensor in s )1 . For the purposeof this work, we used a simplified flow law withonly temperature-dependent viscosity according to thesimple exponential equation:g ¼ C 1 e C 2ð T Þ ;ðA:5Þwhere C 1 and C 2 are coefficients used to calculateeffective viscosity from prescribed range (see Table 2).Density q is given by equation of state:q ¼ q 0 ½1 aðT T 0 ÞŠ; ðA:6Þwhere a is the coefficient of thermal expansion and q 0is the reference density at reference temperature T 0 .Thermal diffusivity j and specific heat capacity C p arerecalculated according to temperature using the followingequations (Whittington et al., 2009):j ¼ 3:19 10 7 þ 1:214 10 6 eð 273:15285:2T Þ ; ðA:7Þ3:224 10 5 2:714 107C p ¼ 1538:39þTT 2 ; ðA:8Þderived from laser-flash analysis to provide realisticvalues for geologically relevant temperatures. Theseequations are solved for temperature and velocity anddescribe the fundamental physics required for modellingthe thermal evolution during crustal diapirism.Calculation methods used for thermodynamic modellingThe pseudosections were calculated using THERMOCALC3.30 (Powell et al., 1998) and the data set 5.5 (Holland& Powell, 1998; November 2003 upgrade), in the systemNa 2 O–CaO–K 2 O–FeO–MgO–Al 2 O 3 –SiO 2 –H 2 O–TiO 2 –O (NCKFMASHTO) with the biotite and meltmodels from White et al. (2007), garnet from Dieneret al. (2008), ilmenite from White et al. (2000), feldsparfrom Holland & Powell (2003), white mica fromCoggon & Holland (2002) and cordierite fromTHERMOCALC documentation (Powell & Holland, 2004).The analysed rock composition of sample H296 (inwt% SiO 2 =71.98, TiO 2 =0.42, Al 2 O 3 =13.53,FeO=2.1, MnO=0.03, MgO=0.73, CaO=1.93,Na 2 O=2.76, K 2 O=4.01, P 2 O 5 =0.15, H 2 O ) =0.22,H 2 O + =0.56, CO 2 =0.03), was modified for modellingby adding 1 mol.% of kyanite to enable a smallamount of aluminosilicate to be stable at the estimatedpeak metamorphic conditions, as is observed in thinsection.The reconstruction of a biotite–muscovite graniteprotolith requires several steps. It involves determinationof the H 2 O content in the final assemblage, considerationof open-system behaviour with respect tomelt, and modification of the whole-rock compositionby adding melt (White & Powell, 2002; White et al.,2004; Sˇtı´pska´ et al., 2008). Tracking of the P–T path istherefore undertaken backwards in time, from thematrix assemblage to the early prograde evolution andto the protolith mineralogical composition. Theamount of H 2 O for the modelling shown in Fig. 12a isset such that it allows the stability of the observedmatrix assemblage with garnet rim chemistry oncooling (not shown; see Franeˇk et al., 2011b; seeHasalova´ et al., 2008b for the approach followed). Asthe whole-rock composition is expected to changealong the P–T path as a result of loss of melt, it has tobe decided from which point on the P–T path the meltcomposition will be taken. Crossing the upper stabilityof muscovite causes an abrupt increase in melt proportionand is considered a likely condition for meltloss (White & Powell, 2002; White et al., 2004);therefore, melt composition reintegrated was undertakenat 16 kbar and 860 °C.REFERENCESAckerman, L., Jelínek, E., Medaris, G., Jezˇek, J., Siebel, W. &Strnad, L., 2009. Geochemistry of Fe-rich peridotites andassociated pyroxenites from Horní Bory, Bohemian Massif:insights into subduction-related melt-rock reactions. ChemicalGeology, 259, 152–167.Aftalion, M., Bowes, D. & Vrána, S., 1989. Early CarboniferousU-Pb zircon age for garnetiferous, perpotassic granulites,Blansky´ les massif, Czechoslovakia. Neues Jahrbuch fu¨rMineralogie-Monatshefte, 4, 145–152.Avouac, J.P., 2008. Dynamic processes in extensional andcompressional settings – mountain building: from earthquakesto geological deformation. In: Treatise on Geophysics, Volumes1-11 (ed. Schubert, G.), pp. 377–439. Elsevier, Amsterdam.Babusˇka, V., Fiala, J. & Plomerova´, J., 2010. Bottom to top lithospherestructure and evolution of western Eger Rift (CentralEurope). International Journal of Earth Sciences, 99, 891–907.Bea, F., 1996. Residence of REE, Y, Th and U in granites andcrustal protoliths; implications for the chemistry of crustalmelts. Journal of Petrology, 37, 521–552.Ó 2010 Blackwell Publishing Ltd200

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J. metamorphic Geol., 2011, 29, 53–78 doi:10.1111/j.1525-1314.2010.00903.xModel of syn-convergent extrusion of orogenic lower crust in thecore of the Variscan belt: implications for exhumation ofhigh-pressure rocks in large hot orogensJ. FRANĚK, 1,2 K. SCHULMANN, 3 O. LEXA, 1,2 Č.TOMEK 1 AND J.-B. EDEL 31 Czech Geological Survey, Klárov 3, 118 21 Prague, Czech Republic (jan.franek@geology.cz)2 Institute of Petrology and Structural Geology, Charles University, Albertov 6, 128 43 Prague, Czech Republic3 Institut de Physique du Globe de Strasbourg, IPGS – UMR 7516, CNRS et Université de Strasbourg (EOST), 1 Rue Blessig,67084 Strasbourg, FranceABSTRACTReflection seismic section, field structural analysis and gravimetric modelling of orogenic lower crust inthe core of a Carboniferous orogenic root reveal details of the polyphase process of exhumation.Subvertical amphibolite facies fabrics strike parallel to former plate margins that collided in the NW.The fabrics are developed in both mid-crustal and lower crustal high-pressure granulite units as a resultof intensive NW–SE intraroot horizontal shortening driven probably by the west-directed collision. Ingranulites, the steep fabrics originated as a result of extrusion of orogenic lower crust in a 20 km widevertical ascent channel from lower crustal depths at 350–340 Ma. The large granulite bodies preserveolder granulite facies fabrics documenting a two-stage evolution during the exhumation process. Surfaceexposures of granulites coincide with the absence of subhorizontal seismic reflectors at depth, suggestingpreservation of the 20 km wide subvertical tabular structure reaching Moho depths. Horizontalseismic reflectors surrounding the vertical channel structure corroborate a dominant flat migmatiticfabric developed in all tectonic units. This structural pattern is interpreted in terms of subhorizontalspreading of partially molten orogenic lower crust in mid-crustal levels (765 °C and 0.76 GPa) at 342–337 Ma. Large massifs of extruded and progressively dismembered felsic granulites disturbed midcrustalfabrics in the surrounding horizontally flowing partially molten crust. The horizontal mid-crustalflow resulted in collapse of the supra-crustal Teplá-Barrandian Unit (interpreted as the orogenic lid)along a large-scale crustal detachment above the extruded lower crustal dome. The presence of felsicgranulites at the bottom of the orogenic root is considered to be a key factor controlling the exhumationof orogenic lower crust in large hot orogens.Key words: exhumation; granulites; Moldanubian domain; orogenic lower crust; Variscan belt.INTRODUCTIONCurrent models for the exhumation of deep-seatedcrustal rocks often focus on processes along suturezones that offer a pre-existing discontinuity alongwhich the deep-seated rocks can be transported upwards.These models mimic a classical concept ofbuoyancy-driven exhumation of subducted crustal slicesalong the subduction channel (e.g. Chemendaet al., 1995) simulated numerically by Gerya &Stockhert (2006), Gerya et al. (2008) and Beaumontet al. (2009) for orogens characterized by continentalsubduction. In contrast, Platt (1993) proposed a cornerflow regime operating on a crustal scale between asubduction zone and a rigid buttress of the overridingplate, which evokes rise of lower crustal material alongits edge. Burg & Podladchikov (1999) offered numericalmodels of lithospheric-scale buckling, where thebuckled crust may undergo fragmentation via thrustzones, which can exhume particular blocks >10 kmvertically (Sokoutis et al., 2005).In many cases at the end of an orogeny the thickenedcrust undergoes rapid tectonic thinning rather thanslow erosion-related decrease of crustal thickness (e.g.Dewey et al., 1993). This process results from gravitydriven extensional collapse of the orogen causingdevelopment of pervasive flat crustal fabrics (Koyiet al., 1999). Additionally, heat redistribution leadingto weakening of the deep crustal rocks via partialmelting may facilitate collapse of a hot and thickenedorogenic root (Rey et al., 2001, 2009). Subsequentlateral ductile- to channel-flow, driven by a combinationof tectonic and overburden gravity force, mayresult in exhumation from high pressure (HP) conditionsas modelled by Beaumont et al. (2004) andJamieson et al. (2007).The Variscan orogen in centralEurope is characterized by numerous occurrences ofHP rocks located far from a suture. The HP rocks areÓ 2010 Blackwell Publishing Ltd 53205

54 J. FRANĚK ET AL.represented mostly by large bodies of HP felsic granuliteswhich form areas sometimes several 100 km 2 insize (OÕBrien & Carswell, 1993; Janousˇek & Holub,2007). These HP granulites are associated with otherHP rocks like mafic granulites, eclogites and garnetperidotites collectively forming the so-called Gfo¨ hlUnit (Fuchs, 1976). This unit does not form a continuousHP body, but occurs in three major belts, whichare surrounded by medium-grade rocks (Schulmannet al., 2009). The protolith of these rocks, their positionat lower crustal depths prior to exhumation andthe exhumation mechanisms represent major problemsof the Variscan belt (e.g. Behr, 1961, 1980). Schulmannet al. (2005) proposed a multistage model of rapidexhumation of orogenic lower crust associated withdevelopment of subvertical ascent channels triggeredby rapid amplification of initial instabilities in front ofan advancing continental buttress. Such a process isenabled by the extremely low viscosity of the Variscanorogenic lower crust at the bottom of the orogenic rootdue to its unusually hot thermal structure (e.g. OÕBrien,2008) combined with its felsic composition (Ru˚zˇeket al., 2007). This model is fundamentally similar tothat proposed by Weber & Behr (1983) who explainedthe exhumation of the Saxonian HP granulites byÔdiapiric foldingÕ. In their model, the deep granulitelayer tends to amplify and pierce through the weakermiddle crust during crustal shortening, to form largescalesteep folds bringing HP rocks to mid-crustallevels.The study region is located in the central part of theorogenic root, far from both the easterly continentalbuttress and westerly suture zone (Fig. 1). Here, weexamine a large portion of orogenic lower and middlecrust using both the published and unpublished part ofthe 9HR seismic line of Tomek et al. (1997). Theseismic profile is combined with a detailed structuralstudy along a SE–NW traverse evaluating mutualrelationships between middle orogenic crust and thethree major granulite massifs of South Bohemia, whichcompletes an earlier study of one of these massifs(Franeˇk et al., 2006). The profile line drawing is combinedwith detailed gravity forward modelling allowingestimation of a vertical ascent channel of orogeniclower crustal rocks. Finally, we discuss the developmentof flat-lying amphibolite facies fabrics throughoutthe traverse, which provides evidence of mid-crustalhorizontal flow. It is shown that this episode is connectedwith the collapse of the orogenic suprastructurerelated to final upwelling of a lower crustal dome.Despite disruption of the linear Variscan orogenictrend by late large-scale transcurrent shear zones (Edelet al., 2003), the wealth of quantitative geological datamakes the Bohemian Massif a suitable field laboratoryfor understanding processes related to exhumation oforogenic lower crust in large hot orogens. The fieldresults are suitable for testing the results from largescaleexhumation-related numerical models referencedabove.GEOLOGICAL SETTINGThe Bohemian Massif (Fig. 1) is traditionally dividedinto the Saxothuringian domain to the west, the Tepla´-Barrandian and Moldanubian domains in the centralpart of the Massif (Kossmat, 1927) and the Brunovistulian(Brunia) Neoproterozoic continent to the east(Dudek, 1980). Schulmann et al. (2005, 2009) interpretedthe Bohemian Massif as a Gondwana-derivedcollisional domain characterized by: (i) relicts of a twostageSE-directed subduction at the Saxothuringian–Tepla´-Barrandian boundary; (ii) a magmatic arcgenetically related to the subduction represented bythe Central Bohemian Plutonic Complex in the centre;and (iii) the rigid foreland represented by the Bruniamicroplate in the SE. In this concept, the Tepla´-Barrandian between the suture zone and the magmaticarc represents the fore-arc domain and the Moldanubiandomain between the magmatic arc and the Bruniamicroplate constitutes the shortened and thickenedintracontinental back-arc region. Large Variscanstrike-slip zones (e.g. the Elbe Fault Zone, seeFig. 1a,b) strike NW–SE and dismember the NNEtrending Variscan structure of the Bohemian Massif(Edel & Weber, 1995).Geology of the Saxothuringian, Teplá-Barrandian andMoldanubian domainsThe SE part of the Saxothuringian domain consists ofthe antiformal Erzgebirge Crystalline Complex, whichcan be divided into a lowermost para-autochtonousdomain that is overlain by crystalline nappes thatrecord peak HP metamorphism at 345–340 Ma(e.g. Franke, 2000; Konopa´sek & Schulmann, 2005).Towards the NW, the antiform passes into a Cambrianto Lower Carboniferous volcano-sedimentarysequence, which crops out in several large-scale latetectonic antiforms and synforms (Franke, 1993). Thesesupracrustal rocks are overthrust by oceanic crustalunits (Mu¨ nchberg, Frankenberg and Wildenfels klippen),metamorphosed at eclogite facies at c. 395–380 Ma (e.g. Dallmayer et al., 1995; Franke, 2000).The Saxonian granulite massif emerges as a large-scaleNE–SW elongated dome structure from below theSaxothuringian Basin (Du¨ rbaum et al., 1999) andconsists of HP felsic granulite, with subordinate lensesof mafic granulite and serpentinized peridotite (Ro¨ tzler& Romer, 2001). Kro¨ ner & Willner (1998) obtainedages of 485–470 Ma from zircon cores interpreted todate the protolith, while c. 340 Ma overgrowths wereinterpreted to date the time of peak metamorphism.The Tepla´-Barrandian domain is separated from theSaxothuringian domain by the Variscan SE-dippingsuture zone represented by the Maria´nske´ La´zněComplex (e.g. Zulauf et al., 1997). This unit consistsof serpentinites and metagabbros of Cambrian andOrdovician age (Timmermann et al., 2004), which werein part eclogitized before Devonian exhumation. TheÓ 2010 Blackwell Publishing Ltd206

EXTRUSIONOFLOWERCRUSTINVARISCANOROGEN 55(a)(b)Fig. 1. (a) Principal divisions of the Variscan chain. RH, Rhenohercynian domain; SX, Saxothuringian domain; MO, Moldanubiandomain. (b) Simplified geological map of the Bohemian Massif modified after Franke (2000).Tepla´-Barrandian domain consists of primitiveNeoproterozoic siltstones to greywackes and volcanicrocks deposited on an oceanic or transitional crust(e.g. Cháb & Pelc, 1973) and weakly metamorphosedduring the Cadomian orogeny (Kettner, 1918). ThisCadomian basement is unconformably overlain byCambrian and Ordovician to mid-Devonian sedimentarysequences that consist of greywackes, shales,sandstones and limestones. The Variscan tectonometamorphicprocesses are most pronounced in thewestern part, where they are restricted to 385–360 Ma(Fig. 2; Appendix S1). The boundary to the east withthe migmatites of the Moldanubian domain is markedby intense strike-slip deformation along the so-calledCentral Bohemian Shear Zone (Rajlich, 1988; Pitraet al., 1999; Scheuvens & Zulauf, 2000), however, thetrue contact is often masked by 375–336 Ma (Fig. 2)intrusions of the Central Bohemian Plutonic Complex(e.g. Zˇa´k et al., 2005a). The boundary betweenMoldanubian migmatites and the SW part of theTepla´-Barrandian domain is marked by the so-calledWest-Bohemian Shear Zone (Zulauf et al., 2002; Do¨ rr& Zulauf, 2008).The Moldanubian domain corresponds to theinternal orogenic root zone of the Variscan orogen(Suess, 1926). It is intruded by numerous Variscanplutons ranging from I-type (e.g. specifically K–Mgrich syenites or calc-alkaline arc-related rocks, Janousˇeket al., 2000) to S-type granitoids (Finger & Steyer,1995). The Moldanubian–Brunia boundary is markedby a several-kilometres-wide zone of highly deformedme´lange derived from both the units (Konopa´seket al., 2002). The Moldanubian domain is traditionallydivided into three tectonic units (Fuchs, 1976). Thestructurally deepest medium-grade MonotonousGroup is overlain by the medium-grade Varied Groupand the structurally highest high-grade Gfo¨ hl Unit.The Monotonous Group consists of biotite-plagioclaseparagneiss with minor orthogneiss, quartzite,amphibolite and locally eclogite bodies (Medaris et al.,1995; OÕBrien & Vra´na, 1995). It contains metagranitoidsof Early Palaeozoic age (see Fig. 2), suggesting aPrecambrian age for the protolith of the surroundingparagneiss (Finger & von Quadt, 1995; Friedl et al.,2004). The Varied Group includes more pelitic protolithsand abundant amphibolite, quartzite, marble andÓ 2010 Blackwell Publishing Ltd207

56 J. FRANĚK ET AL.Fig. 2. Summary of geochronology for the individual units in the Moldanubian, Tepla´-Barrandian and Saxothuringian domains.Histograms of larger datasets are mixed with shaded columns that contain individual age data combined with the stratigraphy ofassociated Palaeozoic sedimentary rocks. Histograms represent the sum of Gaussian probability curves for each radiometric ageincluded. Citations for the 263 plotted ages are given in the Supporting information, Appendix S1. Significant ages are plottedindividually with corresponding blocking temperature. Apatite fission track dating in the Tepla´-Barrandian was performed byGlasmacher et al. (2002), the monazite U–Pb dating from the Moldanubian pegmatite is from Novák et al. (1998). Sx-TB, Saxothuringian-Tepla´-Barrandian;MLC, Maria´nské La´zně Complex; MK, Mu¨ nchberg Klippe; ZEV, Zone Erbendorf-Vohenstrauss; ECC,Eger Crystalline Complex; CBPC, Central Bohemian Plutonic Complex; BPSZ, Bavarian Moldanubicum with Pfahl Shear Zone.calc-silicate intercalations. The protolith of the VariedGroup metasedimentary rocks is supposed to be atleast partly Devonian (Friedl et al., 1993) and Cambrian,based on the 509 ± 27 Ma Nd model age forthe amphibolite layers (Janousˇek et al., 1997). TheGfo¨ hl Unit consists of kyanite-bearing felsic granulite,peridotite, eclogite and migmatitic Gfo¨ hl orthogneissof Cambrian to Early Ordovician protolith age (Friedlet al., 2004; Schulmann et al., 2005).Relevant P–T estimatesThe petrological studies of granulites in the BohemianMassif (Fig. 3) have focused traditionally onpeak metamorphic conditions and amphibolite faciesretrogression involving almost isothermal decompressionfollowed by near isobaric cooling (e.g.Vra´na et al., 1995; OÕBrien & Ro¨ tzler, 2003). Conventionalthermobarometry yields 1000 °C ⁄ 1.6 GPafor the metamorphic peak, followed by retrogressionat 800–900 °C ⁄ 0.8–1.2 GPa (Vra´na, 1989; OÕBrien &Seifert, 1992; Carswell & O’Brien, 1993; Cooke et al.,2000) and 700–800 °C ⁄ 0.5–0.8 GPa (OÕBrien &Carswell, 1993; Vra´na, 1997). However, Sˇtı´pska´ &Powell (2005), Racek et al. (2006) and Tajcˇmanova´et al. (2006) proposed significantly lower temperatureconditions of 750–850 °C at 1.6–1.8 GPa for thepeak metamorphic event followed by retrogression at700–800 °C and 0.5–0.7 GPa. The mantle-derivedrocks enclosed in granulites yield a broad range ofP–T conditions of 800–1350 °C at 2.0–6.0 GPa (seesummary by Medaris et al., 2006). The P–T estimatesfor the Moldanubian mid-crustal rocks yield peakconditions of 750 °C ⁄ 0.7–1.2 GPa for the VariedGroup (Petrakakis, 1997; Racek et al., 2006) and780 °C ⁄ 0.75 GPa for the Monotonous Group(Scheuvens, 2002), with retrogression to 650–720 °C ⁄0.4–0.5 GPa recorded by both the mid-crustal groups(Vra´na et al., 1995; Pitra et al., 1999; Racek et al.,2006).Ó 2010 Blackwell Publishing Ltd208

EXTRUSIONOFLOWERCRUSTINVARISCANOROGEN 57Fig. 3. Relevant P–T data from Bohemiangranulites and surrounding metasedimentaryrocks; from the Moldanubian domain,if not specified from some other unitbelow. The P–T data and related radiometricages (Ma) are assembled from thepublications 1–11, according to the labelsat each polygon. Citations for P–T data –1, Kalt et al. (1999); 2, Kotkova´ (1993) –Eger Crystalline Complex; 3, Kro¨ ner et al.(2000); 4, Linner (1996); 5, Pitra et al.(1999); 6, Racek et al. (2006); 7, Sˇtı´pska´et al. (2004); 8, Sˇtı´pská & Powell (2005);9, Tajcˇmanova´ et al. (2006); 10, Verneret al. (2008); 11, Ro¨ tzler & Romer (2001) –Saxonian Granulites. Citations forgeochronology: 1, Kalt et al. (1997); 2,Kotkova´ et al. (1996) – Eger CrystallineComplex; 3, Kro¨ ner et al. (2000); 5,Gebauer et al. (1989); 7, Sˇtı´pska´ et al.(2004); 9, Schulmann et al. (2005); 10,Verner et al. (2008); 11, Romer & Ro¨ tzler(2001) – Saxonian Granulites.Moldanubian geochronologyMost of the existing U–Pb zircon ages from theBohemian Massif granulites cluster at c. 340 Ma(Fig. 2; for review, see e.g. Janousˇek & Holub, 2007),interpreted as a time of HP metamorphism and subsequentretrogression. The older ages, widely scatteredbetween 450 and 350 Ma, are interpreted usually asprotolith ages (e.g. Wendt et al., 1994; Kro¨ ner et al.,2000; Friedl et al., 2003). Sm–Nd garnet data revealslightly older ages than 340 Ma (e.g. Prince et al.,2000), but their significance is problematic (Romer &Ro¨ tzler, 2001). Janousˇek et al. (2004) suggested anOrdovician granitic protolith with a model age ofc. 450 Ma for the Variscan felsic granulites, supportedalso by radiometric U–Pb zircon ages of Kro¨ ner et al.(2000) and Friedl et al. (2004). The Rb–Sr biotite andmuscovite cooling ages span between 330 and 310 Ma(Van Breemen et al., 1982; Svojtka et al., 2002) similarto 40 Ar– 39 Ar ages on amphibole, muscovite and biotite(Kosˇler et al., 1999). The exposed Moldanubian rocksreached the surface during the Permian, when theywere locally covered by undeformed sedimentaryrocks. Moldanubian Varied and Monotonous Groupsrecord metamorphic events from 355 ± 2 Ma (U–Pbon sphene, Wendt et al., 1993) or the 367 ± 19 Ma forthe Kaplice paragneiss (U–Pb on zircon, Kro¨ ner et al.,1988). A minimum age for the ductile deformationis restricted by syndeformational pegmatites that yieldan age of 331 ± 5 Ma (Rb–Sr on muscovite, VanBreemen et al., 1982).There is a range of syn- to post-tectonic granitoidsintruded into the Moldanubian rocks. The oldest calcalkalineintrusions of the Central Bohemian PlutonicComplex were syntectonically emplaced into uppercrustallevels during regional transpression from c. 354Ó 2010 Blackwell Publishing Ltd209

58 J. FRANĚK ET AL.to c. 346 Ma (Janousˇek et al., 2004) and extension at c.340 Ma (Ža´k et al., 2005a), followed by the intrusionof undeformed granitoids along the boundary betweenthis complex and the Moldanubian domain at c.337 Ma (e.g. Janousˇek & Gerdes, 2003). This corroboratesages of 338–314 Ma for intrusion into theMonotonous Group of mostly post-tectonic granitoidsof the Moldanubian Pluton (e.g. Finger et al., 1997).In northern Bavaria, the Moldanubian rocks exhibityounger metamorphism associated with widespreadmigmatitization dated by U–Pb on monazite at c.326 Ma (Kalt et al., 1997) and c. 315 Ma (Schulzschmalschlager,1984). Syntectonic Bavarian granitoidsyield ages comparable to the MoldanubianPluton (335–322 Ma; e.g. Siebel et al., 2006).GEOLOGY AND STRUCTURAL GEOLOGY OFSOUTH BOHEMIAN MOLDANUBIAN DOMAINThis study focuses on a 130 · 30 km traverse throughthe whole South Bohemian Moldanubian domain fromthe Moldanubian Pluton in the SE to the CentralBohemian Plutonic Complex in the NW following the9HR seismic line (Fig. 4a). The SE part is characterizedby the occurrence of three large neighbouringbodies of felsic granulites (Blansky´ les, Krˇisˇtˇanov andPrachatice granulite massifs, see Fig. 5a). Discontinuousstripes of garnet or spinel peridotites and garnetpyroxenites are systematically found along the marginsof all the granulite massifs, and the Blansky´ les granulitemassif also contains numerous large bodies ofmantle-derived rocks in its interior.The adjacent Moldanubian domain consists of VariedGroup rocks to the SE and Monotonous Group rocks tothe NW. The individual massifs are separated by narrowzones of medium to high-grade paragneisses withaffinity to both Monotonous and Varied Groups. Thenorth-eastern margin of the Blansky´ les granulite massifis marked by a complex association of medium andhigh-grade rocks (Fig. 4a). Farther to the east all theunits become parallel to a major thrust zone defined bythe occurrence of eclogite bodies marking the ductilethrust of the Varied Group over the underlyingMonotonous Group (Rajlich et al., 1986; Vra´na &Sˇra´mek, 1999; Faryad et al., 2006). A zone of lowergrade muscovite-biotite paragneiss intruded by granitoidsof the Moldanubian Pluton occurs in the easternmostextremity of the studied area. The MonotonousGroup in the north-western part of the traverse isintercalated with several large NE–SW elongated bodiesof Varied Group and Gfo¨ hl orthogneiss (Fig. 4a).Earliest granulite facies fabrics S1 and S2The oldest identified fabrics (S1 and S2) are exceptionallywell preserved in an 8.4 · 2.5 km wide ellipticaldomain in the southern part of the Blansky´ lesgranulite massif (Figs 5a,b & 6a). Small relicts arefurther present throughout this massif and rarely alsoin the Prachatice granulite massif. In particular, in theBlansky´ les relict elliptical area the granulites exhibit apenetrative mylonitic foliation S2 (Fig. 5b) whichcontains relicts of S1 compositional layering, bothbearing a stable granulite facies mineral assemblageof Qtz + Kfs + Pl + Grt + Ky + Bt (Sˇtípska´ &Powell, 2005; Franěk et al., 2006). Mineral abbreviationsare after Kretz (1983).The onset of S2 fabric development is characterizedby pervasive recrystallization of the former coarsegrained S1 layered orthogneiss into a fine-grainedmosaic of plagioclase, K-feldspar and quartz, containinga small volume of former partial melt, nowindicated, for example, by cuspate shapes of theseminerals, distributed along feldspar boundaries(Franeˇk et al., 2010). The subsequent folding of S1 isaccompanied by progressive development of S2 axialplanecleavage and penetrative reworking of the S1fabric into the fine-grained, K-feldspar dominatedmatrix containing large quartz ribbons and biotiteflakes oriented into strong subhorizontal stretchinglineation (fig. 3e in Franěk et al., 2006). Plagioclasecoronas developed around kyanite and garnet in thematrix were interpreted as a product of reaction duringdecompression by Tajcˇmanova´ et al. (2007). The S2foliation from the Blansky´ les granulite massif strikesuniformly N–S showing variable dip of 50–90° to theW or less commonly to the E (Figs 4e, 5b & 7).Amphibolite facies D3 deformationIn all the Moldanubian rocks except the granulites, thefirst well-defined metamorphic fabric is a steepNE–SW trending amphibolite facies foliation(Fig. 6b,d). It is assigned as S3 due to a commonstructural concordance with a steep retrogressive S3fabric developed in the granulites.Because of strong later reworking, the S3 fabrics areusually well preserved only in competent lithologies.The only exceptions represent the southern part of theKaplice Zone muscovite paragneiss and the boundaryof the Moldanubian and the Tepla´-Barrandiandomains where steep NW-dipping S3 is present. Rarerelicts of S3 in metasedimentary rocks throughout thestudy area exhibit the same attitude. The S3 foliationoccasionally bears a weak mineral lineation or corrugations,both shallowly plunging NE and SW. In thevicinity of the regionally folded granulite massifs, thegenerally stable orientation of the S3 foliation ismodified by large wavelength late F3 folds.The S3 is defined by preferred orientation of micaand elongated quartz–feldspar aggregates in the Bt +Qtz + Pl + Kfs ± Sil ± Grt ± Ms paragneiss ofthe Monotonous and Varied Groups and by preferredorientation of leucosomes in migmatitic paragneissand the anatectic Gfo¨ hl gneiss. In the medium-gradeKaplice Zone paragneiss the S3 is defined by alignmentof muscovite and biotite and segregation of quartzlenses with cordierite or aluminosilicates (Vra´na et al.,Ó 2010 Blackwell Publishing Ltd210

EXTRUSIONOFLOWERCRUSTINVARISCANOROGEN 59(a)(b)(c)(d)(e)(f)Fig. 4. (a) Structural map of the study region. Structural data from granitoids adopted from Zˇa´k et al. (2005a). (b) Location ofgeological profiles across the granulites. All profiles are of the same scale and not exaggerated. (c) W–E view of Krˇisˇťanov granulitemassif. (d) N–S profile across Prachatice granulite massif, Libı´n Zone and Krˇisˇtˇanov granulite massif. (e) NW–SE Section ofPrachatice and Blansky´ les granulite massifs, roughly following the seismic reflection line 9HR. (f) N–S section of Blansky´ les granulitemassif. Part of profile (e) and profile (f) are modified after Franeˇk et al. (2006).Ó 2010 Blackwell Publishing Ltd211

60 J. FRANĚK ET AL.(a)(b)(c)(d)Fig. 5. (a) Structural map of the SE part of transect, which is dominated by granulites. The S2 fabrics are depicted in the inset (b)in a detail of the relict granulitic domain inside the Blansky´ les granulite massif. (c) Intensity of the L4 lineation in the Prachaticegranulite massif and surroundings interpolated from field data. (d) Steepness of the S4 foliation in the Prachatice granulite massif andsurroundings interpolated from field data.1995; Vra´na & Ba´rtek, 2005). At the boundary of theMoldanubian and Tepla´-Barrandian domains the S3 isdefined by a higher modal muscovite in conjunctionwith a decreasing intensity of partial melting.In the granulites, the S3 represents the dominantplanar fabric in the Blansky´ les and Krˇisˇtˇanov granulitemassifs, whereas in the Prachatice granulite massif,S3 is preserved only along its southern and easternmargins. The S3 results from transposition of S2 viaoutcrop-scale folding accompanied by developmentof S3 axial-plane foliation. The stable assemblage ofQtz + Kfs + Pl + Bt ± Sil ± Grt indicates syndeformationalamphibolite facies retrogression commonlyresulting in weak compositional layering.Syn-D3 partial melting occurred in the outer parts ofthe Blansky´ les and to some extent in the Krˇisˇťanovgranulite massif. At the transition from S2 to S3, theD3 microstructures still resemble the granulitic structure,suggesting that the onset of the D3 deformationoccurred at the granulite–amphibolite facies transition.The timing of D3 in the granulites is best bracketed bythe 337 ± 0.3 Ma crystallization age of a granulitefacies hyperpotassic dyke deformed by D3 (Aftalionet al., 1989; Sla´ma et al., 2008) and the crystallizationage of 340 ± 3 Ma for post-D4 cordierite-bearingleucosomes (Kro¨ ner et al., 2000).In granulites the S3 exhibits arcuate geometry(Figs 5a & 7) with moderate to steep dips (Kodym, 1972;Franeˇk et al., 2006). The arrangement of poles of S3along a great circle in each of the granulite massifssuggests a cylindrical fold geometry. Changes in attitudeof S3 in all the granulite bodies appear as km-scaleflexures, or rarely as several km wide sequences of outcrop-scaleparasitic folds (Fig. 4b–f). Axial planes ofthese large folds in all three granulite bodies revealsimilar orientation (Figs 5a & 7). For the tight fold of thePrachatice granulite massif, we have constructed asubvertical NNW–SSE trending axial surface oriented262 ⁄ 87 (dip direction ⁄ dip in degrees) which is similar inorientation to the 246 ⁄ 89 axial planes of large-scale foldsdefined previously in the Blansky´ les granulite massifby Franěk et al. (2006). For the wide arcuate geometryof S3 in the Krˇisˇťanov granulite massif we are only ableto estimate a steep axial-plane dipping to the W.Ó 2010 Blackwell Publishing Ltd212

EXTRUSIONOFLOWERCRUSTINVARISCANOROGEN 61(a)(b)(c)(d) (e) (f)Fig. 6. Field photographs of structures. (a) Development of S2 axial fabric during F2 passive folding in granulites. (b) Transpositionof S2–S3 via mesoscopic similar folds in granulites. (c) Initial S4 in discrete shear zones affecting S3 in the Prachatice granulite massif.(d) Steep S3 only weakly affected by D4, lower grade Kaplice paragneiss in the SE of the transect. (e) S4 crenulation cleavage in amigmatized paragneiss, Varied Group in the NW of the transect. (f) S3 transposed into S4 via crenulations and shear bands, whichare commonly filled with granitic leucosome. Orthogneiss in NW part of the transect.The biotite or sillimanite L3 mineral lineationplunges north to northwest, parallel to the hinges ofthese large folds in the Blansky´ les granulite massifexcept for the northern part of this body, where theyvary from south to east. The well-developed L3 lineationin the Prachatice granulite massif plunges shallowlysubparallel to the strike of the S3. In contrast,the L3 lineation in the Krˇisˇťanov granulite massif ispoorly developed except in the NW corner, where theplunge is generally to the SW (Fig. 8).In the metasedimentary rocks separating the granulitemassifs, the S3 fabrics are well preserved only inthe south, whereas towards the north and west onlyrare outcrop-scale relicts of S3 occur. Ubiquitous syndeformationalpartial melting of paragneisses resultedin stromatitic migmatitic layering parallel to the S3schistosity. The S3 in these metasedimentary rocksstrikes parallel to the S3 in neighbouring granulites; inthe Lhenice Zone the S3 forms a tight vertical N–Selongated fan-like pattern (Figs 4e & 5a) whereas inthe Libín Zone it dips steeply to the southwestunderneath the Krˇisˇťanov granulite massif (Figs 4d &5a). Rocks of the Lhenice Zone have an intense shallowlyN–S-plunging lineation defined by biotite alignment,whereas in the Libı´n Zone a L3 lineation is notdeveloped.Subhorizontal amphibolite facies fabric S4The most prominent fabric in rocks of the traverseis the S4 foliation that generally strikes NE–SWsubparallel to the Moldanubian–Tepla´-Barrandianboundary, in general dipping gently to the NW. Onlyalong the southern edge of the Central BohemianPlutonic Complex and 10 km northwest of thePrachatice granulite massif does the S4 fabric dip tothe SE, forming large-scale open fold-like flexures inthe Monotonous and Varied Groups; a N–S to NE–SW stretching lineation on the S4 fabric is subparallelto the hinges of these flexures. Similar to the S3 fabric,in the vicinity of the granulite massifs, the S4 fabricexhibits significant perturbation as a result of deformationpartitioning and locally attains a steep attitudesubparallel to the S3 fabric (e.g. at the eastern marginof the Blansky´ les granulite massif).Partial melting associated with the S4 fabric in theMonotonous and Varied metasedimentary rocks(Fig. 6e,f) is of lower volume than that associated withÓ 2010 Blackwell Publishing Ltd213

62 J. FRANĚK ET AL.Fig. 7. Equal area lower hemisphere projections of S2, S3 and S4 fabrics in the granulite-dominated region of the transect contrastedagainst the same fabrics developed in the Moldanubian domain outside the granulite-dominated region.Fig. 8. A krigging interpolation to depictthe prevalence of particular fabrics inferredfrom field study of individual outcrops.Grey levels show the extent of reworkingby particular deformation phases (darkest =S2, lightest = post-D4 late NW-SE crenulationcleavage). Four generations ofcorresponding lineations are depicted.Ó 2010 Blackwell Publishing Ltd214

EXTRUSIONOFLOWERCRUSTINVARISCANOROGEN 63the S3 fabric. In the granulites the S4 is marked bytransposition of the typical dm scale compositionallayering and syntectonic replacement of garnet bybiotite and sillimanite. The transition from S3 to S4proceeds via development of shallowly dipping S4-parallel shear zones, often accompanied by buckling ofS3 at various scales. Occasional cm to dm scale intrafoliationpinch and swell structures with steep E–Wtrending tensional cracks filled by undeformed graniticleucosome develop in the S4 indicating verticalshortening and N–S-oriented stretching operating attemperatures around the granite solidus.The S4 intensity systematically increases from the SEto the NW towards the Central Bohemian PlutonicComplex. Here, the deformation culminates in highlynon-coaxial shearing manifested by a shear zoneseveral kilometres wide intruded by syntectonicgranitoids (Ža´k et al., 2005a) dated at 342–337 Ma,which corroborates a U–Pb age of 341 ± 3 Ma forzircon from leucosome patches cross-cutting the S4fabric of the Prachatice granulite massif (Kro¨ ner et al.,2000). The ambient P–T conditions in this regionduring formation of a flat fabric comparable with theS4 fabric of this study were determined as >720 °C at0.4 GPa by Pitra et al. (1999) and Scheuvens (1999).In the hangingwall paragneisses, at the Moldanubian–Tepla´-Barrandian boundary, the D4 imprint rapidlydies out, documenting a block several kilometres wideof lower grade rocks that are inferred to have sunkbetween the Moldanubian and Tepla´-Barrandiandomains.In the east, the intensity of the D4 deformationgradually decreases so that only the western part of thegranulite bodies shows relatively strong reworking bythe flat S4 fabrics at 765 °C and 0.76 GPa (Fig. 6c;Verner et al., 2008). Whereas most of the Prachaticegranulite massif is strongly reworked by the subhorizontalS4 foliation and weak N–S lineation (Figs 5c,d& 8), the Krˇisˇťanov granulite massif exhibits onlymoderate reworking by a shallow WNW-dipping S4foliation (Verner et al., 2008) and the Blansky´ lesgranulite massif to the east is almost unaffected by D4(Fig. 4b–f). In the Lhenice Zone the S4 forms an opensynform with horizontal N–S axis, and the rocks exhibita strongly developed, shallow N–S-plungingstretching lineation. Within the Libı´n Zone the S4fabric dips moderately to the SW below the Krˇisˇtˇanovgranulite massif; a weak lineation of variable attitude ispresent. North of the Blansky´ les granulite massif theS4 dips shallowly to the SSW. Further southeast the S4is variable and generally has a steeper attitude indicatingthat the D4 intensity decreases. Finally, in thesouth of the medium-grade paragneiss of the KapliceZone the S4 is only weakly developed.NW–SE trending steep cleavage frontThe SW part of the traverse is affected by late ductiledeformation resulting in development of a crenulationcleavage, which develops into a new penetrative foliationdipping moderately to steeply to the NE (Fig. 4a).The intensity of this fabric generally increases southwestwardand it becomes the dominant fabric in theBavarian region characterized by coeval dextral shearzones (e.g. Behrmann & Tanner, 1997; Finger et al.,2007). There is no expression of such structures insidethe granulite massifs implying that the geometry andinternal structure of the granulites were not affected bythis late deformational event.REFLECTION SEISMIC PROFILEMost of the traverse is parallel to a deep seismicreflection profile (9HR) which was shot in a NW–SEdirection through western half of the Bohemian Massifusing explosive sources. The seismic profile, which isonly partly published (Tomek et al., 1997), extendsacross the region between the eastern boundary of theSaxothuringian domain in the NW to the eastern edgeof the Blansky´ les granulite massif in the SE. Thisstudy examines the SE half of the profile (Figs 9 & 10),which offers a good quality record from 2 km belowthe surface to the Moho at 40 km depth, exhibitingin places also reflections in the upper mantle.Characterization of reflection seismic dataFor most of the traverse the seismic profile runsperpendicular to the strike of rock fabrics, whichsuggests that the dips of the reflectors are close totrue dips, at least in the upper crust. The line drawing(Fig. 10) based on the migrated seismic record revealsa striking difference between the seismic properties ofthe Tepla´-Barrandian and the Moldanubian crust.Whereas the Moldanubian crust exhibits severalseismically different domains, the highly reflectiveTepla´-Barrandian upper crust shows one single set ofparallel and very strong reflection packages (B1–B6)dipping to the SE. These packages are interrupted atdepth along a NW-dipping virtual line, below whichthe Tepla´-Barrandian lower crust shows very poorreflectivity, being constrained by subhorizontal, butdiscontinuous Moho reflections (M1) at 11 s TWT(two-way time). The Tepla´-Barrandian–Moldanubianboundary zone, intruded by granitoids of the CentralBohemian Plutonic Complex, shows a sharp decreasein reflectivity of the whole crust in a 10 km widecolumn.The Moldanubian crust shows contrasting seismiccharacteristics in the eastern part, which is dominatedby granulites, and the western part, which is devoid ofthese rocks. The western part is characterized bystronger reflection packages (K1 and K2, V1–V4) andanastomosing reflections down to 5 s TWT, whichresemble pinch and swell structures of 20 km wavelength.The upper crust in the east of the seismic profileis cross-cut by a strong and unusually straight reflectionpackage (G2) dipping moderately to the SE.Ó 2010 Blackwell Publishing Ltd215

64 J. FRANĚK ET AL.Fig. 9. The 9HR seismic profile showing the migrated raw data. Vertical scale approximately equals the horizontal scale. Position of the profile is shown in Fig. 4a, it coincideswith section 3 of the gravity model shown in Fig. 11.Ó 2010 Blackwell Publishing Ltd216

EXTRUSIONOFLOWERCRUSTINVARISCANOROGEN 65(a)(b)Fig. 10. The 9HR seismic profile showing line drawing of the migrated reflections. Position of the profile is shown in Fig. 4a. (a) Upper crust with structural profile that is basedonly on field structural study. Surface extent of major lithological units is plotted above. CBPC refers to Central Bohemian Plutonic Complex, TBU to Tepla´-BarrandianUnit. (b) Seismic reflections with package description, thickness and darkness of lines express intensity of reflections. The reflections in (b) west of the granulites fit very well withthe independent geological interpretation down to 5 km depth. Vertical scale approximately equals the horizontal scale.Ó 2010 Blackwell Publishing Ltd217

66 J. FRANĚK ET AL.At the top, the package terminates abruptly at 1 sTWT below the western edge of the Prachatice granulitemassif, probably on a local N–S fault. Thestrongest reflections in the profile appear below thePrachatice granulite massif; they diminish to the SE at5 s TWT, below the centre of the Blansky´ les granulitemassif. A parallel weaker package (G1) of12 km length appears attached to the top of thispackage of strong reflections. Except for these features,the crust below the granulites exhibits almost noreflectivity and there is only a weakly defined Moho(M4) at 12–13 s TWT. The reflections along the SEedge of the seismic profile are probably just artefacts ofthe migration procedure. This lack of reflections contrastswith a package of strong NW-dipping mantlereflections below 12s TWT (Cˇ. Tomek, personal communication;not shown in Fig. 10). The western part ofthe Moldanubian middle to lower crust exhibits uniformlydistributed horizontal reflections designated asLC, which appear in a lens-shaped domain60 · 15 km in size. There are several packages ofstrong reflections (Z1–Z4) in the underlying lowercrust that show apparent dips of 30–40° to the SE andcontinue through discontinuous MOHO reflections(M2 and M3) into the uppermost mantle.Interpretation of reflection seismic dataIn the high-grade Moldanubian domain, which recordspolyphase deformation, the reflection packages mayrepresent lithological layering, penetrative foliation,faults or boundaries of intrusive bodies. We try toovercome this ambiguity by careful comparison withthe surface extent of major lithological units and withour structural interpretations.The upper-crustal reflection packages (B1–B6) in theTepla´-Barrandian Unit correspond to Proterozoic lowgradesequence of siltstones interlayered with basaltsexposed on the surface that show SE-dipping schistosity.Termination of these packages at the margin ofthe Tepla´-Barrandian Unit may document sidewardintrusions of granitoid bodies related to the CentralBohemian Plutonic Complex, or reworking by laterVariscan deformation. The low reflectivity in thecrustal column directly below the Tepla´-Barrandian–Moldanubian boundary can be best explained by theoccurrence of Central Bohemian Plutonic Complexgranitoid intrusions at depth below the MonotonousGroup exposed at the surface.The anastomosing reflections in the upper crust ofthe western Moldanubian domain correlate well withthe geometry of the S4 fabrics as determined by ourfield research (Fig. 10a). This conformity suggests thatall the anastomosing reflections can be related to thesubsurface continuation of the S4 fabrics. The V1–V4reflection packages correlate well with surface exposuresof km-scale lenses of Varied Group rockssuggesting that the reflectors in this case representinterlayering of paragneisses with marbles, amphibolitesand other lithologies, which are oriented parallelto the S4 fabric. The K1 and K2 reflection packagesdocument probable subsurface occurrence of variedlithological intercalations at the western side of theMoldanubian domain. The nature of the straightreflection package (G2) dipping below the granulitemassifs remains ambiguous, because the observed highamplitude and length of reflections may be causedeither by lithological layering similar to that of theVaried Group or by a thick zone of intensive mylonitization.Based on the single seismic section, theorientation of the G2 reflectors can vary between asteeper dip to the NE through medium dips to the SEto steeper dips to the SW. Such orientations cannot bedirectly correlated to any structure mapped on thesurface. The straight geometry of the G2 packagesuggests that the corresponding reflectors were notsignificantly affected by the D4 deformation, or thatthey may represent an anomalously oriented D4structure. The low reflectivity in the crustal columnbelow the granulite massifs may be caused either bylack of reflection inducing inhomogeneities or extremesteepness of reflectors, which would not be detectableby seismic survey. The low energy of the shots cannotbe responsible for the lack of crustal reflections,because the much deeper mantle reflections have beenrecorded below this seismically transparent zone. Therest of the Moldanubian crystalline crust exhibits significantreflectivity, suggesting that the Moldanubiancrust in general contains abundant reflectors thatwould be recorded if they had a suitable attitude.Below the granulites the reflectors may have beenpartially destroyed by the rise of these HP rocks, but itis argued that this crustal section likely still containsnumerous reflectors similar to the rest of the Moldanubiandomain, but the attitude of these reflectors istoo steep to be detected by the seismic method. Incombination with the dominance of steep fabrics in theexposed granulites, the lack of reflections serves asindirect evidence for the dominance of steep foliationsin the seismically transparent zone throughout thecrust below the granulite massifs. The homogenouslydistributed reflections in the middle to lower crust,designated as LC in Fig. 10b, do not reach the surfacealong the 9HR line. Such middle to lower crustalreflections are classically, but without direct evidence,ascribed to subhorizontal intrusions. These horizontalreflections cannot represent the deeper continuation ofthe S4 fabric, because of the sharp transition zone withthe anastomosing upper-crustal S4-related reflectionsand the general dip of the S4 to the NW. The LCreflections are unlikely to represent a fabric older thanS4, because the D3 phase of horizontal shorteningleads to development of steep NE–SW trending mostlymigmatitic fabrics along the traverse. The abrupttransition in seismic fabric may also mark a rheologicaltransition between the Variscan upper and middle tolower crust, with the age of the LC reflectionsremaining unknown.Ó 2010 Blackwell Publishing Ltd218

EXTRUSIONOFLOWERCRUSTINVARISCANOROGEN 67(a)(c)(b)(d)Fig. 11. (a, b) Three-dimensional gravity model of the Prachatice granulite massif, comparison of measured with modelled gravityfield. (c) Two most significant model sections through the granulite massif. Curves above depict the values of the measured Bougueranomalies (thick dashed) and those modelled in three-dimensional (solid). The thin dashed line shows the two-dimensional gravityeffect of the individual sections. Only 20 km of the horizontal extent of each section relevant to the Prachatice granulite massif areshown. Section 3, which coincides with the reflection seismic profile, contains line drawing of corresponding reflections. White values insection 4 refer to final densities used in the gravimetric model. (d) Map of Bouguer gravity anomalies (modified after J. Sˇvancara,unpublished data) with outlines of geological units, location of gravity model profiles and the 9HR seismic line.The lower crustal reflections Z1–Z4, which areconcentrated in a triangular region in the NW part ofthe Moldanubian domain, dip moderately to the SEand their lower tips appear to cut through the Mohointo the uppermost mantle. Such characteristics arecomparable with seismic images of subduction sutures,Ó 2010 Blackwell Publishing Ltd219

68 J. FRANĚK ET AL.Fig. 12. Interpretative geochronological sketch of Variscan evolution in the Bohemian Massif, as it is related to the traverse.but the speculation that the observed seismic structuremay represent the remnant of a suture zone betweenan unknown plate and the overlying Moldanubiandomain is very poorly constrained. The discontinuousand weakly defined Moho reflections (M1–M4) belowthe Moldanubian crust document systematic deepeningof the Moho from 35 km in the NW to 42 kmin the SE. This deepening is in agreement with theMoho geometry defined by the refraction study ofHrubcova´ et al. (2005) and by the passive seismologyexperiments of Plomerova´ et al. (2005).GRAVITY MODELLING AND VERTICAL EXTENTOF THE FELSIC GRANULITESTo decipher the vertical extent of the granulite massifs,an interpolated grid of Bouguer anomalies coveringmost of the traverse (provided by J. Sˇvancara,unpublished data; Fig. 11d) has been examined. Thearea is dominated by large gravity lows related togranitoids of the Moldanubian Pluton in the SE or tounknown sources in the centre and NW parts. Positiveanomalies are all smaller scale and appear in thesouthern part of the area.Forward gravity modelling in three-dimensions hasbeen undertaken using the IGMAS software (Schmidt& Go¨ tze, 1999), where geological bodies are representedby polyhedrons triangulated between manuallydrawn sections. The only suitable body for such anapproach is the Prachatice granulite massif, the surfaceextent of which correlates with a pronounced circularnegative anomaly. Its lithological homogeneity and thesmoothness of the gravity anomaly suggest that nosignificantly denser bodies (e.g. ultrabasites) areinvolved in this massif. The Krˇisˇťanov granulite massifinduces only a weak negative anomaly possiblyaffected by neighbouring intrusion of denser K–Mgsyenites. The Blansky´ les granulite massif partiallycorrelates with a positive anomaly, despite the densityof the felsic granulites being lower than the surroundingmetasedimentary rocks (Appendix S2),probably as a result of abundance of ultrabasic bodiesÓ 2010 Blackwell Publishing Ltd220

EXTRUSIONOFLOWERCRUSTINVARISCANOROGEN 69(a) (b)(c)(d)Fig. 13. Mid-crustal evolution of the region focused on detailed behaviour during the D3 and D4 phases. (a) Amplification of the complex granulite bulge under dextraltranspression. The driving force is transferred probably from the Saxothuringian subduction and collision. KG, PG and BLG refer to the ancestors of the Krˇisˇťanov, Prachaticeand Blansky´ les granulite massifs, TBU marks the Tepla´-Barrandian Unit and CBPC the Central Bohemian Plutonic Complex. (b) Hardening of felsic granulites caused by coolingresults in amplification of several F3 regional-scale folds due to the dextral shear. (c) Final stage of fold development accompanied by intrusion of Mg-K rich syenites. (d)Immediate change of regional kinematic regime to vertical flattening. Pervasive development of flat S4 fabrics in the NW part of the traverse is combined with normal movement ofthe Tepla´-Barrandian block in the NW partially along the flat S4 foliation.Ó 2010 Blackwell Publishing Ltd221

70 J. FRANĚK ET AL.and underlying dense Varied Group rocks. Theanomalies involved in our gravity model have smallareal extent and exhibit high gradients at their edges,implying that all of them are caused by upper-crustalheterogeneities. Removal of regional long-wavelengthgravity components caused by deeper heterogeneities isthen unnecessary and the original values of theBouguer anomalies have been used in the modellingthat follows. For the same reason, the model concernsonly upper-crustal bodies down to 10 km depth.The geometry of the Prachatice granulite massif hasbeen approximated in six vertical sections 40 km longand 10 km deep, trending NW–SE parallel to theseismic reflection line 9HR (Fig. 11a–c). Unlike thetwo-dimensional gravity model of Vra´na & Sˇra´mek(1999), the three-dimensional approach was chosenbecause the relationship of this granulite massif with itssurroundings suggests a complex non-cylindricalgeometry of structures on the scale of our model and acircular shape of the anomaly also means that a twodimensionalapproach is unsuitable. The model isconstrained by the surface extent of the geologicalunits, structural data extrapolated to 1–2 km depthand the reflection seismic line that cross-cuts thePrachatice granulite massif and the centre of the negativeanomaly. The database of rock densities(Appendix S2) from Hana´k et al. (unpublished data),Chlupa´cˇova´ et al. (unpublished data) and Blizˇkovsky´et al. (unpublished data) provided starting guesses,which were carefully adjusted when the constraintsdiscussed earlier in the article restricted changes ofgeometry for the modelled bodies. This is a commontechnique because rock densities vary in space, e.g. theVaried Group in surface section exhibits significantlithological change on a kilometre scale. The model wasdesigned to be as simple as possible to avoid ambiguoussolutions, but its simplicity naturally causes divergenceof the modelled gravity values from the measured valuesat the model edges. In our model, the Prachaticegranulite massif reaches a maximum depth of 5.7 km,having a thick lensoidal shape. It cannot extend belowthe strong reflection package, because such a packageof reflectors could not be expected in a lithologicallyhomogenous granulite massif. It also cannot form athinner lens, because then an unreasonably low densityfor the felsic granulite would be needed to produce themeasured gravity low. A spectacular detail is the cuspof high-density Varied Group rocks introduced into thegranulite massif from below in the hinge region of thelarge fold (e.g. at 15 km in section 3 in Fig. 11c).Inferring this structure is the only reasonable way toproduce the observed cranked gravity curves. It isprobably structurally analogous to a cusp at the SWedge of the Blansky´ les granulite massif, where a sheetof amphibolites with ultrabasites penetrates deeply insidethe granulite massif along the large fold axialplane. Our results differ from the two-dimensionalgravity model of the Prachatice granulite massif publishedby Vra´na & Sˇra´mek (1999) who expected arectangular rather than lensoidal shape of the massif,reaching to 9 km depth. The discrepancy is causedpartially by the different rock densities used andpartially by the two-dimensional approach. It is arguedthat our model approximates better the reality as aresult of calculation in three-dimensions as wellas using better structural constraints from the nearsurfacegeology.DISCUSSION – KINEMATIC AND MECHANICALSIGNIFICANCE OF THE DEFORMATION FABRICSTectonic significance of granulite facies fabricsThe granulite facies S1–S2 foliations represent aunique example of lower crustal fabrics in respect ofthe whole Bohemian Massif. The decompressional P–Tpath indicates that the S2 fabric originated duringexhumation. The highly discordant relationshipbetween the S2 granulite facies fabric and the S3amphibolite facies foliation is the most importantstructural observation from the South Bohemiangranulites (Franeˇk et al., 2006). The HP granulitesrecord the early deformation, which are not recordedin rocks of either the middle- or the upper-crustalunits. Therefore, there is a possibility that the granulitefacies D2 mylonitization is older and occurred in acrustal unit that was geographically remote comparedto the Lower Carboniferous D3–D4 history recordedin lower-, middle- and upper-crustal units (Fig. 12).Horizontal shortening and development of the regionalvertical fabricThe distribution of the steep S3 foliation in thegranulite massifs shows important variations, whichcan be interpreted as a result of large-scale folding ofgranulite massifs with steep N–S axial planes (Franěket al., 2006). The granulite bodies were morecompetent than the surrounding metasedimentaryrocks due to cooling after exhumation at later stagesof D3, which enabled fold amplification. The differencebetween the asymmetrically folded granulite S3fabrics and the straight NE–SW S3 fabric homogeneouslydeveloped in middle-crustal host rocks indicatesa non-coaxial dextral shear operating duringD3 in both the middle- and lower crustal units. Thesteep attitude of N–S axial planes rules out thepossibility that the S3 folding proceeded during thedevelopment of the flat S4 foliation. The axial planeslie close to the plane of maximum flattening of theinstantaneous D3 strain ellipsoid, while the generalNE–SW trend of the S3 fabric reflects the finitestrain orientation.When the Z-shaped Blansky´ les, the symmetricalKrˇisˇťanov and the V-shaped Prachatice granulitemassifs are unfolded to yield originally NE–SW strikingsheets, then the S3 generally dips to the west in theBlansky´ les, subvertically in the Krˇisˇtˇanov and to theÓ 2010 Blackwell Publishing Ltd222

EXTRUSIONOFLOWERCRUSTINVARISCANOROGEN 71(a)(b)(c)(d)Fig. 14. Geotectonic model of granulite lower crustal exhumation with subsequent stacking and subhorizontal mid-crustal flowdepicted along a two-dimensional NW–SE lithospheric-scale profile. (a) Hypothetical formation of granulites, which are suggested tohave formed during SE-directed subduction leading to intermixing of distinct crustal levels, probably of Late Devonian age (based onnumerical models of, e.g. Gerya & Stockhert, 2006). (b) Early Carboniferous subduction stage with intensive arc magmatism andshortening in the arc and back-arc domain. (c) Prolonged shortening leads to extrusion of the lower crustal rocks from below the arcregion, while the fore-arc block slides to the NW from the amplifying extrusion. (d) Detail of the (c) stage focused on the traversestudied for this paper. Abbreviations used: UC, upper crust (above root); MC, middle crust; LC, lower crust; CC, continental crust(Armorica); EC, eclogitized crust (Armorica); UM, upper mantle; Ast, asthenosphere; PG, Prachatice granulite massif; BLG, Blansky´les granulite massif.east in the Prachatice granulite massif. Such spatialvariations in dip may be interpreted to reflect a20 km wide positive subvertical fan-like structurestriking NE–SW (Figs 13 & 14d). The distribution ofserpentinite stripes inside the Blansky´ les granulitemassif accompanied sometimes by Varied Groupmetasedimentary rocks was interpreted by Franeˇket al. (2006) as a result of isoclinal folding of a granulitesheet and the host rocks accompanied by imbricationduring early D3. This observation indicatesoverall isoclinal folding of lower- and mid-crustal unitsthat dismembered a coherent mass of exhuming granuliteinto several massifs.The original NE–SW trending positive fan-likestructure along the boundary between the Varied andMonotonous Groups can be interpreted as the core ofcrustal-scale vertical extrusion of orogenic lower crustover medium-grade rocks to the NW and SE (Figs 13& 14). The fan-like geometry may result either fromdevelopment of a ductile positive flower structureduring extrusion or due to subsequent flattening andductile thinning, that is responsible for the followingD4 subhorizontal flow.The D3 vertical fabric in the Moldanubian domain isgeometrically and kinematically consistent with thefabrics that are developed in the arc-related CentralÓ 2010 Blackwell Publishing Ltd223

72 J. FRANĚK ET AL.Bohemian Plutonic Complex as well as in the adjacentsupracrustal Tepla´-Barrandian domain. The steepNE–SW trending fabrics preserved in slates and volcanicrocks of the Neoproterozoic unit as well as insyntectonic calc-alkaline (354–346 Ma) intrusions areconsistent with dextral transpression and coeval withsyntectonic contact metamorphism at 349–341 Ma(Scheuvens & Zulauf, 2000; Zˇa´k et al., 2005a,b). Thespatial distribution of fabrics and metamorphic rocksdocumenting D3 deformation reflects a significantexhumation event of HP rocks up to 0.7 GPa. Deeperparts of the subvertical granulite fan-like structureare expressed in the seismic section as a low reflectivityregion below the granulite massifs, which indicates theprobable existence of steep extrusion-related S3 fabricsdown to the Moho and marks the exhumation path ofthe granulite-dominated belt. Thus, the present crustalstructure preserves the geometry of the subverticalextrusion channel related to the flow of HP rocks fromthe base of the crust. The existence of other granulitebodies deeper in this channel is indirectly supportedby gravity modelling suggesting a significant thicknessfor individual granulite bodies (Fig. 11c). In conclusion,the D3 deformation reflects mechanical couplingamong the upper, middle and lower levels of theorogenic crust approximately at 354–342 Ma.Mid-crustal horizontal flow and crustal-scale detachmentThe good correlation of the S4 fabric with the subsurfaceseismic reflections, especially in the NW part ofthe traverse (Fig. 10a), implies that S4 extends at leastto 10 km depth. The surface fabric trajectories combinedwith the anastomosing pattern of seismicreflectors indicate that the large-scale geometry can beinterpreted as a set of 10–20 km wide pinch and swellstructures with subhorizontal NE–SW axes and shallowlyNW-dipping median planes. A characteristicfeature is the systematic decrease of D4 intensity fromthe northwest to the southeast so that the central partsof the granulite massifs are unaffected by this deformationevent, which is in accord with very limitedflattening of the Prachatice granulite massif inferredfrom the gravity modelling. This tectonometamorphicpattern can be interpreted as a result of a large-scaledetachment zone weakened by partial melting, alongwhich the supra-crustal Tepla´-Barrandian Unit sliddown to the W–NW. Scheuvens & Zulauf (2000) andDo¨ rr & Zulauf (2008) suggested that sliding along theCentral Bohemian Shear Zone, supported by limiteddiapiric ascent of migmatites, occurred at c. 340 Maunder dextral transtension. We suggest that a notableamount of vertical movement was achieved along S4planes in the Moldanubian migmatites, based on theextreme intensity of the S4 development and correlationwith the results of Zˇa´k et al. (2005a). The verticaldifferential motion at the Moldanubian–Tepla´-Barrandianboundary is estimated to 10–15 km, basedon data of Scheuvens & Zulauf (2000). The flatgeometry indicates also a component of lateral ductiletransport, which is consistent with numerical models oflate evolution in the core of hot orogens (e.g. Beaumontet al., 2006; Jamieson et al., 2007), but difficultto quantify in the Moldanubian case.D4 fabric maps (Figs 5c,d & 8) show that rheologicallyweaker host gneisses exhibit a structural patternconsistent with viscous flow around granulite bodies,which behaved as rigid ellipsoidal objects (e.g. in thecentral part of the Prachatice granulite massif, the S4exhibits a flat attitude interpreted as the apical part ofan ellipsoidal crustal-scale granulite boudin). At themargins of this massif, the steeper and in the east alsosynformal S4 fabrics represent 2 km wide neck zones.The Varied and Monotonous Group rocks surroundingthe granulites are devoid of (U)HP relicts. Thus,the HP granulite boudins do not represent relicts of anoriginally coherent HP metamorphic unit where theVaried and Monotonous Group rocks would be completelyretrogressed, unlike other orogenic root systems(e.g. Engvik & Andersen, 2000). In conclusion, the S4fabric corresponds to zone of coaxial vertical shortening,with the exception of a narrow highly noncoaxialzone at the Moldanubian–Central BohemianPlutonic Complex boundary in the NW. Here, the D4deformation is partially responsible for decouplingof the Neoproterozoic upper-crustal lid from theMoldanubian middle crust. Such deformation partitioningis of common occurrence across large-scalenormal shear zones related to exhumation (e.g. Lawet al., 1994; Little et al., 1994). The subhorizontal flowis responsible for vertical exhumation from 0.7–0.8GPa to 0.3–0.4 GPa at 342–337 Ma.TECTONIC MODELDevonian subduction of continental crustIn addition to the Carboniferous events, the followingevidence supports an earlier independent Variscantectonic regime throughout the N–NW BohemianMassif in the Devonian. Relicts of a mid-Devonianoceanic subduction zone, e.g. the Mu¨ nchberg Massif,occur in the hangingwall of the main continentalSaxothuringian subduction zone (e.g. OÕBrien, 1997;Konopa´sek & Schulmann, 2005). Also, there are felsicgranulites of Devonian age with a granulite facieslayering reported from several places in the BohemianMassif (e.g. Sowie Go´ry Mountains, OÕBrien et al.,1997). Finally, there is a range of Devonian zirconU–Pb and Pb–Pb ages from granulites (Wendt et al.,1994; Schulmann et al., 2005).Based on this evidence, it is suggested that thegranulite facies rocks in the Moldanubian domainoriginated during a subduction event prior to theSaxothuringian continental subduction, which culminatedin the Carboniferous collisional event at c.340 Ma. The S1–S2 fabrics, which are discordant tothe overall NE–SW Moldanubian trend, could haveÓ 2010 Blackwell Publishing Ltd224

EXTRUSIONOFLOWERCRUSTINVARISCANOROGEN 73developed during such an older episode, possibly inrelation to the early arc history. This model is stronglysupported by the numerical simulations of Gerya &Stockhert (2006), who proposed an origin for HPgranulites in a lithosphere-scale subduction wedge. Inthis model, the rocks of the footwall plate are draggedto depths corresponding to 2.0–2.5 GPa, exhumedbackwards by a return flow and accreted to the base ofthe hangingwall crust (Fig. 14a). Other possible modelsinclude those of Gerya et al. (2008), Warren et al.(2008) and Beaumont et al. (2009) for buoyancy-drivenexhumation of weakened continental rocks in a subductionchannel during the early stages of continentalcollision. Janousˇek & Holub (2007) suggested a similarearly evolution and pointed out the geochemicalaffinity of the felsic granulites with numerous pre-Variscan granitoids of the Saxothuringian domain, i.e.in the footwall of the Saxothuringian subduction. Weinfer that during the D2 deformation the granuliteswere partially exhumed in such a tectonic setting fromthe peak pressure of 1.8–2.0 GPa and accreted to thebase of the hangingwall lower crust (tentativelydepicted in Fig. 14a), because the following D3 deformationcommenced at lower granulite facies conditions(Fig. 3). Buoyancy of the weak felsic granulites presumablyplayed an important role during their transportand exhumation. Alternatively, the low viscosityand low density of the granulites could have allowedtheir rise vertically, through the mantle wedge above thesubduction zone, as suggested for example by Oncken(1998) for the Saxonian granulites. During the ascent,the granulites would also trap the characteristic bodiesof garnet and spinel peridotites. The very limited recordof the early S1 and S2 fabrics precludes distinctionbetween exhumation along the suture zone or thissecond case, as well as a more precise description of theearly granulite transport path.Carboniferous extrusion of granulites and collapse ofcrustal lidThe parallelism of the S3 and S4 strike with the twomain sutures in the Bohemian Massif – the Saxothuringiansubduction zone in the NW and the Moldanubian–Bruniaboundary in the SE (Fig. 1b), indicatesthat at least one of these collisional zones generatedhorizontal shortening inducing the D3 and D4 events.Both of them could have acted as a stiff indentorunderthrusting below the already juxtaposed Tepla´-Barrandian and Moldanubian domains. Geochronologicalarguments (e.g. Schma¨ dicke et al., 1995)indicate activity of the Saxothuringian subduction andsubsequent collision between 385 and 335 Ma (Figs 2& 12; e.g. Konopa´sek & Schulmann, 2005), a timerange involving the radiometric ages of Moldanubiandeformation. The easterly Brunia margin documentsprolonged underthrusting from 330 to 310 Ma (e.g.Hartley & Otava, 2001), too young to cause theMoldanubian shortening.The spatial abundance of S3 steep fabrics (Fig. 5)and their vertical extent (Fig. 10b) suggest homogeneousdevelopment of the steep foliation throughoutthe arc (the Central Bohemian Plutonic Complex) andback-arc domains (Moldanubian domain in the senseof Schulmann et al., 2005, 2009) in a NW–SE dextraltranspressional regime. In the Tepla´-Barrandian Unit,S3 is developed only at the eastern edge (e.g. Zˇa´k et al.,2005a) suggesting that the remote part of western forearcupper crust behaved as a stiff block during the D3.Strain localization below the magmatic arc wasresponsible for the unusually massive amplification ofF3 folds in the lower crust that resulted in localizedexhumation of the deep-seated granulites to midcrustallevels in the form of a 20 km wide ductileisoclinally folded subvertical fan-like structure (Figs 13& 14b–d). The adjacent mid-crustal units in this casewould be forced to develop marginal synclines aroundthe dome, to balance the voluminous vertical masstransfer through the ductile crust (Fig. 14c,d), draggingthe upper-crustal Palaeozoic Varied sequences tomid-crustal depths. This model for exhumation of thegranulites in many aspects is similar to that proposedby Behr (1978), Weber (1984) or Franke & Stein (2000)for exhumation of the Saxonian granulites, but it differsfrom the wide range of numerical models focusedon exhumation in collisional domains (e.g. Burg &Podladchikov, 1999; Gerya & Stockhert, 2006; Jamiesonet al., 2007).The specific tectonic history during exhumation ofgranulites in the Moldanubian domain results from acombination of buoyancy and the felsic composition,which together are responsible for their distinct andtransient mechanical behaviour (Franeˇk et al., 2011;Lexa et al., 2011). The abundance of felsic granulitesamong Variscan lower crustal rocks suggests thattheir presence at the base of continental crust isresponsible for the unique tectonic style presented inthis work. The D4 vertical shortening followedimmediately after the D3 granulite ascent, being mostpronounced near the margin of the upper-crustalTepla´-Barrandian block (Figs 13 & 14c,d). Here, theS4 was lubricated by late arc-related magmas andsyenite magma and attained characteristics of a normalshear zone with a dip-slip lineation (Zˇa´k et al.,2005a). This intensive vertical shortening in thecentral parts of the orogenic root can be explained bya ductile thinning mechanism in which the Tepla´-Barrandian suprastructure slid along the normalshear zone to the NW from the mid- and lowercrustal D3 dome described above.The D4 sliding was a mid-crustal expression ofgravitational spreading of the roof (suprastructure) ofthe crustal-scale dome cored by the granulites (infrastructure).The sliding was localized mainly at thethermally weakened volcanic arc, being driven by thefinal D3 dome amplification (Fig. 14c,d). Indeed,recent work by Gerya et al. (2008) and Beaumontet al. (2009) shows that rapidly exhumed, buoyant,Ó 2010 Blackwell Publishing Ltd225

74 J. FRANĚK ET AL.weak lower and ⁄ or subducted HP crust is mechanicallydecoupled from middle-crustal units until thelater stages of its ascent. This is potentially indicatedby the discordance of early structures in the granuliteswith respect to the Monotonous and Varied Grouprocks. In these numerical models, rapid emplacementof these weak, hot rocks as structural domes in themiddle crust drives lateral flow, leading to extensionand ductile thinning above and adjacent to the dome.With these calculations in mind, the fan-like geometryof the S3 fabrics could have been due to the ductilethinning. Exhumation of the Moldanubian rocksfrom below the Tepla´-Barrandian Unit was suggestedby earlier authors (e.g. Scheuvens & Zulauf, 2000;Do¨ rr & Zulauf, 2008), but without the support ofstructural or seismic data covering the broaderMoldanubian domain.CONCLUSIONSThe structural evolution of the South Bohemiangranulites reveals a complexity not seen in the easternpart of the Moldanubian domain. The SouthBohemian granulites were exhumed along two distinctfabrics, the S2 and S3, instead of in a single verticalchannel known from the eastern Moldanubian(Schulmann et al., 2005, 2008). The S2 indicates adistinct older tectonic episode possibly related to anearly subduction period and emplacement of orogeniclower crust at the bottom of the orogenic root (Franeˇket al., 2011). The gravimetry and structural geologyindicate that the individual South Bohemian granulitemassifs reach several kilometres depth. A reflectionseismic profile additionally depicts a region of probablesteep fabrics through the crust below granulites.This vertical region of low reflectivity represents thetrace of the deformed granulite D3 crustal-scale ascentchannel that probably consists of additional deepergranulite bodies.Partially molten lower crust dominated by felsicgranulites was incorporated into middle crustal levelsin the form of a D3 syn-compressional crustal-scaledome at 342–337 Ma. After emplacement, the domeof felsic granulite deformed together with the surroundingmiddle crust. The subhorizontal S4 fabricimmediately reworked the S3 during the interval342–337 Ma, being induced by gravitational spreadingof the growing dome. During D4, the Tepla´-Barrandianupper crust slid to the NW away from the domeregion and allowed ductile thinning of the mid-crustallevel in the Moldanubian domain.This two-stage exhumation mechanism from HPconditions through the orogenic crust for felsic granulitesmay be applicable to exhumation of other HPfelsic rocks. The initial low viscosity probably enabledthe buoyancy-driven rise, while cooling and continuedshearing resulted in development of a large-scaleme´lange of granulite bodies dismembered withinmid-crustal rocks.The presence of low-density felsic granulites at thebottom of the crustal root is a key factor controllingexhumation of the orogenic lower crust and the tectonicstyle, both driven by buoyancy in addition totectonic far-field forces. The Bohemian Massif representsa field laboratory for many conceptual models toexplain the exhumation of orogenic lower crust ingeneral.ACKNOWLEDGEMENTSThe work was supported by a grant from the CzechScience Foundation (GACˇR 205 ⁄ 05 ⁄ 2187) and aninternal project of the Czech Geological Survey(326700). Visits by J. Franeˇk to ULP Strasbourg werefunded by the French Government Foundation(BGF). K. Schulmann and O. Lexa acknowledgefinancial support from the French National GrantAgency (no. 06-1148784) and from the Ministry ofEducation of the Czech Republic (grant MSM-0021620855). Prof. G. Manatschal is acknowledged forinspiring discussions. We are grateful for the constructiveand inspiring reviews of R. Jamieson andM. Pierce, and we thank M. Brown for careful editorialwork.REFERENCESAftalion, M., Bowes, D.R. & Vrána, S., 1989. Early carboniferousU-Pb Zircon age for Garnetiferous, Perpotassic Granulites,Blanský Les Massif, Czechoslovakia. Neues JahrbuchFur Mineralogie-Monatshefte, 4, 145–152.Beaumont, C., Jamieson, R.A., Nguyen, M.H. & Medvedev, S.,2004. Crustal channel flows: 1. Numerical models with applicationsto the tectonics of the Himalayan-Tibetan orogen.Journal of Geophysical Research B: Solid Earth, 109(B6),B06406.Beaumont, C., Nguyen, M.H., Jamieson, R.A. & Ellis, S., 2006.Crustal flow modes in large hot orogens. 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78 J. FRANĚK ET AL.Weber, K., 1984. Variation in tectonic style with time (Variscanand Proterozoic systems, Europe). In: Patterns of Change inEarth Evolution. Report of the Dahlem workshop (eds Holland,H.D. & Trendall, A.F.), pp. 371–386. Springer-Verlag, Berlin.Weber, K. & Behr, H.J., 1983. Geodynamic interpretation of theVariscides. In: Intracontinental Fold Belts (eds Martin, H. &Eder, F.W.), pp. 427–469. Springer-Verlag, Berlin.Wendt, J.I., Kro¨ ner, A., Fiala, J. & Todt, W., 1993. Evidencefrom zircon dating for existence of approximately 2.1 Ga oldcrystalline basement in southern Bohemia, Czech Republic.Geologische Rundschau, 82, 42–50.Wendt, J.I., Kro¨ ner, A., Fiala, J. & Todt, W., 1994. U-Pb zirconand Sm-Nd dating of Moldanubian HP ⁄ HT granulites fromsouth Bohemia, Czech Republic. Journal of the GeologicalSociety, London, 151, 83–90.Zˇák, J., Holub, F.V. & Verner, K., 2005a. Tectonic evolution ofa continental magmatic arc from transpression in the uppercrust to exhumation of mid-crustal orogenic root recorded byepisodically emplaced plutons: the Central Bohemian PlutonicComplex (Bohemian Massif). International Journal of EarthSciences, 94, 385–400.Zˇák, J., Schulmann, K. & Hrouda, F., 2005b. Multiple magmaticfabrics in the Sa´zava pluton (Bohemian Massif, CzechRepublic): a result of superposition of wrench-dominatedregional transpression on final emplacement. Journal ofStructural Geology, 27, 805–822.Zulauf, G., Do¨ rr, W., Fiala, J. & Vejnar, Z., 1997. Late Cadomiancrustal tilting and Cambrian transtension in the Tepla-Barrandian unit (Bohemian Massif, Central European Variscides).Geologische Rundschau, 86, 571–584.Zulauf, G., Bues, C., Do¨ rr, W. & Vejnar, Z., 2002. 10 km minimumthrow along the West Bohemian shear zone: evidencefor dramatic crustal thickening and high topography inBohemian Massif (European Variscides). International Journalof Earth Sciences, 91, 850–864.SUPPORTING INFORMATIONAdditional Supporting Information may be found inthe online version of this article:Appendix S1. Radiometric ages used for histogramcalculations in Fig. 2.Appendix S2. Mineralogical densities used as startingvalues in three-dimensional gravity modelling.Please note: Wiley-Blackwell are not responsible forthe content or functionality of any supporting materialssupplied by the authors. Any queries (other thanmissing material) should be directed to the correspondingauthor for the article.Received 26 February 2010; revision accepted 16 August 2010.Ó 2010 Blackwell Publishing Ltd230

J. metamorphic Geol., 2005, 23, 649–666 doi:10.1111/j.1525-1314.2005.00601.xContrasting textural record of two distinct metamorphic events ofsimilar P–T conditions and different durationsO. LEXA, 1,2 P. ŠTÍPSKÁ, 1,2 K. SCHULMANN, 1,2 L. BARATOUX 1 AND A. KRÖNER 31 Institute of Petrology and Structural Geology, Charles University, Albertov 6, 128 43 Prague, Czech Republic(lexa@natur.cuni.cz)2 Université Louis Pasteur, CGS/EOST, UMR 7517, 1 rue Blessig, Strasbourg 67084, France3 Institut für Geowissenschaften, Universität Mainz, 55099 Mainz, GermanyABSTRACTA structural, metamorphic and geochronological study of the Stare´ Meˇsto belt implies the existence oftwo distinct metamorphic events of similar peak P–T conditions (700–800 °C, 8–10 kbar) during theCambro-Ordovician and the Carboniferous tectonometamorphic events. The hypothesis of two distinctperiods of metamorphism was suggested on the basis of structural discordance between an undoubtedlyCarboniferous granodiorite sill intrusion and earlier Cambro-Ordovician fabrics of a bandedamphibolite complex. The analysis of crystal size distribution (CSD) shows high nucleation density(N 0 ) and low average growth rate (Gt) for Carboniferous mylonitic metagabbros and myloniticgranodiorites. The parameter N 0 decreases whereas the quantity Gt increases towards highertemperatures progressively approaching the values obtained from the Cambro-Ordovician bandedamphibolite complex. The spatial distribution of amphibole and plagioclase shows intense mechanicalmixing for lower-temperature mylonitic metagabbros. In high-temperature mylonites a strong aggregatedistribution is developed. Cambro-Ordovician amphibolites unaffected by Carboniferous deformationshow a regular to anticlustered spatial distribution resulting from heterogeneous nucleation of individualphases. This pattern, together with CSD, was subsequently modified by the grain growth and texturalequilibration controlled by diffusive mass transfer during Carboniferous metamorphism. The differencesbetween the observed textures of the amphibolites are interpreted to be a consequence of the differentdurations of the Carboniferous and Cambro-Ordovician thermal events.Key words: Cambro-Ordovician and Carboniferous metamorphism; quantitative textural analysis;crystal size distributions; grain contact frequencies.INTRODUCTIONIdentifying distinct metamorphic episodes in domainswith a polymetamorphic history is possible providingthe P–T conditions of a younger metamorphic eventare markedly different from those of a preceding one.The problem becomes more complex when the morerecent metamorphic event affects units which previouslysuffered a metamorphic event of a similar grade.In this particular case it is only the geological contextand geochronological data that can indicate the existenceof distinct tectonometamorphic episodes.Polymetamorphic domains are commonly studiedusing analysis of polyphase deformations and individualtectonic events are attributed to distinctdeformational phases (Turner & Weiss, 1963). Infavourable situations tectonic regimes responsible forthe formation of two distinct deformational phases canbe distinguished. However, polyphase structures commonlyresult from the continuous activation of localmechanical instabilities during a single deformationevent (Burg, 1999).The tools of metamorphic petrology are able toprovide P–T estimates of peak metamorphic conditions,and important fragments of P–T paths can bereconstructed when thermodynamic modelling isapplied (Powell & Holland, 1988). However, metamorphicand phase petrology do not reveal informationabout the duration of metamorphic events.The time span between individual metamorphicdeformationevents and the duration of metamorphicreworking can be determined in principle bygeochronology.When metamorphic rocks are reworked by a followingtectonometamorphic event after a significantperiod of time but under similar P–T conditions, it isunlikely that the duration of events, strain rates andkinematics of deformation are also the same. However,the texture of metamorphic tectonites results fromnucleation, grain growth and various recrystallizationmechanisms that are strongly controlled by temperature,time and strain rate/stress ratio (Hickey & Bell,1996). Therefore, the analysis of metamorphic texturesis a method that is capable of distinguishing betweentwo metamorphic events under similar P–T conditionsof different durations.Here, we consider an example of Cambro-Ordovicianmetamorphism associated with crustal thinningÓ 2005 Blackwell Publishing Ltd 649231

650 O. LEXA ET AL.followed by moderate crustal thickening and magmaunderplating during the Carboniferous (Variscan)orogeny (Sˇtı´pska´ et al., 2001). The results of a structural,petrological and geochronological study allow apreliminary distinction of two metamorphic events,which exhibit similar P–T conditions. The structuraland geochronological arguments presented here arenot sufficiently unambiguous to distinguish betweenmetamorphic fabrics developed during the Cambro-Ordovician thinning and the Carboniferous moderatethickening. Therefore, quantitative textural analysis ofpairs of rock samples with similar mineralogical compositionshas been used to show the influence of theCambro-Ordovician and Carboniferous events on finaltextures.GEOLOGICAL SETTINGThe Stare´ Meˇsto belt was a Cambro-Ordovician rift,separating high-grade gneisses of a thickened continentalcrust (Orlica-Snieznik dome) in the west from theNeo-Proterozoic continental margin (Silesian domain)in the east (Fig. 1), at the eastern margin of theBohemian Massif (Kro¨ ner et al., 2000a,b). The Stare´Meˇsto belt consists of a lower crustal complex whichwas later affected by the Variscan collisional tectonics(Kro¨ ner et al., 2000a,b; Sˇtípska´ et al., 2001). Theoverall structure of this unit is marked by NE–SWtrending lithologies generally dipping to the west(Figs 2 & 3). The top of the tectonic sequence and theboundary with the rocks of the Orlica-Snieznik dome isrepresented by a layer of strongly sheared metagabbro(504.9 ± 1.0 Ma; Kro¨ ner et al., 2000a,b). Thisboundary represents a ductile shear zone along which aVariscan granodiorite sill was emplaced (Parry et al.,1997). Zircon from the northern and southern parts ofthe granodiorite sill were dated at 339.4 ± 1.1 and344.5 ± 0.4 Ma, respectively, reflecting the time ofmagma crystallization. The banded amphibolite complexlocated structurally below comprises a layeredsequence of alternating banded amphibolites and finegrainedquartzofeldspathic rocks (503 ± 2 and502.1 ± 1.7 Ma; Kro¨ ner et al., 2000a,b), subordinatetonalitic gneisses (503.3 ± 0.8 and 501.9 ± 0.6 Ma;Kro¨ ner et al., 2000a,b) and high-grade metasedimentsshowing evidence of anatexis (age of metamorphism –c. 504 Ma; Kro¨ ner et al., 2000a,b). In the northernpart of the Stare´ Meˇsto belt the mylonitic metagabbrosare tectonically repeated underneath the bandedamphibolite complex (Fig. 1).(a)(b)Fig. 1. (a) The schematic outlines of the major units of the Bohemian Massif. The location of the studied area is indicated. Upper leftinset shows the position of the studied area in the frame of European Variscides. (b) Geological map of the Stare´ Město belt based ongeological maps 1:25 000 provided by courtesy of the Czech Geological Survey. Important thrust faults and normal faults areindicated. Frames mark location of maps shown in Fig. 2.Ó 2005 Blackwell Publishing Ltd232

CONTRASTING TEXTURAL RECORD OF TWO METAMORPHIC EVENTS 651(a)(b)Fig. 2. Geological and structural maps of the Stare´ Město belt. Foliations, mineral lineations and major thrust and normal faults areindicated. (a) Northern area, see Fig. 1 for location. A–A¢ is the cross-section shown in Fig. 4. (b) Southern area, see Fig. 1 forlocation. B–B¢ is the cross-section shown in Fig. 3.STRUCTURAL CHARACTERIZATION OFVARISCAN AND CAMBRO-ORDOVICIANFABRICSStructural relationships between the granodiorite sill,mylonitic metagabbros and banded amphibolitecomplex combined with geochronological data byKro¨ ner et al. (2000a,b) allowed Sˇtípska´ et al. (2001)to distinguish Cambro-Ordovician and Variscantectonometamorphic events in the study area. Thishypothesis was based on structural relationshipsbetween the banded amphibolite complex and thesuperposed granodiorite intrusion in the northernpart of the Stare´ Město belt (Sˇtı´pska´ et al., 2001). Inthe south, however, the presumed Variscan andCambro-Ordovician fabrics are concordant, and arealmost indistinguishable from each other in the field(Fig. 3b).Ó 2005 Blackwell Publishing Ltd233

652 O. LEXA ET AL.(a)(b)Fig. 3. Geological cross-sections A–A¢ and B–B¢ (shown in Fig. 2) of the Stare´ Město belt disclosing major structures, lithology of individual units and major tectonic boundaries.Equal-area, lower-hemisphere stereoplots show D 1 ,D 2 and D 3 planar and linear structures. Each stereoplot is contoured at regular multiples of distribution. Drawings are madeaccording to field photographs and field notes and show principal structural features.Ó 2005 Blackwell Publishing Ltd234

CONTRASTING TEXTURAL RECORD OF TWO METAMORPHIC EVENTS 653Structures in Variscan granodioriteThe Variscan granodiorite sill is syntectonicallyemplaced between the mylonitic metagabbro to thewest, and the banded amphibolite complex to the east(Parry et al., 1997). The structural pattern of thegranodiorite sill varies from SW to NE along thelength of the intrusion (Figs 2 & 3). The magmatic tosubsolidus foliation (S 2 ) dips to the NW and bears anintense subhorizontal mineral lineation (L 2 ) defined byan alignment of amphibole (Fig. 3a,b). Rare lock-upshear-bands filled with residual melt indicate a dextralsense of shear. In the south, the granodiorite does notform a continuous body but occurs as numerous sills insurrounding amphibolites and anatectic metasediments(Fig. 3b). The granodiorite locally forms dykesemplaced along conjugate steep brittle–ductile NW–SEor NE–SW trending shear zones at high angle to theinitially flat-lying S 1 foliation (Fig. 3a,b).Structures of mylonitic metagabbrosAlthough the western metagabbro belt was stronglyreworked during the Variscan orogeny, magmaticstructures are preserved in the low-strain domains(Fig. 4c). Here, the metagabbro exhibits either amedium-grained isotropic texture and homogeneouscomposition, or layering marked by alternating layersof variable grain size ranging from a few mm up to 2 cm(Fig. 4c). The mylonitic metagabbros exhibit thedevelopment of S 2 mylonitic foliation, defined byalternating monomineralic ribbons of recrystallizedplagioclase and amphibole (Figs 4c & 5c). This foliationdips at medium to high angles to the WNW.Towards the south, the foliation tends to becomesubvertical (Fig. 2b). A strong mineral L 2 lineation isdefined by recrystallized aggregates of amphibole and isassociated with numerous kinematic indicators suggestingdextral shearing. The S 2 foliation is locallyaffected by late F 3 folds with hinges sub-parallel to L 2lineation or by development of S 3 crenulation cleavage.The complete structural and kinematic coherency betweenthe metagabbros and underlying granodiorite sillsuggests their common Variscan deformation history.The eastern metagabbro was affected by heterogeneousshear zones (Fig. 4d) during the Variscan deformation.The metagabbros are strongly mylonitized atthe upper and lower contacts with the hangingwall andFig. 4. Field photographs showing main structural features of important lithologies of the Stare´ Meˇsto belt. (a) Metamorphic layeringM 1 of banded amphibolite complex. (b) Extensional melt-filled shear-bands D 1 in tonalitic migmatitic gneiss of the banded amphibolitecomplex. (c) Coarse-grained C-O metagabbro (left photograph) mylonitized during D 2 (right photograph) from western belt. (d)Coarse-grained C-O metagabbro mylonitized by localized shear zones during D 2 from eastern belt.Ó 2005 Blackwell Publishing Ltd235

654 O. LEXA ET AL.(a)(b)(c)(d)Fig. 5. Photomicrographs of characteristic mineral assemblages and textures from the Stare´ Meˇsto belt: (a) M 1 assemblage Cpx–Amp–Pl–Qtz of layered amphibolite showing equilibrated annealed texture. (b) M 1 assemblage Grt–Bt–Sil in anatectic paragneiss. (c)Variscan dynamic M 2 recrystallization of C-O metagabbro from the western belt leading to strong mineral banding of amphibole andplagioclase. (d) Variscan dynamic M 2 recrystallization of amphibole and plagioclase in C-O metagabbro from the eastern belt markedby strong cataclastic grain-size reduction.footwall metapelites, respectively, while magmaticstructures are preserved in the inner part. The geometryof S 2 fabric and kinematics of deformation are identicalwith those of the western metagabbro sheet (Fig. 3).Structures of the banded amphibolite complexThe banded amphibolite complex is affected by a hightemperaturemetamorphic event characterized by arange of melt-collecting structures. The S 1 foliation ismarked by a metamorphic compositional layering bestseen in the banded amphibolites and in the tonaliticmigmatites (Figs 3a & 4a). In the northern part of thestudy area, S 1 dips to the NNE at shallow angles(Figs 2a & 3a). To the south, the S 1 moderately dips tothe west in the migmatitic metasediments and is generallyflat-lying in the banded amphibolites (Figs 2b &3b). The S 1 compositional and metamorphic layering islocally folded into rootless isoclinal F 1 folds or isaffected by asymmetrical pinch and swell boudinagewith melt collecting in some neck-zones. In the tonaliticgneisses and more rarely in the layered amphibolitesthe planar fabric is affected by extensional shearzones filled with an amphibole-bearing tonalitic melt,or the planar fabric dissipates in coarse-grained meltpatches (Fig. 4b).The Variscan D 2 deformation in the bandedamphibolite complex is marked by folding of theoriginally flat-lying Cambro-Ordovician S 1 foliation.These open to close F 2 folds strike NE–SW with subhorizontalhinges and wavelengths ranging from a fewmetres to several tens of metres (Fig. 3). Fold hingesare locally cut by subvertical, NE–SW or NW–SEtrending, conjugated shear and fracture zones filledwith granitic melts. The internal foliation within theseshear zones bears a NE–SW trending stretching L 2lineation (Fig. 3). Similar to mylonitic metagabbros,older foliations are locally affected by late F 3 foldswith hinges sub-parallel to the L 2 lineation or bydevelopment of a late S 3 crenulation cleavage subparallelto the S 2 foliation, indicating deformationcontinuity towards lower metamorphic conditions.Ó 2005 Blackwell Publishing Ltd236

CONTRASTING TEXTURAL RECORD OF TWO METAMORPHIC EVENTS 655Table 1. Mineral assemblages and textures of samples used for thermobarometry.Sample Unit Locality Rock type Texture Grt Pl Hbl Cum Cpx Qtz Bt Kfs Sil Rt Ilm Ttn Mag OpxSamples associated with D1 Cambro-Ordovician structureLAC4a LAC 66 g Grt-amphibolite Preferred shape orientation,· · · – – · – – – – · – · –annealedLAC4b LAC S149e Cpx-amphibolite Weak shape preferred orientation, – · · – · · – – – – · – · –annealedLAC4c LAC S149f Melt patches within Random orientation – · · – – · – – – – – – – –amphiboliteTG1a LAC S24-1b Tonalitic migmatite Weak shape preferred orientation, · · · – – · · – – – – – – –annealedTG1b LAC S24-5a Tonalitic migmatite Weak shape preferred orientation, · · · – – · · – – – – – – –annealedTG1c LAC S24-6e Melt patches within Random orientation – · · – – · · · – – – · · –migmatiteMP1 LAC Heg3 Metapelite Weak shape preferredorientation, annealed· · – – – · · · · · · – – –Samples associated with D2 Variscan structureGAHT Western S33b Grt-amphibolite Strong preferred orientation, dynamic · · · · – · – – – – · – · –gabbroic beltrecrystalliation of Pl, QtzGHT3 Western S151 Metagabbro Strong preferred orientation of Amp, – · · · · – – – – – · · – ·gabbroic beltdynamic recrystalliation of Pl, AmpGLT2 Eastern S130 Metagabbro Strong preferred orientation of Amp, – · · – – – – – – – · · – –gabbroic beltdynamic recrystalliation of Pl, AmpT3 Sill Granodiorite Dynamic recrystallization of Pl, Qtz – · · – – · · · – – · · · –PETROLOGY, METAMORPHIC TEXTURES ANDP–T ESTIMATESSample selection for petrological study and analyticalproceduresThe study samples were selected according to theirstructural position with respect to assumed Cambro-Ordovician and Variscan fabrics. Mineral analyseswere carried out on a CAMECA SX 50 at ETH Zu¨ richand a JEOL microprobe at the University of Mainz.Operating conditions were 15 kV acceleration voltageand beam current of 20 nA. Representative sampleswith mineral assemblages are given in Table 1 andcorresponding mineral compositions are summarizedin Table 2. Mineral abbreviations used in text andtables follow Kretz (1983). Representative mineralanalyses are listed in Tables 3–6. Amphibole formulaewere calculated after Holland & Blundy (1994) andclassified according to Leake et al. (1997). The geothermometersand geobarometers used for calculationsare given in Table 7. For each sample, five to 10 sets ofanalyses were used for P–T calculations and the resultsare shown in Table 7 and Fig. 6.Metamorphic textures and mineral compositions of thegranodiorite sill and mylonitic metagabbrosThe granodiorite sill is composed of Pl + Qtz +Amp + Bt + Kfs + Ttn + Mag + Ilm + Ap. Itshows magmatic, sub-magmatic and solid-statemicrostructures marked by dynamic recrystallizationof plagioclase, amphibole and quartz (Parry et al.,1997). Amphibole correspond to magnesiohornblendeand tschermakite (i.e. Si p.f.u. ¼ 6.3–6.7, X Mg ¼ 0.54–0.64) and plagioclase is An 34)37 andesine.The proportions of plagioclase and amphibole in thewestern mylonitic metagabbro vary significantly andthe rock locally consists of up to 90% plagioclase or90% amphibole. Minor clinopyroxene occurs in thecore of large amphibole. Titanite is an abundantaccessory mineral. Non-recrystallized amphibole isinterpreted as magmatic in origin and the compositioncorresponds to a magnesiohornblende (i.e. Si p.f.u. ¼7.00–7.25, X Mg ¼ 0.84–0.89). The magmatic plagioclaseis a An 55)62 labradorite.The mineral assemblage of highly reworked metagabbrowith banded mylonitic structure (Figs 4c & 5c)includes Pl + Amp ± Cum ± Cpx ± Opx ± Grt ±Ttn ± Rt ± Ilm. A granulite facies mineral assemblage,comprising saphirine and corundum, isre-equilibrated during subsequent amphibolite faciesreworking (Baratoux et al., 2005). Both amphibole andplagioclase show strong compositional variationsbetween those of old magmatic and metamorphicgrains (Table 1). Rare garnet-bearing amphibolitedeveloped through the deformation and metamorphismof a tonalitic migmatitic gneiss associatedwith metagabbro in the lower part of the gabbroicsheet. It exhibits mineral assemblage Hbl ±Cum ± Grt + Pl + Qtz ± Mag ± Ilm (Table 1)which allows the pressure during the Variscan metamorphismto be estimated.Recrystallized plagioclase exhibits an increase in theanorthite content from the core towards the rim inthe metagabbro (An 37 fi 60 ) and garnet-amphibolite(An 15 fi 27 ). Recrystallized amphibole in the metagabbrocorresponds to tschermakite, and tschermakite orferrotschermakite in the garnet-amphibolite. Garnetshows an increase in almandine, grossular, pyrope andX Mg (Fig. 7d), accompanied by depletion of spessar-Ó 2005 Blackwell Publishing Ltd237

656 O. LEXA ET AL.Table 2. Summarized compositions of minerals used for thermobarometry.Sample Rock type Locality Grt Pl Amp Cpx Bt IlmXMg Alm Py Grs Sps An Compositional name a XMg Si (pfu) XMg XMg XIlmSamples associated with D1 Cambro-Ordovician structureLAC4a Grt-amphibolite 66 g 0.16–0.18 62–65 12–14 11–21 4–5 31–33 Ts, Fe-Ts 0.47–0.54 6.27–6.37 – – –LAC4b Cpx-amphibolite S149c – – – – – 44–50 Mg-Hbl, Ts, Ed, Prg 0.59–0.68 6.36–6.71 0.67 b – –LAC4c Melt patche within amphibolite S149f – – – – – 38–43 Ts 0.63–0.74 6.25–6.50 – – –TG1a Tonalitic migmatite S24-1b 0.10–0.11 55 (51) 7 (8) 28 (29) 6 (13) 33–35 Fe-Prg, Hs 0.34–0.37 5.8–6.2 – 0.34–0.36 –TG1b Tonalitic migmatite S24-5a 0.11–0.12 58 (50) 8 (6) 25 (31) 9 (14) 31–34 Fe-Prg, Hs 0.25–0.40 5.8–6.2 – 0.33–0.35 –TG1c Melt patche within migmatite S24-6e – – – – – 32–35 Fe-Prg, Hs 0.34–0.38 5.95–6.25 – 0.32–0.33 –MP1 Metapelite Heg3 0.15 (0.26–0.28) b 75 (69) 15 (26) 4 (4) 5 (1) 24–26 – – – – 0.50–0.53 0.95–0.98– – –6.1–6.37.7–7.96.0–6.35.30.46–0.670.48–0.500.82–0.880.64Samples associated with D2 Variscan structureGAHT Grt-amphibolite S33b 0.15 (0.05) 70 (57) 12 (3) 15 (27) 3 (12) 15–27 (15) Fe-Ts, TsCum– – –GHT3 Metagabbro S151 – – – – – 37–60 Ts,CumGLT2 Metagabbro S130 – – – – – 48–62 (25) Mg-Hbl, Ts 0.74–0.88 6.24–6.64 – – –T3 Granodiorite – – – – – 34–37 Mg-Hbl, Ts 0.54–0.64 6.3–6.7 – – –Values in brackets represent core compositions of zoned minerals. Notes: Alm ¼ 100 · Fe 2+ /(Mg + Ca + Mn + Fe 2+ ), An ¼ 100 · Ca/(Ca + Na + K), XMg ¼ Mg/(Mg + Fe 2+ ).a Amphibole compositional name from classification of Leake et al. (1997): Prg, pargasite; Ts, tschermakite; Ed, edenite; Hs, hastingsite; Tr, tremolite; Mg-Hbl, magnesio-hornblende.b XMg ¼ Mg/(Mg + Fe tot ).tine from the core to the rim (i.e. Alm 57 fi 70Grs 2 fi 15 Py 3 fi 12 Sps 12 fi 3 ; X Mg ¼ 0.05 fi 0.15).A relic magmatic assemblage in the eastern myloniticmetagabbro is represented by Pl + Hbl ±Cpx ± Ttn ± Ilm. Magmatic plagioclase (0.5–5 mm)shows zoning marked by a rimward increase ofanorthite from 50 to 60%, whereas recrystallization ofthe plagioclase (0.05–0.1 mm) is accompanied by adecrease in anorthite content (i.e. An 45 ). Magmaticamphibole (0.5–5 mm) with pyroxene relics in thecores is magnesiohornblende (6.8–7.0 Si p.f.u., X Mg ¼0.72–0.77). Recrystallized grains (0.02–0.1 mm) correspondto magnesiohornblende with a slightly lowercontent of Si and X Mg (6.6–6.7 Si p.f.u., X Mg ¼ 0.69–0.74) in comparison with original grains.P–T conditions of Variscan metamorphismThe pressure of crystallization of the granodiorite (T3)was calculated as 6.5–7.0 kbar using the Al-content ofamphibole (Schmidt, 1992) at a temperature of 670–725 °C inferred from the Pl–Hbl thermometer ofHolland & Blundy (1994). In the western mylonitemetagabbro the temperature was estimated as770 ± 50 °C using the thermometer by Holland &Blundy (1994) in metagabbro (GHT3), while in garnetamphibolitethe temperature was estimated as 710–740 °C (GAHT). A similar temperature of 750 ±50 °C is reported by Baratoux et al. (2005) for 40amphibole–plagioclase couples. Pressure conditionsbetween 8 and 10 kbar were calculated only in Qtzbearinggarnet-amphibolite by means of the Kohn &Spear (1990) barometer. Based on 30 amphibole–plagioclase couples from the eastern mylonite metagabbro,Baratoux et al. (2005) estimated temperatureas 650 ± 50 °C using the Pl–Hbl thermometer ofHolland & Blundy (1994). The pressure could not becalculated because of the lack of garnet.Mineral textures and mineral compositions of rocks of thebanded amphibolite complexMineral assemblages and selected mineral compositionsof all study rocks are given in Table 1 and summarizedin Table 2, respectively.Banded amphibolites are characterized by the mineralassemblage Amp + Pl + Qtz ± Cpx ± Grt ±Mag ± Ilm. Three types of microstructures marked bystraightened grain boundaries without indications ofdynamic recrystallization are distinguished: (1) Equigranularaggregates (grain size 0.05–1.0 mm) ofamphibole, plagioclase and quartz (Fig. 5a); (2)amphibole-rich layers alternating with elongateaggregates (grain size 0.05–1.0 mm) of quartz andplagioclase; and (3) coarse-grained aggregates (grainsize 1–5 mm) of randomly distributed plagioclaseand quartz sometimes with amphibole occur in meltpatches. In types (1) and (2) the mineral grains arealigned and elongated parallel to the compositionalÓ 2005 Blackwell Publishing Ltd238

CONTRASTING TEXTURAL RECORD OF TWO METAMORPHIC EVENTS 657Table 3. Representative electron probe analyses of garnet.Garnet analyses recalculated to 12 oxygenEvent C-O Varis.Sample/analyses TG1a/grt23-rim TG1a/grt44-core TG1b/grt27-rim LAC4a/grt43-rim MP1/grt379-core a MP1/grt313-rim a GAHT/grt5-15-core GAHT/grt114-rimSiO 2 37.96 37.65 37.44 37.95 37.69 36.79 37.14 37.47TiO2 0.06 0.13 0.14 0.07 0.08 0.08 0.07 0.04Al 2 O 3 20.64 20.53 20.77 20.46 21.21 20.82 21.17 20.98FeO 26.47 24.70 27.05 30.02 32.60 35.12 25.97 32.46MnO 4.28 5.78 5.19 1.84 0.56 2.25 6.15 1.25MgO 1.80 1.54 1.75 3.06 6.54 3.62 0.85 3.12CaO 10.05 10.28 8.31 7.34 1.33 1.33 9.92 5.28Total 101.26 100.61 100.65 100.74 100.01 100.01 101.27 100.60Si 2.99 2.98 2.98 3.00 2.97 2.97 2.94 2.97Ti 0.00 0.08 0.01 0.00 0.00 0.00 0.00 0.00Al 1.92 1.92 1.95 1.90 1.97 1.98 1.98 1.96Fe 3+ 0.10 0.11 0.09 0.09 0.00 0.00 0.13 0.09Fe 2+ 1.65 1.53 1.71 1.89 2.15 2.37 1.59 2.07Mn 0.29 0.39 0.35 0.12 0.04 0.15 0.41 0.08Mg 0.21 0.18 0.21 0.36 0.77 0.44 0.10 0.37Ca 0.85 0.87 0.71 0.62 0.11 0.12 0.84 0.45a Fe 3+ not calculated.Table 4. Representative electron probe analyses of amphibole.Amphibole analyses recalculated after Holland & Blundy (1994)Event C-O Varis.Sample/analyses LAC4a/amp37 LAC4b/amp16 LAC4c/amp482 TG1a/amp45 TG1b/amp12-6 TG1c/amp64 GAHT/amp115 GAHT/cum6-7 a GHT3/amp48 GLT2/amp-102 T3/amp18SiO 2 42.06 45.27 42.58 39.82 39.12 39.38 41.64 51.85 43.34 45.11 43.59TiO 2 1.59 1.47 1.75 0.91 1.05 0.94 0.53 0.05 0.20 0.61 1.26Al2O3 11.58 9.68 11.73 12.96 13.51 13.40 12.42 1.35 17.66 13.94 11.85FeO 20.83 15.88 17.00 23.31 23.52 23.46 22.17 27.34 11.03 9.82 16.03MnO 0.12 0.32 0.24 0.57 0.54 0.46 0.46 0.59 0.13 0.18 0.24MgO 7.87 11.69 10.54 5.66 5.45 5.55 7.28 13.95 12.09 13.98 10.89CaO 10.62 11.34 10.68 11.13 11.41 11.31 10.57 2.51 11.29 12.42 11.05Na2O 2.11 1.50 1.86 1.45 1.33 1.39 1.77 0.18 1.95 1.71 1.25K 2 O 0.07 0.55 0.78 1.47 1.65 1.63 0.23 0.01 0.27 0.32 1.48Total 96.85 97.70 97.16 97.28 97.58 97.52 97.07 97.83 97.95 98.09 97.64Si 6.38 6.65 6.34 6.16 6.05 6.09 6.30 7.75 6.15 6.43 6.44Ti 0.18 0.16 0.20 0.11 0.12 0.11 0.06 0.01 0.02 0.07 0.14Al 2.07 1.68 2.06 2.36 2.46 2.44 2.22 0.24 2.95 2.34 2.07Fe 3+ 0.58 0.50 0.61 0.66 0.70 0.67 0.91 0.19 0.70 0.34 0.51Fe 2+ 2.06 1.45 1.50 2.36 2.34 2.36 1.90 3.23 0.61 0.83 1.47Mn 0.02 0.04 0.03 0.08 0.07 0.06 0.06 0.08 0.02 0.02 0.03Mg 1.78 2.56 2.34 1.30 1.26 1.28 1.64 3.11 2.56 2.97 2.40Ca 1.73 1.79 1.70 1.84 1.89 1.87 1.71 0.40 1.72 1.90 1.75Na 0.62 0.43 0.54 0.44 0.40 0.42 0.52 0.05 0.54 0.47 0.36K 0.01 0.10 0.15 0.29 0.33 0.32 0.04 0.00 0.05 0.06 0.28a Cummingtonite recalculated on the basis of 23 oxygen and 15 cations + Na + K.layering. Regardless of textural type, the minerals arenot zoned. Almandine-rich garnet reveals flat chemicalprofiles (Alm 62)65 Grs 11)21 Py 12)14 Sps 4)5 ; X Mg ¼ 0.16–0.18; Fig. 7a). Amphibole is magnesiohornblende ortschermakite, rarely edenite or pargasite. Plagioclase isandesine (An 30)50 ) and X Mg of rare clinopyroxeneequals 0.65–0.67.Tonalitic migmatitic gneiss is made up ofQtz + Pl + Bt + Amp ± Grt. Its structure is characterizedby alternations of leucosome, mesosome andmelanosome layers several mm to 1 cm in thickness.Weak layering is defined by amphibole–plagioclaseaggregates alternating with quartz-rich domains(Fig. 4b). Plagioclase and amphibole are elongated(length 0.2–3 mm) in the leucosome and restitic layers,whereas plagioclase is isometric (up to 3 mm). Biotite(0.2–1 mm) is oriented parallel to the foliation togetherwith amphibole in melanosome. Garnet porphyroblasts(0.3 mm) occur in mesosome and melanosome. Largecoarse-grained quartz aggregates occurs in leucosome.Chemical profiles across garnet show slight zoning withcores being more spessartine-rich; the X Mg ratio is constant(0.10–0.12) (Fig. 7c). Amphibole is homogeneousferropargasite or hastingsite, regardless of the texturalposition and type of aggregate. X Mg in biotite rangesbetween 0.32 and 0.36, and the Ti content between 0.14and 0.22 p.f.u. The composition of plagioclase in boththe leucosome and melt patches is An 31)35 andesine.Ó 2005 Blackwell Publishing Ltd239

658 O. LEXA ET AL.Table 5. Representative electron probe analyses of plagioclase.Event C-O VarisSample/analysesLAC4b/pl13LAC4c/pl82-1LAC4a/pl50TG1a/pl46-5TG1b/pl26TG1c/pl2-8MP1/pl137GAHT/pl113 -rimGHT3/pl122-coreGHT3/pl101-rimGLT2/pl122-coreGLT2/pl101-rimGLT2/pl79-coreT3/pl19SiO2 56.92 58.57 60.07 59.94 60.11 59.82 61.64 62.51 59.02 52.36 64.73 56.05 62.00 59.35TiO 2 0.00 0.03 0.00 0.01 0.01 0.00 0.01 0.03 0.00 0.00 0.00 0.00 0.01 0.01Al2O3 27.18 26.20 25.01 25.54 25.26 25.09 24.01 23.64 25.92 30.48 22.01 27.73 23.42 26.19FeO 0.22 0.13 0.35 0.05 0.17 0.17 0.00 0.35 0.00 0.00 0.07 0.18 0.00 0.05CaO 9.74 8.53 6.60 7.09 6.99 7.31 5.28 4.95 7.17 12.18 3.27 10.09 5.26 7.02Na2O 6.30 6.73 8.16 7.65 7.73 7.44 8.76 8.77 7.63 4.50 9.70 5.93 8.66 7.59K 2 O 0.12 0.30 0.05 0.18 0.12 0.35 0.20 0.07 0.00 0.00 0.15 0.03 0.12 0.15Total 100.48 100.49 100.19 100.46 100.39 100.18 99.90 100.32 99.74 99.52 99.93 100.01 99.47 100.36Si 2.54 2.61 2.66 2.66 2.67 2.66 2.73 2.76 2.64 2.37 2.86 2.52 2.76 2.63Ti 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00 0.00Al 1.43 1.38 0.00 1.34 1.32 1.32 1.25 1.23 1.36 1.63 1.15 1.47 1.23 1.37Fe 3+ 0.01 0.00 0.01 0.00 0.01 0.01 0.00 0.00 0.00 0.00 0.00 0.01 0.00 0.00Ca 0.47 0.41 0.31 0.34 0.33 0.35 0.25 0.23 0.34 0.59 0.15 0.49 0.25 0.33Na 0.55 0.58 0.70 0.66 0.66 0.64 0.75 0.75 0.66 0.40 0.83 0.52 0.75 0.65K 0.01 0.02 0.00 0.01 0.01 0.02 0.01 0.00 0.00 0.00 0.01 0.00 0.01 0.01Table 6. Representative electron probe analyses of biotite andclinopyroxene.Bt analyses recalculated to 11 oxygen, cpx to 6 oxygenEventThe mineral assemblage of the migmatitic paragneissesis Pl + Bt + Qtz ± Grt ± Kfs ± Sil ± Rt ±Ilm. Plagioclase and quartz form large isometric grainsup to 10 mm in size. With an increasing degree ofanatexis, quartz and plagioclase (0.3–5 mm) coalesceto form lenses bordered by biotite-rich rims occasionallycontaining sillimanite. Both the leucosome andmelanosome aggregates contain garnet. In thenebulitic stage diffuse pockets of leucosome of differentsize developed. Garnet forms poikilitic crystals(0.1–1 cm) with inclusions of quartz and biotite. ScarceK-feldspar occurs in leucosome lenses. The distributionof elements in garnet cores is homogeneous(Alm 69 Grs 4 Py 26 Sps 1 ; X Mg ¼ 0.26–0.28) but within theouter 100 lm rim or around biotite inclusions garnet isC-OMineral cpx btSample/analyses LAC4b/cpx20 TG1a/bt25 TG1b/bt45 MP1/bt343SiO 2 50.88 33.95 35.10 36.06TiO 2 0.24 2.92 3.15 1.57Al2O3 2.35 15.71 15.76 19.78FeO 11.11 25.22 24.39 18.03MnO 0.35 0.30 0.35 0.05MgO 12.04 7.23 7.37 10.83CaO 22.14 0.02 0.07 0.00Na2O 0.58 0.04 0.05 0.22K 2 O 0.01 9.51 9.40 9.03Total 99.70 94.90 95.64 95.57Si 1.93 2.74 2.80 2.75Ti 0.01 0.18 0.19 0.09Al 0.11 1.49 1.48 1.78Fe 3+ 0.03 0.00 0.00 0.00Fe 2+ 0.32 1.70 1.63 1.15Mn 0.01 0.02 0.02 0.00Mg 0.68 0.87 0.88 1.23Ca 0.90 0.00 0.01 0.00Na 0.04 0.01 0.01 0.03K 0.00 0.98 0.96 0.88enriched with almandine and spessartine and depletedin pyrope (Alm 75 Grs 4 Py 26 Sps 5 ; X Mg ¼ 0.15; Fig. 7b).This zoning is attributed to continuous Fe–Mgre-equilibration between garnet and biotite duringcooling. The X Mg of matrix biotite ranges from 0.50 to0.53 and its Ti-content varies between 0.07 and0.21 p.f.u. Plagioclase is a homogeneous An 28)31oligoclase-andesine.P–T conditions of Cambro-Ordovician metamorphismThe temperature of metamorphism in amphibolites(LAC4a, LAC4b and LAC4c) was estimated with thePl–Hbl thermometer by Holland & Blundy (1994). Thetemperature ranges of 725–825 and 730–860 °C wereinferred for amphibolites and melt patches, respectively.A comparable temperature span of 715–770 and750–770 °C was established for the tonalitic gneissand associated melt patches, respectively (TG1a, TG1b& TG1c). The pressure in amphibolites was estimatedas 8.5–10 kbar, and in tonalitic gneisses as 9–10 kbarusing the Grt–Hbl–Pl–Qtz barometer (Kohn & Spear,1990). A similar range of 8–10 kbar was obtained fromthe aluminium content of amphibole (Schmidt, 1992)within the tonalitic gneisses. Similarly, a pressure 8–9 kbar was estimated from melt patches crystallized inthe small-scale shear zones. A pressure of 7.5 kbar wasobtained from garnet core and matrix plagioclase usingthe GASP barometer (Koziol, 1989), calculated bymeans of the ÔThermobarometryÕ software (Spearet al., 1991). Pressure calculated using the GRAILbarometer (Bohlen et al., 1983) gives a range of 7.5–8 kbar for garnet cores containing rutile and ilmeniteinclusions.QUANTITATIVE TEXTURAL ANALYSISA quantitative microstructural analysis was carried outon three samples of the banded amphibolite complex,assumed to be exclusively the result of the Cambro-Ó 2005 Blackwell Publishing Ltd240

CONTRASTING TEXTURAL RECORD OF TWO METAMORPHIC EVENTS 659Table 7. Results of P–T calculations using the following geothermometers and geobarometers: HB, Holland & Blundy (1994); T, Thompson (1976); FS, Ferry & Spear (1978);NW, O’Neill & Wall (1987); GN, Gasparik & Newton (1984); KS (Mg), Kohn & Spear (1990); KS (Fe), Kohn & Spear (1990); K, Koziol (1989); Bh, Bohlen et al. (1983); S,Schmidt (1992). Errors are given in table headings.Grt–Hbl–Pl–Qtz (KS) Al in Hbl (S) GASP (K) GRAIL (Bh)Pl–Hbl (HB) Grt-Bt (T) Grt-Bt (FS) P calculated at aT (°C) ofSample Rock type Combination T calculated at aP (kbar) ofT (Ed-Tr) T (Ed-Ri) ± 50 °C ± 50 °C P (Mg) P (Fe) – – –±40°C ± 40 °C ± 0.5 ± 0.5 ± 0.6 Not given ± 0.5Samples associated with D1 Cambro-Ordovician structureLAC4a Grt-amphibolite Rims 10 731–770 772–823 – – 700 8.6–9.0 9.5–9.8 – – –LAC4b Cpx-amphibolite Rims 10 727–787 775–825 – – 700 – – – – –LAC4c Melt patches within amphibolite Rims 10 732–765 796–857 – – – – – – – –TG1a Tonalitic migmatite Rims 10 715–755 732–772 700–738 685–737 700 9.7–10.0 9.8–10.1 8.2–10.0 – –TG1b Tonalitic migmatite Rims 10 717–762 712–768 712–727 702–722 700 9.2–9.9 9.5–9.8 8.4–10.0 – –TG1c Melt patches within migmatite Rims 10 750–756 759–770 – – – – – 7.9–8.7 – –MP1 Metapelite Cores 7 – – 788–840 838–905 800 – – – 7.3–7.4 7.6–8.2Rims 7 – – 581–616 549–593 – – – – – –Samples associated with D2 Variscan structureGAHT Grt-amphibolite Rims 10 711–730 724–737 – – 700 7.8–8.4 9.6–10.1 – – –GHT3 Metagabbro Rims 10 716–836 711–837 – – – – – – – –T3 Tonalite Rims 6 671–691 714–724 – – – – – 6.5–7.0 – –Fig. 6. P–T diagram with blank boxes for C-O metamorphism.P–T estimates for Variscan assemblages are presented in greyboxes. Dashed lines show temperature estimates from metagabbrosand from metaperidotites. Numbers correspond tosamples presented in Tables 1, 2 and 7.Ordodvician metamorphism. We investigated also sixsamples representing the Variscan granodiorite, tonalitictectonites and mylonitic metagabbros, in which,only the Variscan metamorphic history is assumed tobe preserved. Polished thin sections of the samplesparallel to lineation and perpendicular to foliation (XZsection) were prepared for optical microscopy andback-scattered electron image (BSEI) analysis. TheBSEI analysis was carried out using the Camscan S4instrument fitted with a high-resolution backscatterdetector. Line drawings of optical microscopy photographsand BSE images were digitized (Fig. 8) andprocessed in ESRI ArcView desktop GIS in order toobtain a map of boundaries between individual grains(Lexa, 2003). The analysis of crystal size distributions(CSD), shape preferred orientation of grains (SPO)and grain boundaries preferred orientation (GBPO), aswell as grain contact frequencies were carried out inPolyLX MATLAB TM Toolbox (Lexa, 2003).Crystal size distributionsThe theory of CSD is a well-established method inchemical engineering, metallurgy and ceramics toreveal information about nucleation, growth ratesand growth times of crystals (Randolph & Larson,1971). CSD in many metamorphic and igneous rocksshow a loglinear relationship between grain size Land population density N according to the followingequationN ¼ N 0 e L=Gt ;ð1ÞÓ 2005 Blackwell Publishing Ltd241

660 O. LEXA ET AL.Fig. 7. Chemical profiles of garnet: (a) inlayered amphibolite showing flat patterns,(b) in sillimanite-bearing paragneiss withwell-developed flat profiles in core andmargin affected by late diffusion, (c) intonalitic migmatitic gneiss and (d) inGrt-amphibolite within metagabbro.Numbering of samples corresponds tonumbers in Tables 1, 2 and 7. See text fordiscussion.where N 0 and Gt are constants and may be related tonucleation density and growth rate of crystal (Cashman& Ferry, 1988; Marsh, 1988).CSD plots of all samples constructed according tothe method by Peterson (1996) exhibit linear correlationsbetween logarithm of population density (i.e.number of crystals per size per volume) and crystalsize (Fig. 9b). Therefore, applying the theory ofCSD, such distributions could be parameterized byzero size intercept N 0 (nucleation density) and slopeGt (growth rate multiplied by time). These twoparameters plotted in a N 0 –Gt space (Fig. 9) exhibitan inverse correlation, and the samples with acomparable metamorphic grade and deformationhistory form distinct clusters. The metagabbros(samples GLT1 & GLT2, Fig. 8b) sheared during theVariscan episode from the eastern mylonitic belt(c. 650 ± 50 °C) form a group with the highest N 0and lowest Gt values, while metagabbros (samplesGHT1 & GHT2) adjacent to the granodiorite sill(c. 750 ° ±50°C) exhibit a decrease in N 0 in conjunctionwith increasing Gt. The samples LAC1,LAC3 and TG2 (Fig. 8a,d) of banded amphibolitesand amphibole-bearing tonalitic gneisses thatoriginated during the Cambro-Ordovician metamorphism(c. 750–850 °C) plot in the centre of the diagram(lower N 0 , higher Gt). The undeformedVariscan granodiorite (sample T1, Fig. 8c) exhibitslowest values of N 0 and highest values of Gt.Grain shapes, shape preferred orientation and grainboundary preferred orientationPlagioclase is a low-symmetry mineral characterizedby weakly elongated to round shapes under mostmetamorphic conditions. Degree of elongation ofplagioclase grains is therefore an indication of intracrystallinedeformation. In contrast, amphibole ishighly elongate in low-metamorphic grade rocks, whileat high-grade ones it becomes spherical (Brodie &Rutter, 1987). Therefore, shape analysis combinedwith SPO provides information about the degree ofdeformation of both plagioclase and amphiboleaggregates.Both grain shapes and SPO were evaluated using thePolyLX toolbox (Lexa, 2003). The grain shapes arecharacterized by preferred orientation of the long axesand by mean axial ratio R of the best-fit ellipse on theindividual grains using the area-moments ellipse fittingmethod. To evaluate an overall single phase SPO,individual linear segments of grain boundaries ofindividual phases were treated as independent vectorscontributing to the bulk Scheidegger–Watson orientation-tensor.The eigenvalue analysis determines twoeigenvalues and assigns them to mutually perpendiculardirections. The ratio of eigenvalues R e is consideredas a rough estimate of the degree of SPO forthe given phase. Grain boundary preferred orientationwas assessed by a similar technique, with the bulkÓ 2005 Blackwell Publishing Ltd242

CONTRASTING TEXTURAL RECORD OF TWO METAMORPHIC EVENTS 661(a) (b)(c) (d)(d)Fig. 8. Representative digitized microstructures used in textural analyses. See text for detailed sample description.Scheidegger–Watson orientation-tensor formed fromthe individual linear segments of boundary tracesbetween the chosen phases. Resulting eigenvalues ratioR b is considered as a rough estimate of the degree ofGBPO.Results of the grain shape analysis and SPO analysisare shown in Fig. 10. Both amphibole (Fig. 10b) andplagioclase (Fig. 10a) show systematic changes. Whilesamples with preserved magmatic textures (T1) andLAC samples (LAC1, TG2) exhibit very weak SPOand lowest values of axial ratios (c. 1.5), samples ofmylonitic metagabbros show a significant increase inthe average axial ratio (2–3) and increase of the SPOdegree with decreasing temperature.Grain contact frequenciesGrain contact frequencies allow the statistical deviationfrom random spatial distribution of grainboundaries to be examined (Kretz, 1994). So far, thedegree of deviation of grain boundaries distributionfrom random distribution have been evaluated byplotting observed/expected ratio of like–like contactsof the two major minerals against each other. Here wepropose a diagram where v valuev ¼ Observed pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiExpected ð2ÞExpectedis plotted against the ratio of orientation-tensoreigenvalues, i.e. the degree of GBPO. The majoradvantage of this diagram is in simple visual evaluationof the degree of deviation from an expected randomdistribution of grain contacts (v ¼ 0).Results of this analysis are presented in Fig. 10. Theamphibolite facies mylonitic metagabbros exhibitalmost a random distribution at higher strains (GLT2)and tend to develop more of an aggregate-type ofmicrostructure (v values are positive for like–like andnegative for unlike contacts in Fig. 10c,d) at lowerstrains (GLT1) accompanied by a significant decreaseof preferred orientation of amphibole–amphibole andamphibole–plagioclase contacts. The high-grademylonitic metagabbros (GHT1 & GHT2) exhibit astrong aggregate distribution and a moderate preferredorientation of grain boundaries. Samples from thebanded amphibolite complex (TG2 & LAC1) show atendency towards a regular distribution (negative like–like v values and positive unlike v values) and a lowdegree of GBPO. The tonalitic gneiss weakly overprintedby the Variscan deformation (LAC3) showsÓ 2005 Blackwell Publishing Ltd243

662 O. LEXA ET AL.Fig. 9. (a) Plot of crystal size distribution (CSD) parameters N 0 and Gt. Parameters are obtained from CSD plots, where straightline is fitted on plotted data. Gt represents the slope and N 0 the intercept of fitted line. Shaded ellipses from upper left to lowerright corner mark domains of LT metagabbros, HT metagabbros, banded amphibolites and magmatic rocks. See discussion.(b) Examples of CSD plots used for N 0 and Gt parameter estimates.transitional values of v and R b between those of thebanded amphibolite complex and the mylonitic metagabbrosamples. The Variscan undeformed granodioritesill shows nearly a random distribution of graincontacts and an absence of GBPO.DISCUSSIONThe study samples of mafic lithologies from theCambro-Ordovician and Variscan structures exhibitremarkably similar mineral assemblages. In addition,the presence of garnet in metamorphosed mafic rocksindicates the medium pressure conditions for bothCambro-Ordovician and Variscan structures. Localoccurrence of orthopyroxene bearing mineral assemblagesin the banded amphibolites as well as in thewestern metagabbro mylonite belt (Table 1) indicatesthat both Cambro-Ordovician and Variscan metamorphicevents reached a boundary between granuliteand amphibolite facies conditions.The metamorphic assemblages and P–T estimates ofthe Cambro-Ordovician metamorphic conditions ofbanded amphibolite complex calculated overlap withinthe likely uncertainties with those of the Variscanmylonitic metagabbros and granodiorite. Moreover, inthe southern part of the region studied, the Variscanand Cambro-Ordovician fabrics are concordant andtherefore slight differences in metamorphic conditionsare indistinguishable. We suggest that in this case thestandard methods of metamorphic petrology cannotdiscriminate the two distinct metamorphic events fromeach other.Interpretation of crystal size distributionsThe grain-size distribution in metamorphic rocks(Cashman & Ferry, 1988; Eberl et al., 1998) is primarilycontrolled by the interaction between nucleationrate which is sensitive to temperature overstepping(Cahn, 1957), and the growth rate which is anapproximately linear function of overstepping (Ridley& Thompson, 1986). A further process that influencesCSD is textural coarsening driven by a tendency todecrease the excess of interfacial energy to reach alower energy state (Voorhees, 1992). The transfer ofmaterial from smaller grains to larger ones occurs bydiffusion and is commonly expressed by the Lifshitz–Slyozov–Wagner (LSW) equation (Lifshitz & Slyozov,1961) for Ostwald ripening, or by the communicatingneighbours theory (CN) by DeHoff (1991) where adiffusion length scale is spatially dependent. This differencebetween the theories is expressed in the waythat the CSD is modified during coarsening (Higgins,1998). In the case of LSW the descriptive parametersN 0 and Gt remain constant, while the CN theoryimplies a decrease in N 0 and Gt during coarsening.Experimental studies have shown that the size ofdynamically recrystallized grains is strongly dependenton recrystallization mechanisms that are controlled bystress and temperature (Twiss, 1977). Low-temperatureÓ 2005 Blackwell Publishing Ltd244

CONTRASTING TEXTURAL RECORD OF TWO METAMORPHIC EVENTS 663Fig. 10. (a, b) Plot of grains shape preferred orientation (SPO) of amphibole and plagioclase in studied samples. The results aresummarized by a boxplot-type plot of axial ratios v. eigenvalue ratios for individual phases. Individual boxes showing median, first andthird quartile values. The whiskers represent statistical estimate of range of data while outliers are not shown. Vertical axis characterizesa shape of grains, while horizontal axis represents area-weighted degree of preferred orientations. (c, d) Grain boundaryfrequencies. Plots of deviations from random spatial distribution v. degree of preferred orientation. The value of deviation fromrandom spatial distribution is obtained by contact-frequency method (Kretz, 1969) and length-weighted degree of preferred orientationis estimated as ratio of eigenvalues of bulk matrix of inertia. Plot of like–like and unlike boundaries is separated into two panels.recrystallization mechanisms such as bulging and subgrainrotation recrystallization (Poirier & Guillope,1979) are processes dominated by heterogeneousnucleation and grain-size reduction. On the contrary,grain-boundary migration recrystallization is oftenconnected with grain coarsening (Urai et al., 1986).Hickey & Bell (1996) suggested that the syntectonicgrowth of minerals is primarily controlled by gradientsin lattice strain energy, which results from inhomogeneousdeformation involving dislocation creep.Decreasing the strain rate to temperature ratio ( _e/T)leads to a decrease in the ratio of nucleation to growthrate (N/G), and coarser grain size can develops. On thecontrary, increase of _e/T ratio leads to increasing N/G,and therefore the grain size decreases. This hypothesishas been supported by Azpiroz & Ferna´ndez (2003)who used CSD plots to analyse dynamicallyrecrystallized metabasites across a metamorphic gradient.The CSD in the mylonitic metagabbros studied arecharacterized by a decrease in N 0 and an increase inGt values with increasing metamorphic temperatureestimates, and are projected in areas of high N 0values and low Gt in the N 0 –Gt plot. If all theseresults are interpreted in terms of CSD theory, theabove-described evolutionary trend indicates adecreasing dominance of nucleation processes overgrain growth with increasing temperature. The correlationof metamorphic temperature and grain-sizedistribution could be explained as a result of texturalcoarsening using the CN theory (Higgins, 1998).However, the microstructural observations of theÓ 2005 Blackwell Publishing Ltd245

664 O. LEXA ET AL.onset of dynamic recrystallization in weaklydeformed metagabbros show that the lower-temperaturemylonites exhibit a smaller initial recrystallizationgrain size than the new recrystallized grainsof the high-temperature mylonites (Baratoux et al.,2005). In addition, inspection of the N 0 –Gt plotshows that highly deformed samples (GLT2,GHT2a,b, T2) exhibit higher N 0 and lower Gt valuesrelative to weakly deformed samples (GLT1, GHT1).Based on these two observations, we suggest that thedifferences in CSD parameters are not the result oftextural coarsening but are merely controlled bytemperature- and strain rate-dependent mechanismsof dynamic recrystallization. Azpiroz & Ferna´ndez(2003) reported an increase in N 0 but a constant Gt(slope) with decreasing temperature and increasing Gtvalues for constant N 0 with increasing finite strain inthe Iberian Massif metabasites. However, it is proposedthat in our study the N 0 and Gt changesimultaneously due to temperature and strainintensity variations and that the temperature changeplays a key role in the resulting CSD shape.Samples from the banded amphibolite complex aremarked by lower N 0 and higher Gt values in comparisonwith the mylonitic metagabbros. The CSD isdeveloped in both amphibolites and tonalitic gneiss,which indicates that it is independent on a relativeproportion of amphibole and plagioclase in both rocktypes. These data in the N 0 –Gt diagram show a continuoustrend together with the above-described samplesfrom the mylonitic metagabbros. In addition,results of Hb–Pl thermometry reveal an increase inestimated temperatures from the eastern mylonitemetagabbros, through the western mylonite metagabbrosto the banded amphibolite complex. Such anevolutionary trend is likely to be interpreted as theresult of a textural coarsening comparable with theresults of Cashman & Ferry (1988) reinterpreted byHiggins (1998).Interpretation of spatial distribution minerals and grainboundariesSeng (1936) and later DeVore (1959) proposed that thespatial distribution of crystals in high-grade gneisses isdominantly determined by interfacial energy. Thisconcept was adopted by Flinn (1969) who explained aprevailing number of unlike boundaries in granulitesthrough the insertion of grains of one phase betweengrains of other phases. Flinn (1969) suggested that thisfeature is a consequence of a smaller interfacial energyof unlike boundaries in comparison with like–likeboundaries. However, Ramberg (1952) suggested thatdifferences in interfacial energies are too small to drivediffusional mass transfer in high-grade rocks.Modern material science experimental studies showthat during the wetting process the low-energy (lowmisorientationangle) boundaries in one phase arepreserved while another phase preferentially precipitateson high-energy (high-misorientation angle)boundaries (e.g. Kim & Rohrer, 2004). In other words,the highest energy boundaries are progressively eliminatedfrom an inherited population by ÔinfiltrationÕ ofthe other phase. This is in agreement with the knownfact that in granular-polygonal aggregates the minorphase precipitates on triple points to achieve lowertotal interfacial energy (Spry, 1969; Vernon, 1974).Such a tendency was documented by Dallain et al.(1999) who showed that the predominance of unlikecontacts in polycrystalline aggregates originatedthrough wetting of grain boundaries by fluids or melts,and subsequent precipitation of other phases on like–like contacts.In contrast, the solid-state differentiation resultingfrom dynamic recrystallization leads to the developmentof monomineralic aggregates or bands due tocoalescence of like phases at high strains (Schulmannet al., 1996; Kruse & Stu¨ nitz, 1999). Therefore, thelike–like contacts prevail and the so-called aggregatetypedistribution develops, which is a typical feature ofhigh-temperature deformation of polyphase rocks suchas gabbros and granites (Dallain et al., 1999; Baratouxet al., 2005).The high-temperature mylonitic metagabbros fromthe study area exhibit high-grain SPO and GPBOassociated with the development of a strong aggregatedistribution. The lower-temperature mylonitic metagabbrosare characterized by extreme values of SPO inconjunction with an almost random grain distribution.We suggest that in the case of high-temperaturemylonitic metagabbros the process controlling thedevelopment of a strong aggregate distribution is solidstatedifferentiation due to a different efficiency ofdislocation creep in hornblende and plagioclase. Thisprocess is likely to be accompanied by some diffusionalmass transfer responsible for preferential heterogeneousnucleation (Kruse & Stu¨ nitz, 1999) of interstitialplagioclase in coarse-grained amphibole aggregates(Baratoux et al., 2005). On the contrary, in the lowertemperaturemetagabbro mylonites, the random mineraldistribution and the lack of crystallographicpreferred orientation of plagioclase were interpreted tobe the result of mechanical mixing due to grainboundarysliding during granular flow.Samples from the banded amphibolite and tonaliticgneiss show a low SPO, a very weak elongation of bothplagioclase and amphibole, a weak GBPO connectedwith a weak dominance of unlike contacts indicating aregular to anticlustered grain distribution. Such a graindistribution and the large grain size of both phasesexclude mechanical mixing as a process explaining thistexture. We are of the opinion that the spatial distributionof plagioclase and amphibole (e.g. sampleLAC1) can result from heterogeneous nucleation ofplagioclase in an amphibole aggregates. However, anoriginal grain-size distribution characteristic of nucleationand growth processes is completely obliteratedby elimination of the small grains. Very low-grainÓ 2005 Blackwell Publishing Ltd246

CONTRASTING TEXTURAL RECORD OF TWO METAMORPHIC EVENTS 665elongation, SPO and GBPO, similar CSD of bothphases and regular grain distribution indicate that theaggregates tend to achieve a state with a minimuminterfacial energy. As mentioned above, the processesof heterogeneous nucleation of minor phase lead tooccupation of the high-energy grain boundaries atearly stages of the texture development. However, inthe rocks studied it is impossible to distinguish a minorphase from the host aggregate, which indicates that theprocess of interfacial energy reduction is more advanceddue to significant diffusional mass transfer. Asthe driving differences in bulk interfacial energies aretoo small (Ramberg, 1952), the only plausible explanationis the long time-scale of the process.Different time-scales of Cambro-Ordovician and Variscanmetamorphic eventsSˇtípska´ et al. (2001) proposed a tectonic model inwhich the Cambro-Ordovician metamorphism is relatedto large-scale rifting while the Variscan metamorphicevent was connected with a thermal effect inducedby the syntectonic granodiorite sill intrusion, and wasspatially restricted to the host rocks of the intrusion.The thermal time constant s is given by the relationship:s / l 2 j ;ð3Þwhere l is the characteristic length of thermal event andj is the thermal diffusivity (Carlslaw & Jaeger, 1959).This equation indicates that duration of thermalequilibration increases with the square of the size of thethermal anomaly. Based on the proposed tectonicmodel, we assume that a relaxation of the perturbedtemperature field generated by the continental riftdiffers in time-scale by at least one order of magnitudefrom that generated by an intrusion of several km insize. In other words, the metamorphic P–T conditionsattained at a similar depth level may yield similartemperatures but the time required for metamorphicequilibration and development of specific metamorphictextures may differ substantially.ACKNOWLEDGEMENTSThis research is a part of an ongoing collaborationbetween Charles University, Prague, Mainz University,Germany, and ETH Zu¨ rich, Switzerland. Financialsupport to P. Sˇtı´pska´ by the Charles University GrantAgency (grant no. 223/2002/B-GEO/PrF) and theCzech National Grant Agency (GA205/00/D043 andGA205/99/1195) is gratefully acknowledged. Themicroprobe work at ETH Zu¨ rich was financed by theSwiss National Foundation (ÔContinuous OrogenesisÕgrant to A.B. Thompson). Two visits of P. Sˇtı´pska´ toMainz University and a part of the microprobe workwere funded by the German Science Foundation(DFG, grant Kr 68-1, Kr 590/35). The grant no.24313005 of the Ministry of Education of the CzechRepublic provided funds for O. Lexa, P. Sˇtı´pska´ andK. 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Contrasting microstructures and deformation mechanismsin metagabbro mylonites contemporaneously deformed underdifferent temperatures (c. 650 ~ and c. 750 ~L. BARATOUX 1'2'5, K. SCHULMANN 3, S. ULRICH t'4 & O. LEXA t~Institute of Petrology and Structural Geology, Charles University,Albertov 6, 12843, Prague, Czech Republic(e-mail: Ika @natur.cuni.cz)2UMR 5570, ENS and Lyon 1 University, 2 rue RaphaYl Dubois,69622, Villeurbanne Cedex, France3Universitd Louis Pasteur, EOST, UMR 7517, 1 Rue Blessig,Strasbourg, 67084, France4Geophysical Institute, Czech Academy of Sciences, Bo(n{ 11/1401,14131 Praha 4, Czech Republic5Present address." Czech Geological Survey, Klarov 3, Praha 1,11821, Czech RepublicAbstract: Deformation mechanisms of amphibole and plagioclase were investigated in twometagabbroic sheets (the eastern and western metagabbros from the Stars M~sto belt,eastern Bohemian Massif), using petrology, quantitative microstructural and electron backscattereddiffraction methods. After the gabbroic pyroxene was replaced by amphibole, bothgabbroic bodies became progressively deformed. The eastern metagabbros were deformedunder temperature of c. 650 ~ and the western metagabbros under c. 750 ~ Subgrainrotation and dislocation creep, characterized by strong crystallographic and shape preferredorientations, operated in plagioclase of the eastern belt during the early stages of deformation.Subsequent randomizing of plagioclase crystallographic preferred orientation isinterpreted to be due to grain boundary sliding in the mylonitic stage. Large (50-150 ixm) grain sizes during the mylonitic stages are interpreted to be due to low strainrates. Amphibole is stronger and deforms cataclastically, leading to important grain sizereduction when the bulk rock strength drops substantially. In the western belt, plagioclasedeformed by dislocation creep accompanied by grain boundary migration (possibly chemicallyinduced) while heterogeneous nucleation and syndeformational grain growth inconjunction with dislocation creep were typical for amphiboles.Microstructural and rheological behaviour ofnatural polyphase rocks is a complex problem,which has been studied in detail mostly inquartzo-feldpathic rocks (e.g. Gapais 1989;Handy 1990). In these rocks, feldspars are generallyconsidered as strong minerals while quartzrepresents the weak phase. Handy (1994) proposeda scheme for rocks containing minerals of contrastingrheology with two end-members: load-beatingframework (LBF) and interconnected weak layers(IWL), based on the assumption that at least onemineral (generally the weaker one) is deformedby the mechanism of dislocation creep. Aspointed out by Jordan (1988) the LBF is notstable with increasing strain, resulting in mechanicallyinduced compositional layering. The strengthof a two-phase material depends on the proportionof the weak mineral and on the rheological contrastbetween the two phases (Jordan 1988; Handy1990, 1994). However, if the deformation mechanismswitches from crystal plasticity to some grainsize sensitive process, the strength of the bulk rockdrops rapidly with respect to the preceding modeldue to strain softening (Knipe 1989). It was alsoshown that generally stronger feldspars maybecome even weaker than quartz if they deformby diffusional creep (e.g. Voll 1976; Simpson1985; Gapais 1989; Martelat et al. 1999).Equivalent comparative microstructural studyof plagioclase and hornblende in naturally deformedmetabasic rocks as well as experimentalstudies of the rheology of amphibole-plagioclaseFrom: GAPA1S, D., BRUN, J. P. & COBBOLD, P. R. (eds) 2005. Deformation Mechanisms, Rheology andTectonics: from Minerals to the Lithosphere. Geological Society, London, Special Publications, 243, 97-125.0305-8719/05/$15.00 ~') The Geological Society of London 2005.249

98 L. BARATOUX ETAL.bearing rocks are rare (e.g. Hacker & Christie1990; Wilks & Carter 1990). However, metabasitesare considered to constitute a major partof the lower continental crust (Rutter & Brodie1992) and the study of deformation microstructuresand mechanisms is therefore crucial forunderstanding the deformation behaviour andrheology of mafic tectonites. Plagioclase rheologyis believed to control the strength of the lowercrust in several models (e.g. Carter & Tsenn1987; Ord& Hobbs 1989). Hornblende isassumed to be a relatively strong phase and tobehave passively during deformation unless itforms a load-supporting framework (Brodie &Rutter 1985).Many studies concerned with deformationmechanisms of plagioclase and hornblendehave been published to date. Experiments suggestthat plagioclase is deformed by dislocation creepunder lower crustal conditions (e.g. Tullis &Yund 1987; Ji & Mainprice 1990; Kruse et al.2001) with commonly active (010)[001] principalslip system (e.g. Olsen & Kohlstedt 1985;Kruhl 1987). Kruhl (1987) also observed a(001) [ 100] slip system in naturally deformed plagioclasesand Stiinitz et al. (2003) have shownthat slip on (001) and {111} in (110) directionare similarly active in experimentally deformedAn60 crystals. Besides these examples, otherless common slip systems were suggested byMarshall & McLaren (1977a, b), Montardi &Mainprice (1987), and Olsen & Kohlstedt(1984, 1985). In addition, grain size sensitiveprocesses in plagioclases have been referred toin some studies (Boullier & Gu6guen 1975;Lapworth et al. 2002). Hornblende generallyundergoes brittle deformation under low temperatureconditions (e.g. Brodie & Rutter 1985;Nyman et al. 1992; Lafrance & Vernon 1993).Many studies document crystal plastic deformationof hornblende with a dominant (100)[001] slip system from experiments and naturalrocks (e.g. Rooney et al. 1970; Cumbest et al.1989a, b; Hacker & Christie 1990). Volume diffusionrates are, however, slow in amphiboles(e.g. Freer 1981) implying that dislocationclimb is limited even at geological strain rates.Syndeformational chemical reactions betweenhornblende and plagioclase are common in metabasicrocks (Brodie 1981; Brodie & Rutter 1985)and involve deformation mechanisms such aschemically induced grain boundary migration(CIGM) (Cumbest et al. 1989a), on nucleationof new plagioclase (Rosenberg & Stfinitz 2003)or hornblende at plagioclase grain boundaries(Kruse & Sttinitz 1999).Metagabbros from the Star6 M6sto belt (easternmargin of the Bohemian Massif) representan example of a two-phase metabasic systemdeformed at different temperatures and straingradients. The aim of this study is to show twotypes of progressive evolution of deformationmicrostructures and textures of dynamicallyrecrystallized plagioclase-hornblende bearingmetagabbros at amphibolite and upper amphibolitefacies conditions with increasing bulk strain.Methods such as quantitative textural analysis,crystallographic preferred orientations (CPO)and study of mineral chemistry were used toconstrain deformation mechanisms for bothminerals. Finally, deformation mechanisms ofrheologically contrasting plagioclase and hornblendeare correlated and the mechanical behaviourof mafic rocks deformed under lowercrustal metamorphic conditions is discussed.Geological settingThe Star6 M6sto (SM) belt in the eastern marginof the Bohemian Massif separates the highgrade gneisses of thickened continental crust ofthe Lugian domain in the west from a Neo-Proterozoic continental margin in the east(Fig. l a). The SM domain was thinned duringCambro-Ordovician rifting and underwent granulitefacies metamorphism (Stfpskfi et al. 2001).Subsequent Variscan tectonics resulted in NE-SW trending structures dipping at relativelyhigh angles to the west (Figs 1 a and b).Strongly deformed 'western' metagabbros ofCambro-Ordovician protolith ages occur at thetop of the SM belt (Fig. 1) (Kr6ner et aL 2000),The metagabbros were pervasively affected bya ductile shear zone along which was emplacedsyntectonically a Carboniferous tonalite silldated at 340 Ma (Stfpskfi et aL 2004). A carboniferousmetamorphism of adjacent metagabbrosreached higher amphibolite facies conditionsbecause of the strong heat input from the tonalitesill. A leptyno-amphibolite complex of Cambro-Ordovician age comprising a sequence of alternatingamphibolites and tonalitic gneissesoccurs in the footwall of the tonalite sill. Thiscomplex is underlined by the 'eastern' metagabbrosheet, which is supposed to form part of thesame lower crust as the western (upper) metagabbrosheet (Stfpskfi et aL 2001). Amphibolitefacies Carboniferous metamorphic conditions ofthe eastern (lower) metagabbro are also attributedto the heat input of a more distant hangingwalltonalite intrusion.Rocks of Cambro-Ordovician age suffered twodeformations: D~ of Cambro-Ordovician age andD2 of Carboniferous age. The latter one prevailsin most lithologies of the SM belt, marked by apenetrative west-dipping $2 foliation bearing a250

TEXTURES OF NATURALLY DEFORMED METAGABBROS 99Fig. 1. (a) Location of the studied area in the frame of the European Variscides. Geological map of the Star6 M~stobelt is based on unpublished geological maps at 1:25 000 provided by courtesy of the Czech Geological Survey(Dr. M. Opletal, author). Important thrusts and faults as well as location of samples and a cross-section A-A' (Fig. lb)are indicated. (b) Schematic cross-section based on field observations shows major structures, lithology and majortectonic boundaries. (Vertical axis not to scale.)subhorizontal or gently inclined N-S trendingmineral lineation. The eastern metagabbros inthe hangingwall of the high-grade rocks of theSilesian domain were affected by localized D2shear zones that developed under amphibolitefacies conditions attained during the peak ofthe Carboniferous metamorphism. Identical geometryof structures and kinematics suggest thatthe eastern and western metagabbro belts weredeformed under a dextral transpressive shear251

100 L. BARATOUX ETAL.regime but under different thermal conditions(Stfpsk~i et al. 2001).Methods and techniquesMineral chemistryMinerals were analysed with a Cameca SX 100electron microprobe equipped with four WDSspectrometers at Blaize Pascal University,Clermont-Ferrand, France. Operating conditionswere 15kV, 10hA beam current, 2-5p~mbeam diameter, 20 s counting time, and naturalmineral standards. Some hornblende and plagioclasewere analysed using a CamScan $4scanning electron microscope and attachedLink EDX microanalytical system, at CharlesUniversity, Prague, Czech Republic. Operatingconditions were 20 kV, 1.8 nA beam current,1-3 p,m beam diameter, 120 s counting time,and mineral standards Structure Probe Instruments(SPI). Ca maps from plagioclase weremade at Claude Bernard University in Lyon,with operating conditions 15 kV, 15 nA beamcurrent and spatial resolution of 512 x 512pixels.Quantitative microstructural analysisQuantitative microstructural analysis of grainboundaries was carried out on traced and digitizedoutlines of grains in ESRI ArcView 3.2Desktop GIS environment. The map of grainboundaries was generated using ArcView extensionPoly (Lexa 2003). The resulting polygonshave been treated by MATLAB TM PolyLXToolbox (http://petrol.natur.cuni.cz/~ ondro;Lexa 2003), in which grain shape and grainboundary preferred orientations (SPO andGBPO, respectively) were analysed using themoments of inertia ellipse fitting and eigenanalysisof bulk orientation tensor techniques (Lexa,2003; modified SURFOR technique by Panozzo(1983) for GBPO). Their degree is expressed asthe eigenvalue ratio (r = E1/E2) of the weightedorientation tensor of grain shapes or boundaries.The orientation is defined by V~ and V2 eigenvectors.The grain sizes of the minerals werecalculated in terms of Ferret diameters of grainsection without stereological corrections.Crystallographic preferred orientationAmphibole and plagioclase crystallographic preferredorientations (CPO) were collected using ascanning electron microscope CamScan $4 inPrague and a JEOL JSM 5600 in Montpellierequipped with Channel5 electron backscatterdiffraction (EBSD) system from HKL Technology(Prior et al. 1999). Thin sections werepolished using 0.25 ~m diamond paste. Toremove all surface damage and minimize reliefbetween minerals, sections were chemicallypolished using a colloidal silica suspension. Allthin sections were carbon-coated. The coatingreduces the quality of the electron backscatterdiffraction patterns (EBSP) so that automaticindexing mode of the EBSP system could notbe used. Most data were therefore collectedmanually. Operating conditions were 20 kV inPrague and 15 kV in Montpellier, 5.6 nA beamcurrent, working distance 39 ram, and 2-5 ~mbeam diameter. For each measurement, threeEuler angles (vl, 05, v2) characterizing thelattice orientation as well as the nature of themineral were determined and stored. Practiceshows that plagioclase diffraction patterns donot change significantly from albite up to atleast An65 (Lapworth et al. 2002). Therefore,the Anorthite48 database was used for indexingplagioclase. Pole figures and inverse polefigures were projected using the software developedby D. Mainprice (ftp://ftp.dstu.univmontp2.fr/pub/TPHY/david/pc).Projectionsof crystallographic axes ([a], [b] and [c]) oroptical indicatrix (o~, /3, and 3/) are generallyused for plagioclase (e.g. Jensen & Starkey1985; Olsen & Kohlstedt 1985; Ji & Mainprice1988, 1990; Prior & Wheeler 1999).The degree of CPO has been quantified usingorientation tensor of crystallographic planesand directions (Mainprice, ftp://ftp.dstu.univmontp2.fr/pub.TPHY/david/pc).The intensityof CPO is given by the I parameter (Lisle 1985):3I = 15/2 x Z(Ei - 1/3) 2i=1where E~, E2, and E3 represent the eigenvalues ofthe preferred orientation of poles to planes anddirections of amphibole and plagioclase grainsplotted in pole figures. The values of I rangebetween 0 (no preferred orientation) and 5 (allfabric elements perfectly parallel one to another).MicrostructuresThe metagabbros from the eastern (lower) beltare affected by localized shear zones (Fig. lb)and all stages from non-deformed rock, protomylonite,mylonite, and up to highly strainedultramylonite are present (Figs 2a, b, c and d).In the western metagabbro (upper) belt, theprotolith stage is not preserved and only twodeformational stages can be distinguished252

TEXTURES OF NATURALLY DEFORMED METAGABBROS101Eastern belt (-650 ~Western belt (-750 ~Fig. 2. Macro photographs of typical metagabbro structures from the eastern (a-d) and the western (e, t3 belts:(a) non-deformed metagabbro (E0); (b) protomylonite (El); (c) mylonite (E2); (d) ultramylonite (E3); (e) augenmylonite (W1); (f) banded mylonite (W2).(Figs 2e and f): mylonites with hornblendeporphyroclasts related most probably to theCambro-Ordovician metamorphic event andcompletely recrystallized mylonites characterizedby a monomineral layering.Metagabbros of the eastern belt consist of plagioclase(40-60%), amphibole (40-60%), relictpyroxene in the cores of some large amphibolegrains (less than 1%), and titanite (less than1%). No substantial change in modal proportions253

102 L. BARATOUX ETAL.is observed with increasing deformation. Modalproportions of amphibole and plagioclase in thewestern metagabbros are more variable thanthose in the eastern belt. Amphibole proportionmay vary between 20 and 80%, most likely dueto original magmatic compositional variations.For the purpose of our study, samples composedof c. 50% of amphibole and c. 50% of plagioclasewere chosen.Deformation of the eastern (lower)metagabbro sheetNon-deformed metagabbro (EO). At low strain,plagioclase and hornblende exhibit euhedral randomlyoriented crystals of 0.5-3 mm in size.Tapering mechanical twins and local undulatoryextinction occur in plagioclase. There is no evidencefor any kind of dynamic recrystallizationor crystallization of new grains. Undulatoryextinction locally attests to some bending ofamphibole grains. Amphibole porphyroclastsshow random spatial distribution and they aregenerally not interconnected.Protomylonite (El). At higher strains, about20-25% of the total volume of plagioclasegrains but only 8-10% of amphibole showevidence of strain, suggesting that the deformationwas accommodated mostly by plagioclaserecrystallization and associated grain-sizereduction. Plagioclase grains show polysynthetictwins according to albite and pericline laws(Tullis 1983) and patchy undulatory extinction.Large plagioclase porphyroclasts of 2-5 mmare cut by brittle fractures (Fig. 3a) reducingthe grain size to 0.5-1 mm. The fractures arefilled with small twin-free recrystallized grainsof 0.02-0.1 mm. Two recrystallization mechanismshave been identified: bulging and subgrainrotation recrystallization (Fig. 3b), as proposedby Poirier & Guillop6 (1979) or Fitz Gerald &StiJnitz (1993), leading to core-mantle structures.Core and mantle structures were observed withan intermediate zone of subgrains and newgrains of similar size developed along porphyroclastboundaries. Boundaries between the neoblastsbecome progressively straight, meeting attriple junctions of 120 ~ .Large porphyroclasts of hornblende revealstrong internal deformation such as kinking,bending leading to sweeping undulatory extinctionand (100) twinning. The twin planes arenot regular and they locally form finger-likestructures. Brittle fractures transecting largegrains are often present. Randomly orientedporphyroclasts of 2-5mm in size rotateFig. 3. Drawing and micrographs (XPL) of the eastern metagabbro protomylonite. (a) Initial stage of plagioclasedeformation characterized by brittle fractures cross-cutting large porphyroclasts. Digitized drawing was used for thequantitative textural analysis. Arrows mark three clasts derived by fracturing of one porphyroclast and showcorresponding grains in the digitized drawing and micrograph. Plagioclase is white, hornblende is light grey, andopaque minerals are black. (b) New grains develop preferentially along fractures by a mechanism of subgrain rotation(SR). Note that the neoblasts are twin-free. (P) refers to porphyroclast.254

TEXTURES OF NATURALLY DEFORMED METAGABBROS 103progressively their cleavage planes { 110} parallelto the foliation. Small needle-like grains (0.02-0.05 x 0.1-0.3 mm in size), arranged parallelto the mylonitic foliation, are present in highstraindomains.Mylonite (E2). Within the eastern metagabbros,plagioclase porphyroclasts are recrystallized intoelongate aggregates or bands wrapped aroundhornblende clasts, and porphyroclasts representonly 5-10% of the total plagioclase volume.Elongate needle-like hornblende grains arearranged into aggregates or bands parallel to themain mylonitic foliation (Fig. 4a). Isolated hornblendeneedles scattered within plagioclase-richdomains are inclined at an angle of 20-30 ~ withrespect to the main fabric, indicating noncoaxialdeformation (Figs 4a and b). Plagioclasegrains attain 0.1-0.25 mm in size. Increase ingrain size with respect to the protomylonitestage and common optical zoning of recrystallizedgrains due to increase in anorthite contentsuggest syndeformational growth. Matrix grainsare often twin-free; the twinned grains representFig. 4. Digitized drawings and micrographs of the eastern mylonite. (a) Mylonitic foliation with alternating layers ofamphibole and plagioclase aggregates (PPL). (b) Plagioclase of subequant shapes and grain size typical for themylonitic stage (E2-pll). (c) High-strain plagioclase domain (E2-p12) adjacent to an amphibole porphyroclast ischaracterized by elongate shapes and strong SPO (XPL). (d) Amphibole porphyroclast (E2) cross-cut by microshearzones filled by small needle-like grains (XPL). Undulatory extinction documents strong internal deformation. Arrowsmark corresponding grains in the digitized drawings and micrographs.255

104 L. BARATOUX ETAL.only 10-20% of the total plagioclase volume.Plagioclases generally exhibit subequant shapeswith straight grain boundaries meeting at 120 ~triple joints (Fig 4b). However, in the vicinityof amphibole porphyroclasts, the plagioclasegrains are more elongate with higher aspectratios up to 2.5 (Fig. 4c).Hornblende grains are arranged in anastomosinginterconnected bands defining the myloniticfoliation. Most of the hornblende matrix grains,0.1-0.3 mm in length, are elongate in XZ sectionsattaining high aspect ratios (up to 6) whilethey are lozenge-shaped in the YZ sectionswith aspect ratios of 1.5-3. A total of 10-20%of the hornblende is present in a form oflocked-up porphyroclasts of 0.2-2 mm in sizeshowing undulatory or patchy extinction characteristicof strong internal strain. These porphyroclastsare elongate parallel to the foliation.Porphyroclasts are often transected by microshearzones (Fig. 4d) and small needle-shapedgrains consequently develop within these zonesby activation of the weak cleavage system{110}.Ultramylonite (E3). The ultramylonitic metagabbrosof the eastern sheet are marked by veryuniform mineral banding of fine-grained hornblendeand plagioclase layers ranging between 1 and5 mm in width (Fig. 5a). Asymmetric fabric ofhornblende and plagioclase is no longer present.Plagioclase forms continuous bands includingabundant interstitial grains of amphiboles(Fig. 5b). In contrast, amphibole-rich bands arealmost monomineral. Plagioclase is completelyrecrystallized, varying in size between 0.05 and0.3mm. The grains have smooth roundedshapes locally elongate parallel to the foliation.Rounded grains of 0.5 mm in size, interpretedas relicts of original porphyroclasts, locallyoccur in the fine-grained matrix. Plagioclasegrains are affected by strong retrograde sericitizationand albitization.Interstitial and matrix grains of hornblendereach 0.03-0.1 x 0.1-0.3 mm in size. Lockedupporphyroclasts of amphibole (0.5-1 mm insize) occur parallel to the foliation. However,they are less common than in the mylonite(E2). All amphibole grains are characterized byvery strong SPO (Fig. 5c).The mylonitic foliation is cross-cut by extensionalveins filled by a mixture of epidote andamphibole. These veins show that fluid activitywas high after the main mylonitic episode. Albitizationand sericitization of plagioclase in somesamples suggest that increased amount of waterwas present in these rocks also during late retrogression,which is most likely not related to themain process of mylonitization.Deformation of the western (upper)metagabbro sheetAugen mylonite (W1). In the western metagabbrobelt, the deformation was pervasive sothat initial protomylonite stages of deformationFig. 5. Drawings and micrograph of the eastern ultramylonite. (a) Plagioclase and amphibole banding (PPL).Plagioclase is darker than amphibole due to strong sericitization. (b) Plagioclase band in ultramylonite (E3)with abundant interstitial hornblende. Small grain size and subequant shapes are typical for these plagioclases.(c) Amphibole layer from ultramylonite (E3) is characterized by very fine grain size and strong SPO.256

TEXTURES OF NATURALLY DEFORMED METAGABBROS 105are absent. The least deformed rock representedby augen mylonite is marked by the presenceof rare 2-3 mm large porphyroclasts of plagioclase(Fig. 6a) and by 1-2ram wide and4-5 mm long porphyroclasts of hornblendewithin the recrystallized matrix (Fig. 6b). Thisrock type is preserved only in the southern partof the metagabbro belt. Plagioclase porphyroclastsrepresent only 5-10% of the total plagioclasevolume and show deformational twinsand undulatory extinction. Recrystallized plagioclasegrains show optical zoning.Modal proportion of amphibole porphyroclastscannot be estimated as they differ fromthe neoblasts only in their chemical composition,making them optically indiscernable. Amphiboleporphyroclasts exhibit patchy or sweeping undulatoryextinction but deformation twins have notbeen observed. The local occurrence of twins isassumed to be growth-related, the twin planesbeing always perfectly straight. Prismatic porphyroclastsattain higher aspect ratios thanmore rounded neoblasts. All rocks in thewestern belt are affected by amphibolite faciesre-equilibration partly destroying the granulitefacies peak metamorphic assemblages.Fig. 6. Drawings and micrographs of the western augenmylonite. (a) Digitized drawing of plagioclase (W1)used for the quantitative textural analysis. Largeporphyroclasts are surrounded by a mantle of finegrainedmatrix grains. (b) Hornblende (W 1) showsvariable grain size and strong SPO. Plagioclase is white,hornblende is light grey, and opaque minerals are black.Banded mylonite (W2). Distinct monomineralbands of 1-10 mm in width are typical of thesehighly deformed mylonites. Small hornblendegrains of irregular shape are common withinthe plagioclase layers, but plagioclase is rarelyincluded within amphibole bands. Both mineralsare completely recrystallized. Hornblende is insome places replaced by later prisrnatic cummingtonite,which grows either parallel orobliquely to the mylonitic foliation and islocally included within the plagioclase layers.Two types of plagioclase microstructuresoccur within one plagioclase band. The firsttype is characterized by grain sizes rangingfrom 0.3 to 1 mm in diameter, subequant orslightly elongate, but always irregular, grainshapes with serrated boundaries (Fig. 7a), andgrowth-related optical zoning. Mechanicaltwins cross-cut the optical zoning in somecases (Fig. 7b), suggesting that at least some ofthem formed later. The second type is characterizedby narrow (c. 0.5 mm) high-strain zonestrending parallel or slightly oblique to the layering(Fig. 7c). Here, the grain size is reduced to0.05-0.2 ram. Plagioclase grains attain higheraspect ratios (up to 2) and exhibit undulatoryextinction and subgrain boundaries elongate parallelto the high-strain zone. The contact betweenthese two deformational domains is either sharpor transitional.Hornblendes of 0.1 - 1 mm in size have straightgrain boundaries locally meeting at high anglessuggesting a high degree of textural equilibration(Vernon 1976) (Fig. 7d). No undulatory extinctionoccurs and twinning is very rare. Wherepresent, the twin planes are always very straight,indicating that twinning is growth-related ratherthan deformational. Aspect ratios vary between1.5 and 2.5, but some grains with aspect ratioclose to 1 occur. Grains with high aspect ratiosare not always parallel to the mylonitic foliation.Bulk rock chemistry, mineral chemistryand zoningBulk rock chemistryTable 1 presents bulk rock chemical analysesstudied by X-ray fluorescence at ClaudeBernard University, Lyon, France. The easternand western metagabbro sheets are regarded tobe a part of the same lower crustal unit prior totheir involvement into the Variscan tectonics(Stfpskfi et al. 2001). However, their chemistryis not identical and bulk rock analyses(Table 1) show that western metagabbros arecharacterized by higher contents in A1 and Caand lower contents in Na and Fe than their257

106 L. BARATOUX ET AL.Fig. 7. Drawings and micrographs of the western banded mylonite. (a) Lobate boundaries document grain boundarymigration (GBM) in coarse-grained plagioclases of banded mylonite (XPL). (b) Mechanical twins (MT) transect thegrowth-related zoning (GZ) in banded mylonites documenting a later deformation phase (XPL). (c) Digitized drawingof two plagioclase domains (W2-pll and W2-p12) used for quantitative textural analysis. (d) Hornblendes (W2) arecharacterized by straight grain boundaries meeting occasionally at triple points at variable angles. Many small grainsindicating high nucleation rate are present. SPO is rather low. Plagioclase is white, hornblende is light grey, pyroxeneis dark grey, and opaque minerals are black in all drawings.eastern counterpart. The increase in Na anddepletion in Ca in both metagabbro belts isrelated either to mylonitization or to primaryvariations due to postmagmatic processes.Trace elements show very similar trends forthe eastern and western metagabbros, suggestingthat they may originate from the same source.Metagabbros from both belts display a slightlynegative Nb anomaly normalized to MORB(Sun & McDonough 1989). Augen mylonitesfrom the western belt are depleted in Zr and Ticompared to eastern mylonites. Contents of Rb,Ba, and K are strongly variable in both belts,which is most probably related to postmagmaticalterations and/or mobility of these elementsduring metamorphism.Mineral chemistry and zoningSyndeformational chemical reactions arecommon in metabasic rocks (Brodie 1981;Brodie & Rutter 1985). Variations in plagioclasecomposition are shown in Figures 8 and 9.Amphibole compositions according to the classificationof Leake et al. (1997) are plotted inFigure 10. Representative analyses of mineralcompositions are listed in Tables 2 and 3.Mineral abbreviations are according to Kretz(1983).It is likely that the studied metagabbros experienceda metamorphic event prior to the maindeformation and metamorphism studied in thiswork. This is supported by existence of incompletelyamphibolized pyroxenes in undeformedrocks of low-strain domains. In addition, thecompositions of cores of large amphibole porphyroclasts(actinolites to actinolitic magnesiohornblendes)deviate from those of recrystallizedmagnesio-hornblendes in the eastern belt. Suchcompositions of primary amphiboles may indicateearly greenschist to amphibolite facies conditionspreceding the main higher gradeCarboniferous deformation.The peak metamorphic assemblages within theeastern metagabbro consist of P1 + Hbl + Qtz +Ttn _+ Ilm __%Mag. The composition of plagioclaseporphyroclasts in protomylonite (El)varies between Anso and An6o (Fig. 8a). Smallrecrystallized grains show a composition similarto the mother host grain corresponding toAnso-6o. More albitic compositions (An4o_45)occur along grain boundaries or triple junctionsof recrystallized grains (Fig. 9a). The geometryand sharp gradient in mineral zoning crosscuttingseveral grains is attributed to post-peak258

TEXTURES OF NATURALLY DEFORMED METAGABBROS107Table 1. Representative analyses of bulk rock compositionsEastern beltWestern beltProtolith (E0) Mylonite (El) Ultramyl. (E3) Augen myl. (Wl) Banded myl. (W2) Banded myl. (W2)SiO2 48.61 49.20 52.69 46.23 50.20 48.92TiO2 0.71 0.82 0.94 0.10 0.35 0.37A1203 17.76 19.74 18.04 21.57 21.43 24.20FeO T 6.14 6.32 7.33 4.86 4.10 3.84MnO 0.12 0.10 0.13 0.10 0.08 0.06MgO 9.31 6.56 5.81 7.91 6.20 4.76CaO 12.47 9.22 8.60 13.15 12.08 11.58Na20 2.64 3.29 4.99 0.82 2.98 3.55K20 0.16 1.52 0.10 1.60 0.51 0.25P205 0.04 0.03 0.03 0.01 0.02 0.07LOI 0.88 1.91 0.57 2.76 1.21 1.65H20- 0.06 0.12 0.05 0.19 0.09 0.10Total 98.90 98.83 99.28 99.30 99.25 99.35Ba 32.1 149.0 52.0 374.3 52.3 54.6Rb 1.6 55.1 4.9 74.7 26.9 11.9Sr 409.4 742.0 226.9 219.1 346.3 397.4Zr 33.0 34.7 39.3 7.8 10.7 29.6Nb 1.3 1.3 1.7 0.6 1.3 1.2Y 12.2 9.4 20.2 3.3 7.2 6.4V 185.9 160.1 203.9 115.0 131.9 92.0Cr 550.5 214.7 222.4 136.0 622.9 269.9Ni 122.8 60.6 58.8 48.9 87.3 90.9CO 36.8 35.4 29.2 32.0 26.6 25.4Sc 42.3 24.9 30.4 41.3 33.6 16.8Cu 8.0 3.6 19.1 2.9 20.4 33.3Pb 0.8 4.5 2.5 6.0 2.4 2.3Ga 13.8 15.0 12.2 12.8 16.9 17.1XMg* 0.730 0.649 0.586 0.744 0.729 0.688Major elements are given in wt%, trace elements are given in ppm. *XMg = Mg/(Fe 2+ + Mg).a) 12- Porphyroclasts Eastern belt b) Western beltPorphyroclaststrogressionII IN!I IIIII8 i . . . . i i i , i , i i , i i , i ig o I] I] , I ,I III~_ Recrystallized grains ~ protomylonite (El) Recrystallized grains6 I ~ mylonite(E2) I r~ ~ Augen mylonite (W1)1Late retrogression [Z] ultramylonite (E3)I---1 Banded rnylonite (W2)l2HI I, 111 t h10 20 30 40 50 60 70 80 90 10 20 30 40 50 60 70 80 90% An % AnFig. 8. Composition of plagioclase in (a) the eastern and (b) the western belts. Porphyroclasts are depicted separatelyfrom matrix grains. Deformation grade increases from dark grey to white colour.259

108 L. BARATOUX ET AL.Fig. 9. BSE images of plagioclase and amphiboleand representative chemical maps of plagioclasecompositional variations. (a) Plagioclase inprotomylonite from the eastern belt (El) suffers fluidrelatedchemical variations modifying the compositionat triple points and grain boundaries. (b) Growth-relatedchemical zoning in mylonite from the eastern belt (E2).(c) Growth-related zoning in coarse-grained plagioclaseof banded mylonite from the western belt (W2).fluid-phase activity along grain boundaries(McCaig & Knipe 1990). Recrystallized matrixplagioclases in the mylonite (E2) have andesiticcompositions (An3o-5o) and show increases inanorthite content towards their rims (Fig. 9b),which is commonly ascribed to increasingmetamorphic conditions. Rare rounded porphyroclast,as well as matrix grains in ultramylonites(E3), underwent strong and irregularly developedlate albitization. Matrix grains correspond toAn3o_43 although oligoclase and albite compositions(An6_ 15) are also common along fractures,cleavages and grain boundaries. These irregularcompositional variations are attributed to late retrogressionunrelated to the main deformationprocess.Amphibole porphyroclasts in protomylonites(El) are magnesio-hornblendes with variablecontent in Si but constant XMg (Fig. 10a).Amphiboles with more actinolitic compositionreplacing former magmatic pyroxenes can befound in the cores of some grains. Small amphibolegrains show similar compositions comparedto the porphyroclasts, some of them being moretschermakitic. The amphiboles in mylonites(E2) are magnesio-hornblendes with slightlylower Si contents than those in the protomylonites.High Si/low A1 domains documentingincomplete transformation of pyroxenes intohornblendes are present in the cores of large porphyroclastsas well. Locked-up grains in theultramylonites (E3) vary between magnesiohornblendeand tschermakitic composition.Small new grains have slightly lower A1 contentscompared to the less deformed stages.In the western belt, the peak metamorphicassemblages correspond to P1 § Hbl _ Cpx _+Opx +_ Grt ___ Ttn + Ilm. A granulite faciesmineral assemblage, comprising also sapphirineand corundum, is not stable and re-equilibratedduring subsequent upper amphibolite faciesreworking. In the south, granulite facies assemblagesare absent and probably completelyre-equilibrated. In the north, a late metamorphicphase is documented by the presence of prismaticorthoamphibole (gedrite).Plagioclase porphyroclast compositions arevery similar to those of the eastern belt and varybetween An4o and An65. Recrystallized matrixgrains in the augen mylonite (W1) show locallyincreasing anorthite content up to An6o_9o comparedto the porphyroclasts, indicating thatchemical reactions took place during recrystallization.Sharp limited domains of An4o_45 compositionalong grain boundaries are also common,suggesting late fluid activity (Knipe & McCaig1994). Recrystallized plagioclase grains inbanded mylonites (W2) show stronger variabilityin compositions (An35-9o). Growth-relatedzoning due to increasing metamorphic grade isdocumented in Figure 9c. The cores have morealbitic composition (An34) compared to the rims(An6o_65) and are attributed to the early metamorphicstage. Modifications of plagioclasecomposition characteristic of fluid-related compositionalchanges are also developed in thebanded mylonite.Amphibole porphyroclast compositions (Fig.10b) were analysed in two different thin sectionsof augen mylonite (W1) and the difference intheir mineral chemistry may be explained bylocal bulk rock chemical variations. The amphiboleporphyroclasts correspond to magnesiohornblendeswith similar compositions to thosein the eastern belt. Porphyroclast rim compositionsshow a decrease in Si content withrespect to the cores. Neoblasts are marked260

1.0--0.9__Mg~(Mg2++Fe2+)9 Core of porphyroctast9 Rim of porphyroclastb, Core of neoblasto Rim of neoblastTremoliteActinolileTEXTURES OF NATURALLY DEFORMED METAGABBROS 109I1,00.9TremoliteAclinoliteMagnesio-homblendefoo:oI Core of o oyroclast~Magnesio-hornblende ! TschermakiteCore of A i oporphyroclast [....Tschermakite9 9 A9 A,~ Ultramylonite (E3)i Protomylonite (El)I]-schermakiteMylonite (E2)irrotschermakite} . - - - -{ 5,75!iI~rrotschermakite5.75O.C8.00Ferro-actinolite7.50Ferro-hombtendeFerrotschermakite i6.50 5,751.0 i0.9iTremoliteActinoliteMagnesio-hornblende ~, ~.~ Tschermakiteo o ,.1~, ~ oo ~1.0I 0.9_Ng~ _(Mg2++Fe2+)[:008.004TremoliteActino}iteFerro*actinoliteMagnesio-hornblendeTschermakitePorphyroctast 1 ~ . . _ ~oO~ooPorphyroclast 2 ~ o o.... 0,9o-- -Augen mylonite (Wl)Ferro-homblendeFerrotschermakite7.50 )6.50 5.75~.50Bandedmylonite (W2)Ferrotschermakite5,75Fig. 10. Composition of amphiboles in (a) the eastern and (b) the western belts. (Classification of Leake et al. (1997)was used, (Na + K)A < 0.5.)by sharp compositional differences expressedby decreasing Si and XMg contents. In thebanded mylonite (W2), larger grains showhigher Si content than small neoblasts (verylow Si content), both being of tschermakiticcomposition.Temperature (T) was calculated using thehornblende- plagioclase thermometer byHolland & Blundy (1994; thermometer B).Estimated temperatures are based on 30 and 40amphibole-plagioclase couples in the easternand western belt, respectively. In the eastern(lower) sheet, calculated T was estimated tobe 650 4-50~ while T in the west was750 + 50 ~ Deformation is interpreted to havetaken place under or close to these metamorphicconditions. Pressure (P) could not be estimated asgarnet is absent in the mineral assemblage of themetagabbros. However, P of 9 kbar was estimatedat the lower contact between the metagabbrosheet and tonalitic sill, where garnetshave been formed (St/pskfi et al. 2001).261

110 L. BARATOUX ETAL.r~t'-,IMdMdddd~dd ~ d ~ d d d d d d ~r~

TEXTURES OF NATURALLY DEFORMED METAGABBROS 111e4Mo('4,,.0 ~ I'~ ~0 oo e4 oo ~1- P--. e~h ~"~ ',4~ ~ ',~ t"q ~0 ',,0 0", tt3 c4 0", ,--~9~-~.~~84-~dd dd,~,~d~ ~dd,~ddd4~dddII-F+d4-o.1...o0~8r~~ .~ . ~ .qq ~ . q ~ . ~ .~263

112 L. BARATOUX ET AL.Quantitative textural analysisGrain size distributionThe average grain sizes are presented as medianvalues of the Ferret diameter and the grain sizespread is expressed by the difference of thirdand first quartile (Q3- Ql). The statisticalvalues of the quantitative microstructural andgrain size analyses are summarized in Table 4.Plagioclase. Plagioclase in XZ section of theprotomylonitic stage in the eastern metagabbrobelt is characterized by an average grain size of45 ixm (El) (Fig. l la). Matrix grains in mylonitesare marked by an increase of the averagegrain size, reaching 74 pore in less (E2-pll) and58 Ixm in more (E2-p12) strained domains.Grain coarsening in the mylonite is consistentwith the syndeformational growth of grains originatedfrom subgrain rotation in the protomyloniticstage. This is supported by a greater amountof larger grains in grain size frequency histograms.The grain size is reduced to ~45 Ixm inthe ultramylonite (E3).Plagioclase grains in the augen mylonite ofthe western metagabbro belt (W1) are 43 Ixmon average, which is the statistical valueTable 4, Statistical values of the quantitative textural analysisEastern beltWestern beltProtomyl. Mylonite Mylonite Ultramyl. Augen myl. Banded myl. Banded myl.pl amp, pll p12 amp, pl amp, pl amp, pll p12GBPO Pl-pl 1.23 1.30 1.58 1.44 1.39 1.24 1.59Eigenvalue Amp-amp 2.50 3.53 2.26 1.96r = el/e2 Amp-plt 1.41 1.94 2.60 2.18 1.55 1.50Eigenvector V1 Pl-pl 7 24 23 17 3 - 3 5Orientation (~ Amp-amp 11 2 4 - 2Amp-pl * 16 12 18 7 -5 1SPO P1 1.45 1.46 1.73 1.85 1.86 1.38 1.64Eigenvalue Amp 1.83 4.40 2.89 2.21r = el/ea Amp t 1.81 2.67 3.43 3.26 1.96 1.76Eigenvector V1 P1 - 14 21 22 8 1 2 8Orientation (~ Amp 16 2 8 - 2Amp t 18 9 21 8 - 1 9Aspect ratio P1 1.59 1.56 1.82 1.84 1.68 1.50 1.72(median) Amp 4.46 3.79 2.36 2.12Amp ~ 2.67 2.59 3.02 3.00 2.28 1.84Grain size-Ferret diameterMedian (Ixm) P1 45 74 58 45 43 92Amp 49 41 119 94Amp t 49 39 43 33 59 50Q1 (~m) P1 33 52 43 30 30 57Amp 36 30 62 61Amp t 38 29 31 24 34 35Q3 (~m) P1 63 103 80 72 63 147Amp 68 62 211 142Amp t 64 52 60 44 102 74Q3 - QJ (~m) P1 30 51 37 42 32 90Amp 33 32 150 81Amp t 26 23 29 20 68 39Skewness "+ PI 1.44 - 0.34 - 0.08 - 0.09 1.15 0.01Amp 0.62 0.25 0.02 0.36Kurtosis * P1 7.30 2.86 2.77 2.53 6.99 2.98Amp 4.86 3.12 2.43 2.74R 2 P1 0.9151 0.9909 0.9979 0.9951 0.9442 0.9977correlation Amp 0.9768 0.9942 0.9899 0.9883coef.*Positive values are oriented anticlockwise with respect to the horizontal line.tin plagioclase domains.*The R 2 values, skewness close to zero, and kurtosis close to 3 signify well-fitted lognormal distribution.264725010656

' ~'-----2~--~----TEXTURES OF NATURALLY DEFORMED METAGABBROS 113b)Amphibole~.. ~ Amphibole in plagioclase domainsAmphibole,~ Amphibole in plagioclase domains\ A0 50 100 150Grain size - Ferret diameter (pm)9 Protomylonite- E1"~ A Mylonite- E2-pllLU z~ Mylonite- E2-pl2/x Ultramylonite- E39 ~ 9 Augen mylonite- Wlr o Banded mylonite- W2-pllo Banded mylonite- W2-p12Plagioclaseo50200 181614~8'Plagioclase ~ East ' "2Em Protomylonite - E1I. ~ Mylonit~- E2ttl [~ Ultramylonite- E36420fill L.. Plagioclase - West~21]il-grained plagioclasedlllllllllllllllilLlUL J~nn n~l][]nlIJ]n~Amphibole - Eastt 3~I ~ Mylonite- E2b9 11~ I I Ultramylonite- E3llilJLAmphibole - Westm Augen mylonite W1- [~ Banded mylonite- W250 100 150 200 250 3000 50 100 150 200 250 300 350 400Grain size - Ferret diameter (~m)Grain size - Ferret diameter (p.m)Fig. 11. Grain size evolution in the eastern (triangles) and western (circles) metagabbros. Degree of defomlationincreases from dark to white colour. (a) Calculated medians and standard deviations. (b) Histograms showing thecharacteristic frequencies of grain size for each deformational stage.representing the size of the recrystallized matrixgrains. Plagioclase grain size is substantiallyhigher in the coarse-grained domain (W2-pll)of the banded mylonite (92 ~zm). The grain sizeis reduced to 72 ~m on average in the highstraindomain (W2-p12).The grain size distributions (Fig. 1 lb) in protomylonitesof the eastern belt (El) show slightlybimodal distribution due to presence of largeporphyroclasts surrounded by small recrystallizedgrains. A higher amount of large grainscharacterizes the western banded mylonite(W2) compared to the augen mylonite (W1)and eastern mylonites (El-E3). Lognormal distributionscorrelate with intensity of deformationin both belts.Amphibole. The amphibole grain size is 49 p~min the eastern metagabbro mylonite (E2) and41 ~m in the ultramylonite (E3). However, thegrain size spread is always lower than that ofplagioclase in the same rock.The size of amphibole grains included withinplagioclase domains decreases systematicallywith increasing deformation, from 49 ~m in theleast deformed sample (El) to 39 ~m and43 ~m in the mylonite (E2-pll and E2-p12,respectively) and 33 p~m in the ultramylonite(E3). The grain sizes are always slightly lowerthan those of amphibole matrix grains in thesame thin section.Amphibole grain size in the western metagabbrobelt is highest in the sample including porphyroclasts(W1), reaching 119 Fm on average.Very high variance is typical of this mylonite,which is due to a relatively low number of newgrains and a high number of large porphyroclasts.In the banded mylonites (W2), grain sizedecreases to 94 Fm on average.In the eastern belt, amphibole grain size distributions(Fig. 11 b) are similar to those of plagioclase.They are characterized by well-developedlognormal distribution in both mylonite andultramylonite stages. In contrast, the distributionof amphibole grain size in the western belt exhibitsweaker fits of lognormal distribution withincreasing deformation, which is attributed to ahigher number of large grains of variable size(Table 4).Shape preferred orientationPIagioclase. Recrystallized plagioclase grainsin the protomylonite of the eastern belt (El)and mylonite (E2-pll) have aspect ratio (AR)of 1.59 and 1.56, respectively (Fig. 12a). Bothaspect ratio (AR = 1.82) and SPO increase inthe analysed high-strain plagioclase aggregateadjacent to the rigid amphibole porphyroclastin mylonite (E2-pl2). The ultramylonite (E3)exhibits further strengthening of the SPO andsimilar high aspect ratios (AR = 1.84).265

114 L. BARATOUX ETAL.o ._,

TEXTURES OF NATURALLY DEFORMED METAGABBROS 115maximum eigenvector V1 is only slightlyinclined with respect to the foliation, suggestingminor component of non-coaxial deformation. Inthe western belt, ramp_am p weakens from 2.26 inthe augen mylonite (W1) to 1.96 in the bandedmylonite (W2), being still stronger than rpl-plin similar rock types. The eigenvector V~ is parallelto the foliation in both types of mylonites.Crystallographic preferred orientationFive thin sections, covering all the deformationstages, were analysed using the EBSD technique.In the eastern mylonite (E2) and the westernbanded mylonite (W2), two plagioclase domainswere investigated.Besides pole figures of [100] directions andpoles to (010) and (001) planes, inverse polefigures, calculated parallel to stretching lineationand pole to foliation, are used in the eastern belt inorder to identify orientation maxima, which couldcorrespond to less common slip systems reportedin the literature. The inverse pole figures may beemployed because the orientations of lineationand foliation follow strong SPO and CPO ofamphibole. For the description of inverse polefigures, only the slip systems previouslydescribed in the literature will be mentioned(for summary see e.g. Kruse et al. 2001 ; Sttinitzet al. 2003). The I parameters representing theintensity of CPO are shown in Table 5.PlagioclaseWithin the protomylonite of the eastern belt (El),the plagioclase CPO is not random (Figs 13 and14a). In the inverse pole figure, the lineations aredistributed subparallel to the horizontal planewith a maximum around [i 10]. Poles to foliationare arranged in a N-S girdle with maximaaround (001) and a* (100). The (001) [1 i0] slipsystem (Olsen & Kohlstedt 1984, 1985) couldTable 5. Intensity I of crystallographic preferredorientationE1 1.51tE2-pl 1E2-p12E3 2.906Wl 1.422W2-pll 0.948W2-p12HornblendePlagioclase(100) [001] {110} [1001 (010) (001)1.926 0.897 0.116 0.666 0.3190.095 0.109 0.1000.108 0.347 0.1711.591 1.454 0.083 0.033 0.1781.854 0.846 0.111 0.074 0.0201.606 0.653 0.457 0.552 0.9480.590 0.299 0.741thus be possibly active. The orientation of porphyroclastsis different from that of recrystallizedgrains, suggesting that the recrystallization wasnot host-controlled (e.g. Ji & Mainprice 1990;Kruse et al. 2001). The mylonite (E2-pll andE2-p12) shows weak CPO in both plagioclasedomains. A weak maximum of lineations is situatedaround [110] in the E2-pll _domain, and someclusters of lineation around [1_12] directions andpoles to foliation close to (201) can be observedin the high-strain domain in sample E2-p12. Theseorie_ntationsmay indicate the possible activation of(201)1/21112] (Marshall & McLaren 1977a, b) inthe high-strain fine-grained domain adjacent to thehornblende porphyroclast (Fig. 4d). The ultramyloniteCPO (E3) displays maximum for polesto foliation clustering near the (021) and maximumfor lineations near the [112] direction. SuchCPO would suggest activation of (021)1/21112]slip system (Olsen & Kohlstedt 1984; Montardi& Mainprice 1987). The I parameters attain lowervalues in more deformed stages for [100], (010),and (001), respectively.The CPO of the augen mylonite in the westernbelt (W1) (Fig. 15) is very weak and no slipsystems could be extracted even by examiningthe inverse pole figures, which are not shownhere. Within banded mylonite (W2-pll and W2-p12), both coarse-grained (pll) and fine-grained(p12) plagioclase domains exhibit strong maximaof [100] directions close to the lineation. Polesto (010) and (001) planes are distributed close tothe pole of the foliation, which is defined by compositionallayering. The domain W2-p12 representslate shear zones marked by importantgrain refinement (Fig. 7c). The 1 parameters arehigher in the banded mylonite (W2) comparedto the augen mylonite (W1), indicating strengtheningof plagioclase CPO. It is likely that thislate shearing was not entirely coaxial with respectto the main flow represented by coarse-grainedaggregate fabrics (W2-pl 1). Therefore, the differenceof (010) and (001) maxima in pole figurescould result from slightly oblique orientation oflate shear plane with respect to dominant compositionallayering. The shear zone (W2-p12) ismarked by an inclination of [100] maximum ofabout 10 ~ with respect to the lineation, whichcould be attributed to non-coaxial deformation (Jiet al. 1988). Activity of slip systems (001)[100](Marshall & McLaren 1977a, b) and (010)[100](Montardi & Mainprice 1987) may be inferredfrom such orientation patterns.AmphiboleSmall recrystallized hornblendes adjacent totheir host grain in the eastern protomylonite267

116 L. BARATOUX ETAL.Plagioclase - eastern belt[1001 (010) (001).,7, ...... .. . . . . . ~,,. . :~':. .' ..~_s kQ...'........,....._~...:/r ....' .:.-,~.~ :- -.~ '...: ,'9 MD = 3.36 9 MD = 6.24 9 MD = 3.63.e-...........,, ~ .......-.-" :? ,::..,...9 MD = 3.28 9 MD = 2,83 9 MD = 2.54.,..,"E --9 MD = 3.33 9 MD = 3.99 9 MD = 3.33-.; ~&~-41 . --.,.."" :'. ".,. f" """~.-. ):: ..........., ,,~ ..Lower hemisphere 9 MD = 2.71 9 MD = 2.67 9 MD = 2.63Fig. 13. Point and contour pole figures of plagioclase CPO in the eastern belt (lower hemisphere, equal areaprojections). Diagrams are contom'ed as multiples (0.5, 1.0, 1.5, 2.0, 2.5 .... x ) of uniform distribution. MD ismaximum density of data in the contour diagrams. Foliation is represented by the horizontal line, lineation istrending E-W.(El) show progressive reorientation of the crystallographicplanes (100) towards the foliationdirection and c-axes [001] towards the lineation(Fig. t6a). Cleavage planes { 110} cluster parallelto the foliation as well. The asymmetry of theCPO fabric is consistent with the dextral shearzone cross-cutting the porphyroclast under theangle of ~20 ~ with respect to the mylonitic foliation.The CPO of hornblende in the ultramylonite(E3) displays a strong maximum of polesto the (100) and {110} planes (Fig. 16b) clusteringnormal to the foliation (I--2.9 for (100)).The c-axes trend in the direction of lineation.Such CPO patterns suggest that the (100)[001]slip system could be activated. Inclination ofdata with respect to the mylonitic foliation dueto dextral shear is present. Similar fabric asymmetry,interpreted as the result of a non-coaxialstrain, has been described in eclogite faciesglaucophanites (Zucali et al. 2002).Hornblende CPO in the augen mylonite (W 1)from the western belt form a cluster of (100)planes subparallel to the mylonitic foliationand c-axes parallel to the lineation. However, afew grains occur with a different orientation(Fig. 16c). In the banded mylonite (W2), the(100) planes and c-axes reveal similar orientationas in the augen mylonite but the maxima are lesspronounced (Fig. 16d). The I parameters areslightly lower in the banded mylonite (W2) comparedto the augen mylonite (W1). This CPOpattern is consistent with activation of the(100)[001] slip system. No asymmetry withrespect to the shear plane occurs.Discussion and conclusionsRecrystallization and deformationmechanisms of plagioclaseThe early deformation stage in plagioclase in theeastern amphibolite facies metagabbro belt ismarked by undulatory extinction, twinning, andfracturing, which are typical for low temperature268

TEXTURES OF NATURALLY DEFORMED METAGABBROS 117Lineationsa) upper hemisphere a-coo) a-(~oo)Protomylonite {.-'..:: "_~.'.~.~E1Plagioclase - eastern belt=219Poles to foliationsa*(100)a*(100)~ M D = 3.91 9D = 3.60, 9a*(100) a'(100) a*(100) a*(100)N=160 N:'.~-...:.Mylonite , .....E2-pl 1 ~!-'".:=o .;MyloniteE2-p12D=3,53 9D=2.839a*(lOO) a*(lO0) a*(lOO) a*(lOO)= 160, ", o-N=I ~ 0oUltramyloniteE3MD = 2,95 9D = 3.20 9a'(~00) a'(10o) a'0 00) a'(100)N = 200D=4, 9 D=4.b) ~ IPF Reference frame 12~~0Directions [uvw] _ Poles to planes (hkl) /o/o21O1 ~' 1~1~10,~~ O l O O l O ~ o i - oFig. 14. (a) Inverse pole figures of plagioclase CPO in the eastern belt (upper hemisphere, equal area projections).Projections of lineations are presented as point and contour diagrams in the left, projections of poles to foliations are inthe fight side of the figure. N is the number of plotted points. Diagrams are contoured as multiples (0.5, 1.0, 1.5, 2.0,2.5 .... • ) of uniform distribution. MD is maximum density of data in the contour diagrams. Large circle and square inthe protomylonite represent two porpbyroclasts. (b) The crystallographic reference frame.and/or high differential stress (Tullis & Yund1987, 1992; Rosenberg & Sttinitz 2003). Smallgrains concentrated in these fractures exhibittypical features of dynamic recrystallizationsuch as bulging and subgrain rotation. However,some of them might have developed bycrushing of the host grain or by heterogeneousnucleation (StiJnitz et al. 2003; Rosenberg &Sttinitz 2003). Twin-free grains adjacent totwinned host grains (see Fig. 3b), compositionalresemblance of recrystallized grains to hostgrains, and strong CPO (Lister et al. 1978) mayindicate a recrystallization mechanism by subgrainrotation. This recrystallization mechanismis also supported by small average grain sizes,small aspect ratios and weak SPO of newgrains. A fairly small degree of GBPO for plagioclase-plagioclaseboundaries is in agreementwith subequant shapes of equidimensional newgrains, thus favouring the hypothesis of subgrain269

118 L. BARATOUX ETAL.Plagioclase - western belt(010) (001)N= 1509 MD = 3.64 9 MD = 3.93 9 MD = 3.169 MD = 5.23 9 MD = 4.39 9 MD = 5.41Lower hemisphere 9 MD = 6.00 9 MD = 4.38 9 MD = 4.77Fig. 15. Point and contour pole figures of plagioclase CPO in the western belt (lower hemisphere, equal areaprojections). Diagrams are contoured as multiples (0.5, 1.0, 1.5, 2.0, 2.5 .... • ) of uniform distribution. MD ismaximum density of data in the contour diagrams. Foliation is represented by the horizontal line, lineation is trendingE-W.rotation recrystallization mechanisms. The CPOis strong and indicates a possible activity ofthe (001)[110] slip system.The mylonitic stage is characterized by subequantshapes of recrystallized plagioclasegrains (low aspect ratios) with straight boundaries,weak shape preferred orientation andslightly higher degree of GBPO compared tothe protomylonite. Recrystallized grains showincreasing grain sizes compared to protomyloniticstage. Compositional zoning in plagioclaseindicates syndeformational growth of matrixgrains (Sodre Borges & White 1980). The plagioclaseCPO in the mylonite is fairly weak. Theseobservations suggest that grain boundary sliding(Boullier & Gu6guen 1975; Lapworth et al.2002), accompanied by some component ofcrystal plastic deformation in plagioclase, wasthe dominant deformation mechanism. Mixingof amphibole and ptagioclase grains could beanother argument for the presence of grainboundary sliding. Rosenberg & Sttinitz (2003)suggested for the syntectonically cooled anddeformed Bergell tonalite that mixing of plagioclaseand biotite in a fine-grained matrix togetherwith weakening of CPO implied a mechanism ofdiffusion-accommodated grain boundary sliding.A switch of deformation mechanism from dislocationcreep to grain size sensitive (GSS) flowdue to strain softening and grain size reduction iswell known from quartzo-feldspathic rocks (e.g.Walker et al. 1990; Tullis & Yund 1991). Verysmall grain size is required for activation ofGSS deformation: 'less than 10 ~m' byBoullier & Gu~guen (1975), 2-16 &m by Tullis& Yund (1991), and 3-30 p~m by Stfinitz &Fitz Gerald (1993). However, some studiessuggest grain boundary sliding deformation forgrain sizes of 100-150 I-~m for plagioclase(Jensen & Starkey 1985), 24-41 I~m for mixedplagioclase-hornblende layers (Kruse & Stfinitz1999), and 100-250 p,m for plagioclase(Lapworth et al. 2002). In our samples withweak or random CPO, the grain size variesbetween 30 and 150 txm. The occurrence ofgrain-boundary diffusion creep for such grainsize may be explained by very low strain rates(Fliervoet et al. 1999; Lapworth et al. 2002).Fine-grained plagioclases from high-straindomains adjacent to amphibole porphyroclastsand matrix plagioclase grains in ultramylonitesshow a decrease in average grain size coupledwith increased aspect ratios, strong SPO andGBPO. The CPO is strong for the ultramylon_itesample, suggesting activation of a (021) 1/2 [ 112]slip system. These features may indicate thatat very high strains the grain size sensitiveflow becomes less important in plagioclase.270

TEXTURES OF NATURALLY DEFORMED METAGABBROS119[a](lOO)Hornblende[c][001]Lower [a]{110} hemisphere (100)[c][001] {110}a) Protomylonite - E 1 b) Ultramyionite - E3- a,~,~ -. .~- | 9 . .." 9N = 209N=989 MD = 7,69 9 MD = 9.57 9 MD = 4.619 MD = 15.98 9 MD = 12.08 9 MD = 5.55c) Augen mylonite - Wl d) Banded mylonite - W2N= 105 N= 107m MD = 7.93 9 MD = 9,94 =MD = 5.11 =MD = 5.15 = MD = 9,46 9 MD = 4.91Fig. 16. Point and contour pole figures of amphibole CPO in (a and b) the eastern belt and (c and d) the western belt(lower hemisphere, equal area projections). Contours are counted as multiples (1, 2, 3, 4 .... • ) of uniform distribution.MD is maximum density of data in the contour diagrams. Foliation is represented by the horizontal line, lineationis trending E-W. The CPO of porphyroclast is marked by a circle.A possible switch from diffusion-dominated flowin the mylonite to dislocation creep in the ultramylonitemay be due to a decrease in temperatureof deformation as described from felsic granulitesby Martelat et al. (1999). The switch fromthe GSS to the GSI regime for decreased grainsize is possible and implies an increase ofstrain rate and/or a decrease of temperature inthe shear zone (Handy 1990). A similar transitionfrom diffusion-dominated creep in plagioclaseaggregates in augen mylonites to dislocationcreep in banded ultramylonites was reportedby Schulmann et al. (1996) from deformedmetagranitoids.Dominant (010)[001] slip was identified inplagioclase from both CPO patterns and/orTEM observations (e.g. Olsen & Kohlstedt1984, 1985; Montardi & Mainprice 1987; Ji &Mainprice 1988, 1990). However, the observedCPO patterns for amphibolite facies metagabbrosare not compatible with activation of the classicslip system. Various other slip systems, whichmay be activated in plagioclase also, weredescribed by Marshall & McLaren (1977a, b),Olsen & Kohlstedt (1984, 1985), and Montardi& Mainprice (1987). In addition, St(initz et al.(2003) concluded that during experimentaldeformation of plagioclase several slip systems271

120 L. BARATOUX ET AL.are contemporaneously activated. The presenceof several maxima in inverse pole figures ofour samples may indicate that several slipsystems probably operated simultaneously andtheir unequivocal identification from CPOpatterns is therefore difficult.In the western augen mylonites, plagioclaserecrystallizes dynamically by subgrain rotation.Aggregates of plagioclase matrix grains showmicrostrnctures indicative for dislocation creep,such as subgrain formation, strong SPO, highaspect ratios and relatively strong GBPO ofplagioclase boundaries. Large differences inchemical compositions between porphyroclastand matrix may imply that some of the grains originatedby heterogeneous nucleation (Rosenberg& Sttinitz 2003). Surprisingly, the CPO ofrecrystallized grains is very weak, which maysuggest that some GSS process took place atthis stage of deformation.In the banded mylonites, the plagioclasematrix grains show strongly migrated boundaries,and lower SPO, aspect ratios and GBPOcompared to the augen mylonite. Plagioclase ischaracterized by a relatively high average grainsize and a common growth-related, chemicallyzoned composition. The plagioclase cores areovergrown by a new plagioclase with a morecalcic composition at upper amphibolite to granulitefacies conditions. A similar mechanism wasnumerically modelled by Jessell et al. (2003)showing that some original grain cores aredirectly in contact with rims of new grains dueto preferential growth of grains at the expenseof others. This sharp compositional differencemust necessarily introduce a strong chemicalpotential between these two grains. Serratedboundaries indicate a mechanism of grain boundarymigration. A process called chemicallyinduced grain boundary migration (CIGM) wasintroduced by Hillert & Purdy (1978) and inthe case of naturally deformed amphibolesdescribed by Cumbest et al. (1989a). Such aprocess requires a strong chemical potential andis driven by a reduction in the free chemicalenergy associated with the chemical change(Yund & Tullis 1991). However, Yund &Tullis (1991) observed important compositionalchanges in experimentally deformed plagioclase,which were not assumed to be the driving forcefor associated dynamic recrystallization. TheCPO in banded mylonites is relatively strongand (001)[100] and (010)[100] slip may beinferred from the pole figures. Martelat et al.(1999) reported a strong CPO in plagioclasedeformed by high-temperature diffusion creepin coarse-grained felsic granulites. We propose,however, that a difference in chemical potentialcould have been one of the main driving forcesfor grain boundary migration in the bandedmylonites.Fine-grained shear zones, cross-cutting thecoarse-grained plagioclase bands, developedlater at somewhat lower temperatures. Thisis supported by the mechanical twins crosscuttingchemically zoned plagioclases fromthese shear zones. Higher aspect ratio, strongerSPO and GBPO, and strong CPO suggest dominantactivity of dislocation creep with activationof (001)[100] slip in these domains.Deformation mechanisms of amphiboleMicrostructural evidence for cataclastic flow waspresented by Nyman et al. (1992) from amphibolitesthat were previously interpreted to havebeen deformed plastically. Amphiboles are relativelyresistant to crystal plastic deformation evenat high metamorphic grades (Brodie & Rutter1985). The dominant slip system (100)[001] inamphibole was inferred from both experimentaland natural studies (Dollinger & Blacic 1975;Rooney et al. 1975; Cumbest et al. 1989b;Hacker & Christie 1990; Skrotzki 1990). The(hk0)[001] slip system was determined byBiermann & Van Roermund (1983) andMorrison-Smith (1976). The (010)[100] and(001)[100] slip systems were observed byMorrison-Smith (1976) in experimental studiesbut not in nature. All these studies concludedthat the only possible active Burgers vectoroperating in amphiboles WaSoin [001], which isalso the shortest one (5.299 A). [100] and [010]Burgers vectors (9.885 A and 18.169 A, respectively)were thought to be too long to be activated(Rooney et al. 1970).Hornblende microstructures from the easternmetagabbro belt (protomylonite and mylonite)exhibit twinning on (100) planes (Rooney et al.1975), kinking, and strong undulatory extinctionrelated to bending/twins of the crystalline lattice.Microshear zones transect large porphyroclastsinto bookshelf-like segments and these segmentssubsequently rotate with their c-axes parallel tothe shear direction. This is consistent with polefigures showing CPOs of small grains that aredifferent from the host porphyroclast. In YZ sections,the grains have lozenge shapes, with theirgrain boundaries parallel to the {110} cleavage.The SPO and GBPO are relatively weak forthese strongly elongate grains when comparedto grains from more advanced stages of deformation.This is interpreted to be the result ofdevelopment of new grains along intracrystallinemicroshear zones and along margins ofporphyroclasts.272

TEXTURES OF NATURALLY DEFORMED METAGABBROS 121Some observations are consistent with brittledeformation of amphibole. For instance, hornblendehost grains are often rimmed by 'subgrains'forming core-mantle-like structures.Grain boundaries between new grains and porphyroclastsare very sharp, indicating that newgrains formed by fracturing rather than by subgrainrotation (Nyman et al. 1992). Interstitialplagioclase decorate grain boundaries betweensmall recrystallized grains demonstrating openingof spaces between amphiboles. However, theoperating deformation mechanism in hornblendecannot be unequivocally determined fromoptical observations and EBSD data alone. Thestrong CPO patterns may be interpreted by(100)[001] slip with dislocation creep being themajor operative deformation mechanism, bytranslation gliding along {110} planes indicatingdominant microfracturing or by mechanicalrotation of anisotropic grains. The strong asymmetryof CPO reflects the development ofantithetic microshear zones operating duringbookshelf-like rotations associated with dextralshear.Hornblende crystals in ultramylonites arecharacterized by higher aspect ratios, moreuniform grain sizes, and stronger SPO andGBPOs compared to amphiboles from protomylonites and mylonites. Strong CPO of (100)and { 110} cleavage planes subparallel to the foliationand c-axes subparallel to the lineation areconsistent with activation of the (100)[001 ] and{110}[001] slip systems. As mentioned above, astrong CPO of amphibole may also be producedby rigid body rotation.The most important observation in the westernaugen mylonitic hornblendes is that the originalporphyroclasts and new grains have differentchemical compositions (magnesio-hornblendeand tschermakite with higher Na and K content,respectively). Chemical differences between oldand new grains suggest that new grains originatedby heterogeneous nucleation under highermetamorphic conditions (e.g. Binns 1965; Laird& Albee 1981), which may be chemically and/or deformationally induced (Berger & Sttinitz1996). The porphyroclasts show high aspectratios and SPO, GBPO, and CPO maximaoriented parallel to the mylonitic foliation. Arigid body rotation of large porphyroclasts inthe weaker plagioclase matrix is the most likelymechanism to explain the strong preferred orientationof these grains. New grains have significantlylower aspect ratios, strong SPO, CPOand high degrees of GBPO of like-like boundaries.The orientation of CPO maxima is consistentwith the possible activation of (100)[001]slip.In the banded mylonites, primary hornblendeporphyroclasts are missing; however, slightchemical zoning within large and smaller newgrains indicates their syndeformational growth.Amphiboles, in monomineral hornblende bands,show straight boundaries decorated with smallamounts of tiny interstitial plagioclase. Amphibolegrains show no signs of internal deformationand have straight to lobate grain boundaries,which is indicative of high degrees of texturalequilibration (Brodie & Rutter 1985). The grainsizes and the degree of SPO and GBPO decreasewith respect to less deformed rocks (augen mylonite).However, the CPO is strong and verysimilar to that from less deformed augen mylonite,indicating possible activation of (100) [001 ]slip. Hornblende grains, which occur in plagioclaseaggregates, are marked by significantlysmaller grain sizes compared to host matrix crystals(Fig. 7c). They are often located at highenergysurfaces or unstable triple points, andshow globular shapes typical for a very differentstructure of host and inclusion grain. In someplaces, where the minor phase becomes abundant,it is forced to occupy less favourableplanar interfaces (Vernon 1976). All these criteriasuggest that these grains developed byheterogeneous nucleation, namely in a highsurface energy plagioclase aggregate (Spry1969; Dallain et al. 1999). The mechanism ofsyndeformational growth is more pronouncedin hornblende layers. This is supported byshapes of grain size distribution histograms thatshow important numbers of large grains. Possiblemechanisms to explain strong CPO in the studiedrocks is a nucleation and growth of new hornblendegrains in a non-lithostatic stress field aswas proposed for pyroxenes by Helmstaedtet al. (1972) or Mauler et al. (2001).Implications for the theology of atwo-phase amphibole-plagioclase systemExperimental deformation of amphibolite andgranulite facies natural samples of metabasites(Wilks & Carter 1990) reveals that the rheologyof such polyphase systems depends on temperature,rock composition, deformation mechanisms,water content, and to a minor extent,pressure. The relationship between these factorsand differential stress and strain is, however,complex and Wilks & Carter (1990) concludedthat it is necessary to estimate the contributionof each phase to the creep rate of the bulk rockin order to establish a convenient flow law.Rosenberg & Sttinitz (2003) proposed thatinterconnected monomineralic networks of273

122 L. BARATOUX ET AL.recrystallized plagioclase in middle or deepcrustal metabasic rocks are rare and restrictedto rocks with very high plagioclase modal abundancessuch as anorthosites (Ji et al. 1988).However, as mentioned by other authors (e.g.Brodie & Rutter 1985), banded metabasic mylonitesare common, which is confirmed by thiswork. We show that plagioclase may easilyform interconnected weak layer networks inhornblende gabbros under upper amphibolitefacies conditions.Our microstructural study of an amphibolitefacies metagabbro (650 _+ 50 ~ shows thatthe load-bearing framework structure (Handy1990, 1994) is restricted to the lowest deformationintensities. The non-deformed metagabbroshows coarse-grained ophitic structure composedof randomly distributed hornblende and plagioclase,where amphibole grains are only locallyin contact. With ongoing strain, the deformationis mostly concentrated in the plagioclase, whichis at the transient region between brittle andplastic behaviour, leading to development of afine-grained matrix (Tullis & Yund 1987).Amphibole grains showing high internal strainand local fracturing behave as rigid bodies surroundedby interconnected layers of plagioclasegrains. Such microstructures may be interpretedas an interconnected weak layer structure(IWL) with a high viscosity contrast betweenrigid clasts of amphibole and weaker, finegrainedplagioclase layers (Handy et al. 1999).In the amphibolite facies metagabbroic mylohires,the monomineralic hornblende layers areobserved, while plagioclase-rich layers showalmost perfect mixing with hornblende. Thegrain size of plagioclase and amphibole in theplagioclase-rich matrix areas is fairly similar,the latter showing slightly higher elongation. Thephase distribution of plagioclase-hornblendemixture, the absence of CPO in plagioclase,and its aspect ratio suggest that the dominantmechanism is granular flow (grain boundarysliding). Based on this interpretation, wesuggest that the phase mixing is probably amechanical process. The microstructures andCPO of amphibole forming monomineraliclayers indicate either dislocation creep or cataclasticflow. The absence of boudinage and progressivemixing of plagioclase and amphibolesuggest that the diffusion-dominated flowprocess operating in plagioclase aggregates ismechanically as efficient as dislocation or cataclasticflow in the hornblende layers. The finalstructure resembles the interconnected weaklayer structure with low viscosity contrast(Handy 1994). A switch in deformation mechanismfrom dislocation creep towards a grain sizesensitive process is thought to be responsiblefor the convergence of mechanical properties ofamphibole and plagioclase in the mylonite,resulting in a drop of bulk rock strength(Etheridge & Wilkie 1979; Kirby 1985; Rutter &Brodie 1988).Two deformation stages were observed in theupper amphibolite facies (750 -t- 50 ~ metagabbros:(1) augen mylonite with locally preservedporphyroclasts of both plagioclase and hornblende;and (2) banded mylonites with completelyrecrystallized amphibole and plagioclase,each arranged in monomineralic layers. Theinitial stages of deformation are characterizedby tectonic grain size reduction of plagioclasewhile hornblendes represent strong objects floatingin the weak plagioclase matrix. The deformationof the metagabbro is interpreted to haveoccurred via dislocation creep accompanied bydiffusion mass transfer mechanisms, responsiblefor moderate mixing of plagioclase and amphibole.The banded fabric, only developed at highstrains, can be defined as an interconnectedweak layer structure with low viscosity contrast(Handy et al. 1999). The layered structureshows that the strengths of amphibole monornineralicaggregates and plagioclase-rich bandsare similar, suggesting convergence of rheologiesof both minerals at high strains (Jordan1988; Handy 1994).In conclusion, amphibolite and upper amphibolitefacies metagabbroic mylonites are characterizedby layered low-viscosity IWL structures.This indicates that at high strains this bandedstructure in metagabbros forms a so-calledsteady-state foliation (Means 1990). Mechanicalmixing of phases is more important in lowertemperature (eastern belt) than in higher temperature(western belt) amphibolite facies mylonites.The bulk strength of amphibolite andupper amphibolite facies mylonitic metagabbrosis controlled by an equal contribution of bothrock-forming minerals showing contrasting butequally efficient deformation mechanisms.We are indebted to D. Mainprice for his help with theEBSD analyses. Fruitful discussions with F. Holub, D. J.Prior, H. Stiinitz and J. Wheeler are gratefully acknowledged.We thank P. T~)cov~, J. Haloda, P. Grandjean andP. Capiez for the help with microprobe and bulk rock analyses.K. Brodie, H. Van Roermund and D. Gapais arethanked for thorough reviews, which improved significantlythe original manuscript. The project was fundedby grants of Czech National Grant Agency No. 42-201-204 to K.S. and 42-201-318 to P. Stipskfi, by Czech GeologicalService assignment No. 6327 to P. Mixa, and by aPhD financial support attributed by the French governmentto L.B.274

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TEXTURES OF NATURALLY DEFORMED METAGABBROS 125ROONEY, T. P., RIECKER, R. E. & Ross, M. 1970.Deformation twins in hornblende. Science, 169,173-175.ROONEY, T. P., RIECKER, R. E. & GAVASCI, A. T.1975. Hornblende deformation features. Geology,3, 364-366.ROSENBERG, C. L. & STONITZ, H. 2003. Deformationand recrystallization of plagioclase along atemperature gradient: An example from theBergell tonalite. Journal of Structural Geology,25, 389-408.RUTTER, E. H. & BRODIE, K. H. 1988. Experimental'syntectonic' dehydration of serpentinite underconditions of controlled pore water pressure.Journal of Geophysical Research, B, 93, 4907-4932.RUTTER, E. H. 8,: BRODIE, K. H. 1992. Rheology of thelower crust. In: FOUNTAIN, D. M., ARCULUS, R. &KAY, R. W. (eds) Continental Lower Crust. Developmentsin Geotectonics, Elsevier, Amsterdam, 23,201-267.SCHULMANN, K., ML(~OCH, B. 8,: MELKA, R. 1996.High-temperature microstructures and rheology ofdeformed granite, Erzgebirge, Bohemian Massif.Journal of Structural Geology, 18, 719-733.SODRE-BORGES, P. & WHITE, S. H. 1980. Microstructuraland chemical studies of sheared anorthasites,Roneval, South Harris. Journal of StructuralGeology, 2, 273-280.SIMPSON, C. 1985. Deformation of granitic rocksacross the brittle ductile transition. Journal ofStructural Geology, 7, 503-511.SKROTZKI, W. 1990. Microstructure in hornblende of amylonitic amphibolite. In: Knipe ROBERT, J. &RUTTER, E. H. (eds) Deformation Mechanisms,Rheology and Tectonics. Geological Society,London, Special Publications, 54, 321-325.SPRY, A. 1969. Metamorphic Textures. PergamonPress, Oxford.STiPSKA, P., SCHULMANN, K., THOMPSON, A. B.,JE~EK, J. & KRONER, A. 2001. Thermo-mechanicalrole of a Cambro-Ordovician paleorift during theVariscan collision; the NE margin of the BohemianMassif. Tectonophysics, 332, 239-253.STfPSKA, P., SCHULMANN, K. & KRONER, A. 2004.Vertical extrusion and middle crustal spreading ofomphacite granulite: a model of synconvergentexhumation (Bohemian Massif, Czech Republic).Journal of Metamorphic Geology, 22, 179-198.ST15NITZ, H. & F/TZ GERALD, J. D. 1993. Deformationof granitoids at low metamorphic grade; II, Granularflow in albite-rich mylonites. Tectonophysics,221, 299-324.STUNITZ, H. FITZ GERALD, J. D. & TULLIS, J. 2003.Dislocation generation, slip systems, and dynamicrecrystallization in experimentally deformedplagioclase single crystals. Tectonophysics, 372,215-233.SUN, S. S. & MCDONOUGH, W. F. 1989. Chemical andisotopic systematics of oceanic basalts; implicationsfor mantle composition and processes. In:SAUNDERS, A. D. & NORRY, M. J. (eds) Magmatismin the Ocean Basins. Geological Society,London, Special Publications, 42, 313-345.TULLIS, J. 1983. Deformation of feldspars. In:RmBE, P. H. (ed) Feldspar Mineralogy. Reviewsin Mineralogy. Mineralogical Society of America,Washington, DC, 2, 297-323.TULLIS, J. & YUND, R. A. 1987. Transition fromcataclastic flow to dislocation creep of feldspar;mechanisms and microstructures. Geology, 15,606-609.TULLIS, J. & YUND, R. A. 1991. Diffusion creepin feldspar aggregates; experimental evidence.Journal of Structural Geology, 13, 987-1000.TULLIS, J. & YUND, R. 1992. The brittle-ductile transitionin feldspar aggregates; an experimentalstudy. In: EVANS, B. & WONG TENG, F. (eds)Fault Mechanics and Transport Properties ofRocks; A Festschrift in Honor of W. F. Brace. AcademicPress, San Diego, United States, 89-117.VERNON, R. H. 1976. Metamorphic Processes. Murby,London, Wiley, New York.VOLL, G. 1976. Recrystallization of quartz, biotite andfeldspars from Erstfeld to the Leventina Nappe,Swiss Alps, and its geological significance. SchweizerischeMineralogische und PetrographischeMitteilungen, 56, 641-647.WALKER, A. N., RUTTER, E. H. & BRODIE, K. H. 1990.Experimental study of grain-size sensitive flow ofsynthetic, hot-pressed calcite rocks. In: KNIPE, R. J.& RUTTER, E. H. (eds) Deformation Mechanisms,Rheology and Tectonics. Geological Society,London, Special Publications, 54, 259-284.WILKS, K. R. • CARTER, N. L. 1990. Rheology ofsome continental lower crustal rocks. Tectonophysics,182, 57-77.YUND, R. & TULLIS, J. 1991. Compositional changesof minerals associated with dynamic recrystallization.Contributions to Mineralogy and Petrology,108, 346-355.ZUCALI, M., CHATEIGNER, D., DUGNANI, M.,LETTEROTTI, L. & OULADDIAF, B. 2002. Quantitativetexture analysis of glaucophanite deformedunder eclogite facies conditions (Sesia-Lanzozone, Western Alps): comparison between X-rayand neutron diffraction analysis. In: DE MEER, S.,DRURY, M. R., DE BRESSER, J. H. P. &PENNOCK, G. M. (eds) Deformation Mechanisms,Rheology and Tectonics: Current Status andFuture Perspectives. Geological Society, London,Special Publications, 200, 219-238.277

ClickHereforFullArticleJOURNAL OF GEOPHYSICAL RESEARCH, VOL. 112, B10210, doi:10.1029/2006JB004820, 2007Extreme ductility of feldspar aggregates—Melt-enhanced grainboundary sliding and creep failure: Rheological implications forfelsic lower crustProkop Závada, 1,4 Karel Schulmann, 2 Jiří Konopásek, 3 Stanislav Ulrich, 1,4and Ondrej Lexa 1,2Received 25 October 2006; revised 17 July 2007; accepted 13 August 2007; published 27 October 2007.[1] High-grade orthogneisses from granulite-bearing lower crustal unit show extremefinite strains of both K-feldspar and plagioclase with respect to weakly deformedquartz aggregates. K-feldspar aggregate in the most intensely deformed sampleshows interstitial grains of quartz and albite, which also mark some intragranular fractureswithin K-feldspar grains. Both interstitial grains and fractures are oriented mostlyperpendicular to the sample stretching lineation. Quartz and albite grains withinK-feldspar bands are interpreted as crystallized from interstitial melt and the petrologystudy shows that the melt was produced by a metamorphic reaction inplagioclase-mica bands. Thermodynamic Perple_X modeling shows that melt volumeincrease was negligible and melt amount was too small to generate considerablemelt overpressure for calculated PT conditions. It is therefore suggested that dilation ofK-feldspar aggregates and fracturing of its grains represent a final creep failure state,which resulted from the cavitation process accompanying grain boundary slidingcontrolled diffusion creep. The consequence of cavitation-driven dilation of K-feldsparaggregates is the local underpressure resulting in infiltration of melt from plagioclasebands. Analogy with metallurgy experiments shows that the cavitation process,exclusively developed in cryptoperthitic K-feldspar, can be attributed to its lower puritycompared to more pure plagioclase. Contrasting rheological behavior of feldsparswith respect to quartz prior to fracturing is attributed to different deformationmechanisms. Feldspars appear weaker due to grain boundary sliding accommodated bycoupled melt-enhanced diffusion creep along grain boundaries and dislocation creepwithin grains, in contrast to quartz deforming via grain boundary migrationaccommodated dislocation creep.Citation: Závada, P., K. Schulmann, J. Konopásek, S. Ulrich, and O. Lexa (2007), Extreme ductility of feldspar aggregates—Meltenhancedgrain boundary sliding and creep failure: Rheological implications for felsic lower crust, J. Geophys. Res., 112, B10210,doi:10.1029/2006JB004820.1. Introduction[2] It has been experimentally demonstrated that smallamount of interstitial melt increases creep rates of deformingrocks and can induce switch of deformation mechanisms[Cooper and Kohlstedt, 1984; Dell’Angelo et al., 1987;Dell’Angelo and Tullis, 1988]. Dell’Angelo et al. [1987]have shown transition from dislocation creep to meltenhanceddiffusion creep in fine-grained granitic aggregates1 Institute of Petrology and Structural Geology, Charles University,Prague, Czech Republic.2 Centre de Geochimie de la Surface, UMR 7516, Université LouisPasteur, Strasbourg, France.3 Czech Geological Survey, Prague, Czech Republic.4 Geophysical Institute, Czech Academy of Sciences, Prague, CzechRepublic.Copyright 2007 by the American Geophysical Union.0148-0227/07/2006JB004820$09.00at small volume fractions of melt F = 0.01–0.03 contemporaneouslywith strength drop below the limit of detection.Rosenberg and Handy [2005] have shown that the ‘‘meltconnectivity transition’’ marked by melt fraction F = 0.07 iscritical in mineral aggregate strength drop at experimentalconditions. These results imply that small amount of meltcan be responsible for considerable weakening of crustalrocks, which can explain, e.g., the development of largescaleshear zones bounding regions of rapid uplift [Hollisterand Crawford, 1986].[3] Melt bearing microstructures in deformed rocks canbe also used for reconstruction of grain-scale melt migrationpathways, because deformation and melt extraction arecoupled [Brown and Rushmer, 1997; Rosenberg and Handy,2000; Rosenberg, 2001]. A typical feature of deformationexperiments at low melt volumes is the preferential distributionof melt along grain boundaries oriented at low angleto principal compressive stress direction [Dell’Angelo et al.,B102102791of15

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Figure 1. Simplified geological map of the studied area around the Eger river valley, NE of Stráž nadOhří town [Hradecký, 2002] (scale 1:25000) and the position of the studied area in the Bohemian Massif[Franke, 2000] and European Variscides. AM, Armorican massif; MC, Massif Central; BM, Bohemianmassif.1987; Daines and Kohlstedt, 1997; Gleason et al., 1999;Rosenberg and Handy, 2000]. This phenomenon can beexplained either by dilation of granular aggregates due tomelt overpressure [Renner et al., 2000], when hydrostaticmelt pressure overcomes the confining pressure, or bypassive melt migration into the incipient voids during grainboundary sliding (GBS) [Čadek, 1988].[4] However, studies of natural microstructures showingevidence of large finite strains and interstitial quenchedmelt, which would be indicative for pronounced meltenhancedweakening and grain-scale melt migration arepractically absent. This is due to two reasons: (1) Naturalmelt-bearing microstructures equilibrated at high-grade conditionsare likely to be overprinted by subsequent static ordynamic recrystallization [Rosenberg and Riller, 2000;Rosenberg, 2001]. (2) High finite strains, slow creep ratesand high homologous temperatures necessary to introducehigh-temperature melt-enhanced deformation mechanisms(e.g., diffusion creep or dislocation climb accommodatedGBS [Boullier and Gueguen, 1975]) cannot be attainedusing conventional experimental facilities. There are threeapproaches that may help to understand the rheologicalbehavior of minerals deformed at high-grade conditions inthe presence of melt and their influence on rheology ofpolyphase rocks: (1) Investigation of high-grade rocks,where microstructural record was preserved due to rapidcooling rates, (2) comparison of high-grade rock microstructureswith results of various experiments on creepproperties of metals, alloys and steels, and (3) applicationof more efficient rock deformation laboratory facilities (e.g.,Paterson rig apparatus).[5] In this work, we characterize deformation microstructuresin a strongly deformed high-grade banded orthogneiss,where melting was concurrent with deformation. Meltbearing microstructures have been preserved due to rapidcooling that effectively increased quenching of melt instudied rocks [Zulauf et al., 2002]. Deformation mechanismsof constituent mineral phases are investigated usingquantitative microstructural analysis. The mechanism responsiblefor preferred distribution of interstitial melt insample with maximum strain intensity is critically evaluatedusing modeling of melt volume produced by a metamorphicreaction and molar volume change of the system. The effectof mobile melt phase on deformation mechanisms and bulkrock rheology is discussed. Finally, we consider the apparentstrength relationship between quartz and feldspars.2. Geological Setting[6] The rocks studied belong to a lower crustal hightemperaturemetamorphic unit that was rapidly exhumedalong the major Variscan tectonic boundary in the BohemianMassif called the North Bohemian Shear Zone (NBSZ ofZulauf et al. [2002]). The trend and spatial limits of theNBSZ are not defined in the field, because it is mostlycovered by Neogene volcanics and sedimentary sequences(Figure 1). The sampled unit forms the uppermost thrustsheet within a crustal nappe stack that displays typicalmetamorphic inversion and is called the ‘‘Upper CrystallineNappe’’ (UCN) [Konopásek and Schulmann, 2005]. TheUCN in the studied area consists of banded orthogneisses,high-pressure granulites and migmatites (Figure 1). PTcalculations of metamorphic conditions for the partiallymolten orthogneisses in this study reveal pressures of900 MPa and temperatures 700 ± 20°C. In contrast,the granulites exhibit peak metamorphic conditions of1600 MPa and 800°C [Kotková etal., 1996] indicating thatthe granulites and orthogneisses have been juxtaposed atmiddle crustal levels during differential exhumation. Furthermore,very fast cooling rates of 50°C +25°C/ 17°CMa 1 were deduced [Zulauf et al., 2002] from the U-Pbdating of formation of zircons and monazites (342 ± 1 Ma)and 40 Ar- 39 Ar cooling ages of muscovites and biotites(341 ± 4 Ma) from the orthogneisses.2of15280

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Figure 2. (a) Traced and digitized mineral phase boundaries in XZ and YZ rock sections of themylonite type 3 sample (B109) and XZ section of a metagranite type 1 sample (B108) for comparison(mineral abbreviations after Kretz [1983], Mat means matrix and represents indiscernible mineral phases).(b) Flinn diagram of the phase aggregate’s shapes. R xy and R yz designate X/Y and Y/Z ratios of the strainellipsoid axes, respectively [Flinn, 1962].[7] Intensive north-south Variscan compression resultedin the development of vertical foliations in both the UCNand adjacent middle crustal rocks (LCN, Lower CrystallineNappe) [Konopásek et al., 2001] and also vertical stretchinglineations in the UCN. Zulauf et al. [2002] proposed thatthis vertical fabric is related to the rapid vertical motion oflower crustal rocks along the NBSZ and to intense deformationof gneisses and granulites. The UCN orthogneissshows significantly higher finite strains of feldspars comparedto relatively weakly deformed quartz suggestingextreme ductility of feldspars. In this work we focus onthe evaluation of the competency relationship betweenfeldspars and quartz, which can be of great importance forthe rheology of felsic lower crust.3. Finite Strain Analysis[8] In order to evaluate the influence of individual mineralson the rheology of UCN orthogneiss, the specimenshave been sampled according to the degree of finite strain.Plates cut along the XZ and YZ planes of the rock fabricellipsoid (XZ, parallel to lineation and perpendicular tofoliation; YZ, perpendicular to lineation and foliation) werecolored [Gabriel and Cox, 1929] in order to distinguish K-feldspar from plagioclase aggregates. The boundaries ofmineral phases in both XZ and YZ sections were thenmanually traced and digitized from photographs in ArcViewGIS environment (Figure 2a). The phase aggregate objectswere then statistically evaluated using the PolyLX Matlabtoolbox [Lexa et al., 2005] (http://petrol.natur.cuni.cz/ondro/polylx:home). Because both feldspars in bandedorthogneiss samples formed apparent infinite prolate-shapedaggregates, their strain was calculated using a method basedon normalization of their volume by the mean volume of thephase aggregates in sample with low finite strain intensity(where the phase objects are closed, sample B108). It isassumed that initially nonbanded feldspar aggregates coalescedtogether during the deformation. Quartz aggregatesin the XZ section appeared to retain similar cross-sectionalarea as in the sample with lowest strain intensity. Wetherefore suggest that quartz aggregates did not undergoextensive pinching by other phases during deformation,which would decrease their cross-sectional area and finitestrain estimate. Bulk deformation was calculated from allphases together as object area weighted geometrical mean.[9] Results of shape analysis are demonstrated in theFlinn diagram [Flinn, 1962] (Figure 2b). Sample B108shows the lowest degree of deformation and represents a‘‘metagranite’’ type 1. The ‘‘banded orthogneisses’’ type 2(samples B13, B109-2) show rod-shaped aggregates of bothfeldspars and plane strain to slightly prolate symmetry ofquartz. This type shows bulk intensities around D = 5 (D =(R 2 xy +R 2 yz ) 1/2 [Ramsay and Huber, 1983], D of K-feldsparclose to 10 and of plagioclase around 6. Maximum bulkintensity D = 16 was attained in sample B109 with K-feldspar and plagioclase reaching D = 26 and D = 50,respectively. This sample represents the final member ofmicrostructural evolution and will be designated the‘‘mylonite’’ type 3 sample. All stages of deformation showhigher finite strains in K-feldspar and plagioclase aggregatescompared to quartz.4. Deformation Microstructures[10] Microstructures and mineral textures of all rock‘‘types’’ were studied in detail in order to evaluate theevolution of deformation mechanisms with respect to increasingbulk rock strain. Quantitative microstructural analysiswas carried out using manually digitized grainboundaries [see Lexa et al., 2005] from several sequentialmicrophotographs of colored XZ thin sections. Texturalanalysis comprises quantification of the grain size and grainsize spread, grain shape (axial ratio) and shape preferredorientation (SPO), expressed as the ratio of eigenvalues of3of15281

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Table 1. Microstructural Analysis Data From Digitized ThinSections a Ferret, mm Axial RatioKfs Pl Qtz Kfs Pl QtzB108 0.32 0.25 0.47 1.40 1.44 1.51B13 0.30 0.26 0.42 1.48 1.42 1.49B109-2 0.26 0.17 0.47 1.42 1.45 1.65B109 0.12 0.09 0.90 1.43 1.49 1.58Q 3 –Q 1 , mmSPOKfs Pl Qtz Kfs Pl QtzB108 0.21 0.16 0.35 1.20 1.18 1.14B13 0.22 0.18 0.23 1.30 1.21 1.08B109-2 0.20 0.15 0.25 1.17 1.07 1.12B109 0.07 0.05 0.93 1.18 1.13 1.45a The grain size represented by the Ferret diameter is the diameter ofcircles with the same area as the measured grain. SPO is the eigenvalueratio of the shape preferred orientation tensor. Q 3 –Q 1 designates grain sizespread quantified as the difference between third and first quartile.the shape preferred orientation tensor [Lexa et al., 2005].These are important parameters when determining thedeformation mechanisms in deformed rocks [Kruse et al.,2001; Schmid et al., 1999; Ulrich et al., 2002].[11] Crystallographic preferred orientation (CPO) of quartz,plagioclase and K-feldspar was determined using the electronbackscattered diffraction method (EBSD using the CHAN-NEL5 software) in manual mode. Each thin section wascalibrated by a silicone monocrystal and only measurementswith mean angular deviation below 1 were accepted and saved.The output data were plotted in pole figures (ftp://www.gm.univ-montp2.fr/mainprice//CareWare_Unicef_Programs/) andpresented on lower hemisphere stereographic projections inwhich the trace of foliation is oriented in E-W direction. In caseof K-feldspar and plagioclase, pole figures of published slipplanes and slip directions [Tullis, 1983;Kruse et al., 2001]were plotted in order to select and present possible operativeslip systems.4.1. Microstructural Analysis[12] Microscopic observations, grain size and shape characteristicsare described below for three different typesmarked by increasing degree of strain intensity. The type1 (metagranite) and type 2 (banded orthogneiss) revealcompletely recrystallized equigranular mineral aggregateswith similar shapes of mineral grains without relics oforiginal magmatic porphyroclasts. Quartz aggregates consistof large grains with highly lobate mutual quartz boundariesand numerous ‘‘left-over grains’’ [Jessell, 1987]. Grainboundaries with neighboring feldspars as well as feldsparinclusions in quartz exhibit cuspate-lobate phase boundaries,typical for diffusional mass transfer along interphaseboundaries in high-temperature quartzo-feldspathic rocks[Gower and Simpson, 1992; Martelat et al., 1999]. Monomineralicfeldspar ribbons are composed of equidimensionalrecrystallized grains marked by straight grain boundariesshowing 120° triple point junctions. Average grain size,expressed as geometrical mean, ranges between 0.26 and0.32 mm for K-feldspar and between 0.17 and 0.26 mm forplagioclase. Grain size distribution of these minerals ismarked by low grain size spread (K-feldspar 0.2–0.21mm and plagioclase 0.15–0.18 mm) quantified as thedifference between third and first quartile (Table 1). Quartzexhibits average grain size ranging between 0.42 and 0.47mm and significantly larger grain size spread (Q 3 –Q 1 =0.23–0.35 mm) in comparison with feldspars. Axial ratio ofall mineral phases is very low in conjunction with low SPO(see Table 1).[13] The microstructure of type 3 sample shows alternatingbands of feldspars and quartz. Elongate quartz aggregatesand one-grain thick ribbons consist of large quartzgrains with gently curved to straight mutual boundariesperpendicular to ribbon boundaries. Both feldspars formmonomineralic bands with polygonal mosaic of feldspargrains with a triple junction network similar to microstructureof types 1 and 2. Grain size of both minerals hasconsiderably decreased to 0.12 mm and 0.9 mm for K-feldspar and plagioclase in comparison with type 1 and 2samples. The grain size spread also decreased (from averageQ 3 –Q 1 = 0.21–0.07 mm for K-feldspar and from averageQ 3 –Q 1 = 0.16 mm to 0.05 mm for plagioclase) and the axialratio together with SPO remain low (Table 1). In contrast,quartz exhibits significant increase in grain size (from0.47 mm in type 2 to 0.9 mm in type 3), grain size spread(from average 0.28 in types 1 and 2 to 0.93 in type 3) andSPO (Table 1).4.2. Crystallographic Preferred Orientations[14] The crystallographic preferred orientation of grainsof all mineral phases was measured in two samples; amoderately deformed type 2 banded orthogneiss and ahighly strained type 3 mylonite using the EBSD method.Only quartz texture in sample B13 (type 2) was measuredusing the U stage.[15] K-feldspar and plagioclase in type 2 sample exhibitweak CPO that do not show any coincidence with publishedslip systems [Kruse et al., 2001; Tullis, 1983] (Figure 3).The only exception can be seen in case of K-feldspar thatshows maxima compatible with the activity of [110](001)slip system [Willaime et al., 1979]. C axes pattern of quartzis characterized by incomplete crossed girdles with twomaxima distributed along the periphery of the diagramhaving mutual distance of 90° and weak central submaximum.This type of quartz CPO corresponds to type I crossgirdle pattern of Lister and Hobbs [1980] that is typical forcombination of basal hai and prism hai slip systems active athigh-temperature plane strain deformation [Lister andDornsiepen, 1982; Morgan and Law, 2004]. The form ofquartz c axes pattern is consistent with the shape ofmeasured plane strain ellipsoid of quartz aggregates(Figure 2b) [Passchier and Trouw, 1996].[16] K-feldspar in the type 3 sample shows crystallographicpreferred orientation marked by [100] directionsforming weak maximum close to the lineation (Figure 4).This slip direction operates on (010) planes [Willaime et al.,1979] that exhibit weak maximum in the YZ plane at anangle of 30° with respect to Y axis of the specimen’scoordinate system. Inspection of K-feldspar CPO patternrevealed that [001](120) slip system can be also active[Willaime and Gandais, 1977]. Plagioclase grains showweak CPO and reveal some activity of three slip directions,namely, [201], [100] and [101] on the (010) slip planereported by Marshall and McLaren [1977] and Olsen andKohlstedt [1984]. Quartz CPO is characterized by two broadvertical parallel girdles of c axes indicating again combined4of15282

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Figure 3. CPO data for K-feldspar, plagioclase and quartz from the type 2 (banded orthogneiss) sample(B13). Contoured at multiples of uniform distribution, lower hemisphere, equal-area projection. X and Zdesignate direction of lineation and pole to the foliation, respectively. Mineral abbreviations are afterKretz [1983].Figure 4. CPO data for K-feldspar, plagioclase, and quartz from the type 3 (mylonite) sample (B109).Contoured at multiples of uniform distribution, lower hemisphere, equal-area projection. X and Zdesignate direction of lineation and pole to the foliation, respectively. Mineral abbreviations are afterKretz [1983].5of15283

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210activity of basal hai and prism hai slip systems [Lister andHobbs, 1980]. This CPO pattern is typical of constrictionaltype of deformation that is consistent with prolate symmetryof quartz aggregates from this sample [Passchier andTrouw, 1996].Figure 5. BSE image of the type 3 (mylonite) bandedmicrostructure. Note the circular inclusions of K-feldspar inquartz. Scale bar is 500 mm. Inset shows plagioclasecomposition (Or, orthoclase component; Ab, albite component).Rhombs indicate band-forming plagioclase ofoligoclase composition. Triangles indicate interstitial albitein K-feldspar aggregates or oligoclase rims. Mineralabbreviations are after Kretz [1983].5. Melt Topology[17] Investigation of melt topology is important for understandingthe possible influence of melt phase on operativedeformation mechanisms of its host aggregates. Melttopology also reflects grain-scale melt migration pathways[Marchildon and Brown, 2001; Rosenberg, 2001]. Sawyer[2001] reviewed criteria for recognition of former presenceof melt on grain scale in deformed rocks. A typical featureof some rapidly quenched melting experiments is thedevelopment of melt pools with cuspate margins [Jurewiczand Watson, 1984] or thin melt films along crystal faces[Daines and Kohlstedt, 1997]. Melt phase crystallized asalbite, quartz or K-feldspar grains at triple point junctions orat crystal faces in residual aggregates have been commonlyreported from natural examples and interpreted in terms ofmelt topology in partially molten granite [e.g., Brown et al.,1999; Rosenberg and Riller, 2000; Marchildon and Brown,2001; Rosenberg, 2001].[18] In our study, polygonal mosaic of K-feldspar grainscontains numerous interstitial quartz and albite grains up to50 mm wide (Figures 5 and 6a) that extend along singleK-feldspar facets in the XZ section. Locally, narrowalbites (An 0.02 ) margin residual oligoclase (An 0,17 ) grainsFigure 6. Details of BSE images (type 3, mylonite) showing microstructures interpreted to mimic thetopology of crystallized interstitial melt. (a) Interstitial grains of quartz and albite between grains of K-feldspar aggregate developed especially on grain boundaries perpendicular to lineation (L). K-feldspar,light grey; quartz and albite, dark grey; scale bar 500 mm; XZ section. (b) Scarce triangular grain of K-feldspar between mica and oligoclase grain. Note fine exsolution lamella of albite adjacent to oligoclasegrain (arrow). K-feldspar, white; scale bar 250 mm; XZ section. (c) Albite rims on residual oligoclasegrains adjacent to K-feldspar aggregates (arrow 1), partial replacement of oligoclase grain by K-feldspar(arrow 2), amoeboid quartz grains at triple point junctions of plagioclase, adjacent to K-feldsparaggregates (arrow 3). Scale bar 250 mm; XZ section. (d) K-feldspar in host plagioclase grains (arrow 1),wedge-shaped albite grains within muscovite grains (arrow 2). Dark patches, holes in the specimen; scalebar 250 mm; XZ section. Mineral abbreviations are after Kretz [1983].6of15284

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Figure 7. Block diagram of the mylonite sample microstructure.The melt field in the legend represents grains ofphases assumed to have crystallized from interstitial melt.Note the melt topology; grain face pockets in XZ sectionand triple-point pockets in YZ section. Mineral abbreviationsare after Kretz [1983].(Figure 6c). Numerous interstitial round quartz grains withhigh dihedral angle occur in triple point junctions of thehost plagioclase aggregates often close to adjacent K-feldspar bands (Figure 6c). We suggest that these grainsdo not represent crystallized melt but resulted from annealingof disintegrated myrmekite fronts [Hanmer, 1982]. Locally,triangular K-feldspar grains with thin albitic ‘‘exsolution’’rim adjacent to oligoclase grains occur in plagioclasebands (Figure 6b). Thin wedge-shaped albite grains aredeveloped along (001) cleavage planes of muscovite grainsin plagioclase aggregates (Figure 6d). Plagioclase grainsadjacent to muscovite grains exhibit a rectangular networkof K-feldspar and albite, which is interpreted in terms ofmelt penetration along the cleavage planes of plagioclase(Figure 6d) [Mehnert et al., 1973; Dell’Angelo and Tullis,1988]. Quartz–K-feldspar boundaries show cuspate-lobateshapes and these boundaries are decorated with numeroussmall amoeboid grains of K-feldspar with high dihedral angleadjacent to the K-feldspar cusps (Figure 5).[19] Quartz and albite grains in K-feldspar aggregate,albite rims and scarce K-feldspar ‘‘pools’’ with albite‘‘exsolutions’’ in plagioclase bands, K-feldspar in plagioclasegrains, together with albite grains within mica crystalsand amoeboid grains of K-feldspar in quartz bands satisfythe criteria for presence of crystallized mineral phasesassumed to mimic residual melt topology [Rosenberg andRiller, 2000; Sawyer, 2001] and will be designated as ‘‘themelt’’ for simplicity. K-feldspar component crystallizedfrom melt probably grew mostly onto the older grainswithin K-feldspar bands, which would also explain theirperthite-free margins [Zulauf et al., 2002]. The crystallizationof melt as unlike phases in residual aggregates commonlyresults in disappearance of the banded rock texture.Transitions from a banded orthogneiss into a homogeneousnonfoliated rock can be observed on a single outcrop in thestudied area. However, in type 2 and type 3 samples, themelt distributes preferentially in K-feldspar bands.[20] The SEM imagery allowed identifying melt topologyalong intergranular voids as well as in intragranular fractures(Figures 5 and 6a). Quantitative analysis of meltpockets and their orientations allows creating a threedimensional(3-D) geometrical reconstruction of melt distributionin K-feldspar aggregates and analysis. Furthermore,we can define the geometrical relationship of intragranularfractures with crystal orientation and rock fabric.5.1. Intergranular Melt Topology[21] In order to depict and evaluate in detail the grainscalemelt distribution, grain boundaries on backscatteredelectron (BSE) images were traced and digitized from bothXZ and YZ sections (Figure 7). In moderately deformedorthogneiss type 2, melt in K-feldspar aggregate in XZsection occupies triple point junctions and extends into thinwedge-shaped melt films (maximum 30 mm in width) alonggrain faces perpendicular to the stretching lineation. Inhighly deformed orthogneiss type 3, the intergranular meltpockets (or seams [Rosenberg and Riller, 2000]) of aspectratio 2–3 show two submaxima in R f /f graph inclined at±20° relative to the apparent Z axis of the rock fabricellipsoid (Figure 8). Melt topology in YZ section of type3 orthogneiss is characterized by presence of melt ‘‘droplets’’at triple point junctions of K-feldspar grains. Rarely,the melt forms seams (30 mm in width) that line grainboundaries perpendicular to the foliation. These wide pocketsare likely to result from oblique sections of seamsoriented at high angle to the principle stretching direction.The melt topology from two perpendicular sections allowscreating a simplified 3-D geometrical model of interstitialmelt distribution in K-feldspar aggregate. This model ischaracterized as an interconnected network marked bymelt walls parallel to YZ plane, connected by tubes parallelto X direction along triple point boundaries between threegrains (Figure 9).5.2. Intragranular Fractures Locally Filled by Melt[22] A closer observation of XZ section BSE images ofthe type 3 (mylonite) sample revealed that some K-feldspargrains are crosscut by intragranular fractures (Figure 10a)and some of these fractures are filled with wedge-shapedmelt films. Orientations of traces in XZ section of theseFigure 8. R f /f diagram of the intergranular melt pocketswithin the K-feldspar aggregate from the XZ section.7of15285

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Figure 9. Simplified 3-D melt topology model in K-feldspar aggregate of the mylonite type 3 orthogneiss. Seetext for explanation.fractures are represented by a rose diagram in Figure 10band show strong maximum subparallel to the Z direction ofthe rock fabric ellipsoid and perpendicular to the horizontalstretching lineation (X direction). Furthermore, we havemeasured the CPO of grains affected by intragranularfractures, which exhibit clustering of poles to (001) planesin stretching lineation direction, while poles to (010) revealmaximum in the Y direction of the sample coordinatesystem (Figure 11). In order to evaluate relationship betweenorientations of crystallographic directions and intragranularfractures within corresponding grains, theorientations of fractures measured by U stage and graincrystallographic orientations obtained by means of EBSDhave been compared. Poles to all measured fractures exhibitstrong maximum close to the X direction, which is extendedtoward the Z direction (Figure 11). Analysis of anglesbetween pole of fracture and pole to (001) of K-feldsparcrystal shows that small angles characterize grains with theirpoles to (001) planes subparallel to the stretching lineation.In contrast, this angle is high for grains with (001) poles athigh angle to the stretching lineation (Figure 11). Thestatistics of angles between fracture orientations and (001)planes clearly shows that the majority of fractures aresubparallel to the (001) plane (in 52% of the fracturedgrains the angle is less than 30°, see Figure 12). Morphologyof the fractures shows distinctive features for bothgroups of grains with low (60°) angle.In group 1 (60°), the grains show single curved fractures orinward tapering melt intrusions (Figure 10a). In addition,there exist intragranular curved fractures that were initiatedfrom triple point junctions and locally adjoin to the cleavageof the grain at their tips (Figure 10c).6. Melt Producing Reaction and Melt VolumeEstimates[23] An important issue is the origin of the interstitial meltdescribed in section 5. The melt could have been producedby in situ melting reaction or by infiltration from externalsources, which is reflected by the melt composition underestimated PT conditions and possible reaction textures[Sawyer, 2001]. Another important issue is the melt volumeproduced and volume change of the melting reaction, asboth these variables control the embrittlement of rock atdynamic conditions during breakdown reactions of somehydrous phase and/or at sufficient heating rates [Connolly etal., 1997; Rosenberg, 2001; Rushmer, 2001; Holyoke andRushmer, 2002]. These data have critical impact on thedescribed melt topology, deformation mechanisms andrheology of studied orthogneiss and will be further discussed.Figure 10. (a) Detail of two fractured grains. Arrows and numbers designate group 1 and 2 fracturedgrains. Scale bar 100 mm. (b) Rose diagram of traces of the intragranular fracture orientations in K-feldspar bands (XZ section of the sample coordinate system). (c) Sketch of curved and splittingintragranular fractures initiated from triple point junctions. Melt is gray; straight lines in the grains markthe cleavage.8of15286

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Figure 11. Stereographic projection of CPO data and fracture poles of 100 fractured grains. X and Zdesignate direction of lineation and pole to the foliation, respectively. Symbols in the center projectiondesignate the misorientation angle of (001) planes with respect to the corresponding fracture pole (right).Lower hemisphere, equal-area projection, contoured at multiples of uniform distribution.[24] The rocks studied are represented by banded orthogneissesconsisting of K-feldspar (Kfs) (39%), plagioclase(Pl) (25%), quartz (Qtz) (26%), biotite (Bt), and muscovite(Ms) (9%) and minor garnet (Grt) (1%). P-T pseudosectionfor this rock was calculated with the Perple_X software set[Connolly, 1990; Connolly and Petrini, 2002] using phaseend-members thermodynamic data by Holland and Powell[1998]. Mixing properties of phases used for the calculationwere taken from Berman [1990] for garnet, Newton et al.[1980] for plagioclase, Thompson and Hovis [1979] for K-feldspar, Powell and Holland [1999] for biotite and clinopyroxene,and White et al. [2001] for melt.[25] The obtained pseudosection calculated inNCKFMASH system with composition taken from thewhole rock analysis of a typical banded orthogneiss (sampleB13) is shown in Figure 13a. Stable mineral assemblageKfs-Pl-Qtz-Ms-Bt-Grt-melt, and composition of coexistingFe-Mg phases suggest that the highest metamorphic temperaturesassociated with melting are 700°C at a pressure of9.5 kbar. Calculated isopleths of the melt mode indicatethat for the given melt model and estimated P-T conditions,the rock contained 2–4 vol % of the melt phase(Figure 13a) at its metamorphic peak. This is, however, aminimum estimate, because the chosen system compositiondoes not take into account the amount of water releasedduring melt crystallization. The model of White et al. [2001]anticipates that the melt does not contain any CaO componentat the estimated PT conditions, which is in goodagreement with albitic composition of plagioclase crystallizedfrom the interstitial melt in the rock. Generalizedreaction leading to melt production in the stability field ofassemblage Kfs-Pl-Qtz-Ms-Bt-Grt-melt can be derived fromthe changes in modal proportions of phases in the stabilityfield of interest (Figure 13b). Such changes suggest that theincrease in melt content is associated with the crystallizationof garnet as a result of the reaction: Bt + Ms + Pl + Qtz = Grt +Kfs + melt.[26] The inferred melt-producing reaction correspondswell with the distribution of mineral phases in the rockmicrostructure. Small garnet grains (50 mm) form clusters onboundaries between plagioclase bands and mica aggregatesin the vicinity of quartz grains (Figure 14). Agreement ofmodeled melt composition with that encountered in intergranularvoids of K-feldspar indicates that the melt migratedonly in between adjacent feldspar aggregates. Therefore wecan exclude melt loss or melt infiltration from externalsources, which would produce different reaction texturesand could result in crystallization of interstitial plagioclasewith different composition.7. Deformation Mechanisms[27] Microstructural features in quartz observed in samplesof all strain intensities reveal migrated grain boundariesand large grain size that increases even more during thedevelopment of type 3 microstructure. Microstructures andCPO patterns of quartz in type 2 and 3 samples show clearactivity of grain boundary migration (GBM) accommodateddislocation creep mechanism [Jessell, 1987; Hirth andTullis, 1992] and CPO patterns correspond to the shapesof deformation inferred from the strain analysis.[28] Microstructural data show isometric shapes of polygonalfeldspar grains, absence of shape preferred orientationand weak or no CPO for both feldspars in type 2sample. These microstructural features and high finite strainintensities attained are indicative for grain boundary diffusionaccommodated grain boundary sliding (D gb -GBS) inboth feldspars [Poirier, 1985; Wadsworth et al., 1999].Although large grain size (K-feldspar 300 mm and plagioclase200 mm) is not characteristic for such deformationmechanism at experimental strain rates [Tullis and Yund,Figure 12. Statistics of angles between poles to (001)planes and corresponding fractures crosscutting the grainsrepresented by frequency histogram and a doughnutdiagram. Colors in the doughnut diagram represent thethree groups of grains as depicted in the histogram.9of15287

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Figure 13. (a) P-T pseudosection for the sample B13 of granitic orthogneiss calculated in theNCKFMASH system and presented with quartz in excess. Isopleths of the volume percent of the melt inthe system are shown as solid black lines. Estimated PT conditions based on the position of compositionisopleths of coexisting garnet and biotite in the Kfs-Pl-Ms-Bt-Grt-melt stability field are marked by theblack field in the lower right inset. (b) Isobaric section (at 9 kbar) showing changes in modal proportionsof stable phases with increasing temperature. Changes in stable mineral assemblages are shown as labeledsolid vertical lines. Mineral abbreviations are after Kretz [1983].1991], at natural conditions and presence of melt, whicheffectively enhances diffusion creep, it may be possible[Dell’Angelo et al., 1987]. In type 3 microstructure, increaseof finite strain in both K-feldspar and plagioclase is associatedwith significant reduction of grain size (in contrast toquartz) and an important strengthening of CPO, which isremarkable especially in K-feldspar, while the shape preferredorientation (SPO) remains low. This is explained byincreased activity of dislocation creep [e.g., Tullis, 1983] ofboth feldspars in comparison with type 2 orthogneiss. Theopposite trend of grain size increase in quartz probablyreflects lower relative stress and strain dissipation in quartzwith respect to both feldspars. However, strong contributionof GBS accommodated by grain boundary diffusion flow(D gb -GBS) is considered to be responsible for accommodationof extreme finite strains in feldspars.[29] The suggested operative deformation mechanism infeldspars of type 3 sample (D gb -GBS + dislocation creepand climb within grains) was reported to result in extremefinite strains in several alloys deformed at relatively highstrain rates [Wadsworth et al., 1999; Weietal., 2003]. Finegrainedalloys undergoing D gb -GBS mechanism can beelongated up to several hundreds percents prior to creepfailure [Poirier, 1985; Čadek, 1988]. This behavior is called‘‘superplasticity’’ and it was also attributed to be responsiblefor deformation of some natural mylonitic rock bands composedof fine (10 mm), isometricaly shaped and CPOlacking grains [Boullier and Gueguen, 1975; Behrmannand Mainprice, 1987]. In general, it is very little knownabout high-temperature GBS mechanism operating inquartzo-feldspathic rocks, although it is regarded as one ofthe most important strain accommodation mechanisms activein mineral aggregates (Langdon and Vastava [1982] as citedby Zhang et al. [1994]; also Ranalli [1995]).8. Discussion[30] Detailed analyses of different microstructural typespresented in this work provided a unique view on processesoperating in the studied rock at high metamorphic conditionsand extreme finite strains. The mylonite type 3microstructure shows intergranular voids filled by meltand intragranular fractures both developed exclusively inK-feldspar aggregates. Microstructural observations showFigure 14. Micrograph of the studied orthogneiss indicatingthe melting reaction; scale bar 500 mm. Mineral abbreviationsare after Kretz [1983].10 of 15288

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Figure 15. Illustration of the dilatancy mechanism andfracturing of K-feldspar grains. Lineation parallel to s 3 isvertical in geographic coordinates. (a) Low proportion ofmelt shuffled in intergranular spaces during GBS accommodatedby melt-enhanced diffusion creep along grainboundaries and dislocation creep within grains in bothfeldspars. (b) Dilatancy driven by cavitation or by meltoverpressure in K-feldspar that extracts the melt fromplagioclase intergranular films. Arrows designate thedirection of melt flux.that melt was present in both plagioclase and K-feldsparaggregates during deformation. At first, origin of melt-filledseams and intragranular fractures developed within K-feldsparaggregates is critically discussed using petrologicaldata. We further consider the processes responsible forenhancement of solid-state flow, given by the presence ofmelt phase, and the role of its redistribution on the rheologyof polyphase and strongly deformed rock in terms ofdraining and accumulation of melt in different mineralaggregates. Finally, we discuss the apparent rheologicalrelationship between ‘‘weak’’ feldspars and ‘‘strong’’quartz.8.1. Cavitation Versus Fracturing Driven by MeltOverpressure[31] In studied rocks, the melt topology is characterizedby preferential distribution of melt in K-feldspar intergranularpockets, which is similar to experiments conductedwith low amounts of melt and at relatively high differentialstress [Daines and Kohlstedt, 1997; Gleason et al., 1999;Rosenberg, 2001; Holtzman et al., 2003]. In addition, K-feldspar grains are affected by fractures of the same orientation,which show affinity to the (001) cleavage. Ourpetrological investigations have shown that the melt wasproduced by metamorphic reaction within mica-plagioclasebands (source) and migrated into K-feldspar bands (sink).This implies that individual mineral aggregates building therock can be characterized as open systems (Figure 15).There are two possible mechanisms, which could haveproduced the observed melt topology and intragranularfractures in K-feldspar.[32] If a fluid is introduced to the rock, e.g., by meltproducing metamorphic reaction, its pressure may overcomethe least principal stress s 3 as well as cohesion ortensile strength of suitably oriented planes in the aggregate[Hubbert and Rubey, 1959; Price and Cosgrove, 1990].These planes would in our case be represented by grainboundaries and (001) crystallographic planes oriented perpendicularto the stretching lineation (s 3 direction). As aresult of melt overpressure, the K-feldspar aggregate dilatesand the melt is accumulated in intergranular pockets andsome intragranular fractures. The aggregate hardens due toincreased frictional stress on melt free boundaries, becausemodeling of melt productivity suggests that no significantamount of melt was gained from external sources (‘‘dilationhardening’’ of Renner et al. [2000]).[33] The second model, explaining production of intergranularmelt pockets and intragranular fractures is theprocess of cavitation. Cavitation, formation of submicroscopiccavities driven primarily by diffusion and accumulationof vacancies, is the direct consequence of GBS. WhenGBS cannot be fully compensated by diffusion and/ordislocation creep controlled grain shape change, cavitiesnucleate at first on grain boundaries at high angle to thetensional direction (direction of s 3 )[Čadek, 1988; Kassnerand Hayes, 2003]. Cavities then grow and coalesce witheach other to form intergranular voids. Further cavitycoalescence can be caused by nucleation and growth ofcavities on grain boundaries at low angle to the tensionaldirection due to (1) formation of tensile and compressiveledges, where boundaries are not straight due to inhomogeneousplastic deformation of the grains or (2) dislocationpileups at grain boundaries (Zener-Stroh mechanism) oraround impurities in crystal lattice (Figure 16) [Vollbrecht etal., 1999; Kassner and Hayes, 2003]. Plastic deformation isfurther being localized to the ‘‘bridges’’ (intact grain boundarysegments) affected by further cavitation. The bridgesloose its stability and locally contract to form finally afracture which is driven by coalescence of cavities at its tip(Figure 16). In this way, clear macroscopic intragranularfractures develop [Čadek, 1988].[34] According to the cavitation model, opening of intergranularvoids and intragranular fractures during GBSproduces local underpressure (the volume of the systemincreases due to creation of voids in K-feldspar and thegrains support the voids as a load bearing framework),Figure 16. (a) Illustration of the cavitation mechanismgenerated by GBS (grain boundary sliding) and formationof intergranular and intragranular fractures at the onset offinal creep failure of the aggregate. T and C designate thetensional and compressive sector of a grain boundary ledge.Z-S designates illustration of the Zener-Stroh mechanism ofcavity formation [Kassner and Hayes, 2003]. See descriptionin text.11 of 15289

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Figure 17. PT plot with contours of molar volume in Jbar 1 (equivalent to 10 1 cm 1 mol1 ) calculated for thesame system as shown in Figure 13. Mineral abbreviationsare after Kretz [1983].which soaks up the melt produced within plagioclase bands(the melt is passively redistributed, Figure 15). Consequently,plagioclase could not deform any further due to the loss ofintergranular melt that worked as a creep enhancing medium.[35] It has been shown that creation of melt-bearingporosity can be induced by fast melt production due tobreakdown of some hydrous phase in the rock (e.g., viadiscontinuous reaction) or due to high heating rate causingsubstantial overstepping of continuous melting reactionsand thus fast melt production as a result of the system’seffort to reach equilibrium state [e.g., Connolly et al., 1997;Rushmer, 2001]. In contrast, change in molar volume of therock during slow progress of continuous reactions will beprobably small.[36] In order to test this hypothesis, the P-T phasediagram section in Figure 13a was contoured for molarvolume of the system to see, whether the observed meltingreaction at estimated PT conditions may lead to rapidvolume increase. The result (Figure 17) shows that the onlyrapid increase in molar volume of the rock in given PTrange results from dehydration melting of muscovite at PTconditions higher than those reached by studied orthogneisssamples. In addition, the role of melt volume change andvelocity of the reaction is speculative for aggregates with‘‘incoherent’’ grain boundaries like both feldspars undergoingGBS in studied orthogneiss than rather ‘‘intact’’ solidphases around reaction sites in the experiments of Connollyet al. [1997].[37] The calculated volume of melt (2–4 vol %) suggeststhat the melt formed isolated melt films, pools or pockets. Inthe case of ‘‘overpressure’’ model, sufficient amount of meltneeds to be produced to induce considerable melt overpressureand melt redistribution. This critical melt volumecorresponds to creation of interconnected network of meltbetween the framework of grains, so that the melt can startincreasing its hydrostatic pressure toward the level ofmaximum compressive stress [Renner et al., 2000]. Thismelt volume is given by the melt connectivity transition(MCT) of F =0.07[Rosenberg and Handy, 2005] or similar‘‘liquid percolation threshold’’ (LPT) of F =0.08[Vigneresseet al., 1996], although this critical level is likely to depend onseveral variables (e.g., dihedral angles). It is possible that themelt migrates from locally overpressured isolated pockets athigh angle to the maximum compressive stress (s 1 )tointergranular boundaries subparallel with this direction(s 1 ). However, it is unlikely that the cohesion of grainboundaries and tensile strength of (001) intragranular cleavageof K-feldspar are significantly smaller than intergranularcohesion of plagioclase, which would be necessary conditionfor preferential hydraulic failure of K-feldspar, although themelt was produced in plagioclase bands.[38] In contrast, during cavitation, underpressure arisingfrom opening of incipient intergranular voids in K-feldsparis theoretically susceptible to ‘‘extract’’ melt from isolatedpools and films within plagioclase aggregates, in spite of itslittle amount. If distortion of plagioclase grains is easierthan for K-feldspar grains to accommodate GBS, then plagioclaseaggregates should also ‘‘extrude’’ excessive intergranularmelt into possible sink sites. This is in agreementwith analogue modeling results of Walte et al. [2005], whereaggregates composed of ‘‘weaker’’ grains (more susceptibleto plastic deformation) show higher ‘‘GBS locking transition’’than relatively stronger grains. Melt migration alongmutual boundaries of quartz was probably negligible incomparison to both feldspars. This is again controlled byrelative ‘‘weakness’’ of quartz grains with respect to ‘‘stronger’’grains of both feldspars and ‘‘welded’’ boundaries ofquartz. The viscosity ratio between melt and host aggregategrains thus increases from quartz to plagioclase and ishighest for K-feldspar [Walte et al., 2005].[39] Considering little amount of melt produced by themelting reaction and absence of intergranular melt pocketsperpendicular to the stretching lineation in plagioclase, weconsider the ‘‘cavitation’’ model to be the most likelymicrophysical mechanism explaining the dilation, microfracturationand preferential melt distribution in K-feldsparaggregates of type 3 microstructure.8.2. Impact of Impurities and MechanicalAnisotropy of Crystals on GBS[40] Experiments with aluminum in GBS regime withrelatively high-purity metals (e.g., aluminum of purity 4N(where N expresses the degree of purity of a material; 2N =99% purity, 4N = 99.99% purity, etc.), 650°K, strain rate10 9 s 1 ) or at high strain rates (e.g., 4N copper at 10 2 s 1 ,873°K [Čadek, 1988]) are marked by high relative elongationse, distinct contraction at the area of failure (necking)associated with intracrystalline fracturing. On theother hand, experiments with relatively low-purity metals(2N aluminum at 10 9 s 1 , 650°K) or at low strain rates(4N copper at 10 9 s 1 , 873 K) show small relativeelongations to failure, regular contraction along the deformedspecimen (homogeneous strain) and intergranularfracturing (Figure 18) [Sklenička et al., 1977; Čadek, 1988;Mohamed, 2002]. Therefore the higher purity material canwithstand higher relative elongations to creep failure (and atfaster strain rates) than its lower purity equivalent. Incontrast, stress accumulations around impurities and grainboundary ledges enhance diffusion-driven cavitation, cavitieseasily form at dislocation pileups around impuritieswithin the grains and creep failure will occur after smallerrelative elongations (and slower strain rates).12 of 15290

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB10210Figure 18. Final creep failure microstructure and specimendistortion from tensile experiments with metals atconstant strain rate. [Sklenička et al., 1977; Čadek, 1988].Material of purity 2N contains more impurities than 4Nmaterial.[41] These results can be compared with different intensitiesof deformation and microstructures of both feldspars inthe studied orthogneiss (type 3 microstructure) sample.Feldspars can be also regarded as solid solutions like alloysor impure metals. We consider the three component K-feldspar with clearly defined solvus as a ‘‘less pure’’ equivalentto the two component oligoclase. This interpretation issupported by the presence of cryptoperthite exsolutionswithin some K-feldspar grains (lower left part ofFigure 10a), along which the tips of cavitation drivenfractures could nucleate and propagate. The perthite intergrowthsare commonly oriented along (001) plane [Spry,1969], which is fully compatible with orientation of most ofthe intragranular fractures reported in this work. In contrast,plagioclase shows higher finite strains in type 3 orthogneissthan K-feldspar and no intragranular fractures, which isconsistent with higher purity aluminum experiments orhigher purity metals in general [Čadek, 1988; Mohamed,2002]. Fractured grains in K-feldspar exhibit strong crystallographicpreferred orientation marked by (001) planesperpendicular to the axis of stretching direction (X direction)and their (010) plane coinciding with the Y direction of therock fabric’s coordinate system. Because propagation offractures (group 1 fractures, Figure 10a) driven by coalescenceof cavities is initiated from boundaries oriented at highangles to the maximum principal compressive stress direction,we can suggest that cavity coalescence was mosteffective along the K-feldspar (001) plane in [100] crystallographicdirection. Selection of suitably oriented K-feldspargrains for intragranular cavitation reflects the strong mechanicalcrystallographic anisotropy of K-feldspar [Smith,1974]. Group 2 fractures (Figure 10a) and curved or splittingfractures (Figure 10c) could develop due to cavitation andpropagation of fractures from triple-point junction boundaries.The propagation of fractures in K-feldspars in the caseof ‘‘melt overpressure’’ model would probably follow similarpathways as cavitation-driven fractures.8.3. Effect of Melt Phase on GBS[42] Several tensile experiments on creep properties ofmetals have shown dominantly GBS accommodated superplasticbehavior with peak of finite elongations at temperaturescoinciding with occurrence of melt phase along grain boundaries[Mabuchi et al., 1997]. This weakening was associatedwith effective contribution of the melt phase to annihilation ofgliding dislocations, therefore prevention of dislocation pileups,increased velocity of diffusive mass transfer and accommodationof stress concentrations that was also reflected bylower cavity densities in comparison with ‘‘melt-free’’ experiments[Koike et al., 1998]. Melt-enhanced weakening isfavored by low wettability of melt (high dihedral angles) andits relatively small amount. In contrast, high wettability (lowdihedral angle) melts thickly coating grain boundaries producedpremature necking and creep failure due to the loss ofgrain boundary cohesion [Mabuchi et al., 1997; Koike et al.,1998]. The same transition from superplastic flow to prematurefailure coinciding with almost complete wetting of grainboundaries (at length proportion of 70%) was demonstratedalso during compressive experiments [Pharr et al., 1989;Baudelet et al., 1992].[43] In summary, the material science experimentalresults show distinct rheological transition between melt(liquid)–enhanced GBS and collapse marked by increasedcavitation velocities at a melt volume threshold coincidingwith complete wetting of grain boundaries. This threshold issimilar to the concept of melt embrittlement at the onset ofmelt connectivity transition (MCT) proposed by Rosenbergand Handy [2005] from experimental data on quartzofeldspathicrocks. Collapse (rather than weakening) at theMCT was associated with localized intergranular and intragranularmicrocracking, frictional sliding and limited bodyrotation leading to the development of cataclastic zonesand produces heterogeneous and restricted deformation[Rutter and Neumann, 1995; Rosenberg and Handy,2005]. Rosenberg and Handy [2005] have regarded MCTas a rheological threshold even more important than thepreviously emphasized rheological critical melt percentage(RCMP) at F = 0.1–0.3 [Arzi, 1978; van der Molen andPaterson, 1979] at experimental conditions. However, theweakening mechanisms associated with MCT are incompatiblewith our observations.[44] Our strain measurements and material science experimentalresults show that weakening associated with meltenhancedGBS at low melt volumes (below the MCT) issubstantial and the aggregate is deformed relatively homogeneously.We therefore suggest that the degree of weakeninginduced by melt-enhanced GBS can be higher andmore important on large scales at natural conditions than theweakening mechanisms operative at experimental strainrates at the onset of melt connectivity transition as proposedby Rosenberg and Handy [2005].[45] The critical melt volume for transition from meltenhancedGBS to premature creep failure due to increasedmelt volume is likely to depend on several variables, such asdifferential stress [e.g., Bordeaux et al., 1994; Rosenberg,2001], strain rate, grain size [Bordeaux et al., 1994; Renner etal., 2000], melt volume and melt framework viscosity ratio[McKenzie, 1984; Pharr et al., 1989; Walte et al., 2005].[46] The presence of water without melt in the systemcould effectively enhance diffusion and dislocation creepaccommodated GBS of feldspars and produce similarmicrostructures as in this study [Tullis and Yund, 1991].However, microstructural observations and petrological datashow that melt was present in the system during deformation.The combined effect of water and melt on GBSenhancement is speculative (according to the weakeningmechanisms of melt discussed above), because the influenceof water on dihedral angles of melt in quartzo-feldspathicrocks is not yet well understood [Rushmer, 1996].13 of 15291

B10210ZÁVADA ET AL.: EXTREME DUCTILITY OF FELDSPAR AGGREGATESB102108.4. Comparative Rheology Between Feldspars andQuartz[47] Experimental studies show that the creep strength offeldspars deformed at dislocation creep is higher than that ofquartz for the entire range of temperature conditions investigated[Shelton and Tullis, 1981; Jaoul et al., 1984; Handy,1994]. This assumption is essential in assessing rheology ofpolyphase rocks with different proportion of weak quartzand ‘‘strong skeleton’’ formed by feldspars in both loadbearingframework and interconnected weak layer microstructures[Handy, 1990, 1994; Schulmann et al., 1996; Ji etal., 2004; Rybacki and Dresen, 2004]. However, change ofdeformation mechanism in feldspars from dislocation todiffusion creep has never been considered in term ofmineral strength contrast between quartz and feldspars.Exceptionally high finite strains of both feldspars surroundingmuch less elongated quartz ribbons shows well thatquartz represents the strong phase according to Handy[1990, 1994] and feldspars form interconnected weaklayers. The explanation is seen in important contributionof dislocation creep accommodated GBS mechanism to thedominant melt-enhanced grain boundary diffusion creepoperative in both feldspars that significantly decreases theirstrength with respect to quartz simultaneously deforming viadislocation creep-controlled recrystallization mechanisms.[48] Acknowledgments. This work has been kindly supported byGrant Agency of the Czech Academy of Sciences (GAAV) projectIAA3111401. 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Valverde-Vaquero (2002), Evidence for high-temperature diffusional creep preservedby rapid cooling of lower crust (North Bohemian shear zone,Czech Republic), Terra Nova, 14(5), 343–354.J. Konopásek, Czech Geological Survey, Prague, Czech Republic.(konopasek@cgu.cz)O. Lexa and K. Schulmann, Centre de Geochimie de la Surface, UMR7516, Université Louis Pasteur, Strasbourg, 67000 Cedex, France.(lexa@illite.u-strasbg.fr; schulman@illite.u-strasbg.fr)S. Ulrich, Geophysical Institute, Czech Academy of Sciences, Prague,Czech Republic. (stano@ig.cas.cz)P. Závada, Institute of Petrology and Structural Geology, CharlesUniversity, Prague, Czech Republic. (zavada@natur.cuni.cz)15 of 15293

ClickHereforFullArticleJOURNAL OF GEOPHYSICAL RESEARCH, VOL. 113, B10406, doi:10.1029/2007JB005508, 2008Evolution of microstructure and melt topology in partially moltengranitic mylonite: Implications for rheology of felsic middle crustKarel Schulmann, 1 Jean-Emmanuel Martelat, 2 Stanislav Ulrich, 3,4 Ondrej Lexa, 4Pavla Štípská, 1 and Jens K. Becker 5Received 19 November 2007; revised 27 April 2008; accepted 30 July 2008; published 24 October 2008.[1] The deformation study of midcrustal porphyritic granite reveals exceptionally highstrain intensities of feldspar aggregates compared to stronger quartz. Three types ofmicrostructures corresponding to evolutionary stages of deformed granite wererecognized: (1) the metagranite marked by viscous flow of plagioclase around strongalkali feldspar and quartz, (2) quartz augen orthogneiss characterized by development ofbanded mylonitic structure of recrystallized plagioclase and K-feldspar surroundingaugens of quartz, and (3) banded mylonite characterized by alternation of quartz ribbonsand mixed aggregates of feldspars and quartz. The original weakening of alkali feldspar isachieved by decomposition into albite chains and K-feldspar resulting from aheterogeneous nucleation process. The subsequent collapse of alkaline feldspar anddevelopment of monomineralic layering is attributed to the onset of syn-deformationaldehydration melting of Mu-Bi layers associated with production of 2% melt. The finaldeformation stage is marked by mixing of feldspars which is explained by higher meltproduction due to introduction of external water. An already small amount of melt isresponsible for extreme weakening of the feldspar because of Melt ConnectivityThreshold effect triggering grain boundary sliding deformation mechanisms. The grainboundary sliding controls diffusion creep at small melt fraction and evolves to particulateflow at high melt fractions. Strong quartz shows a dislocation creep deformationmechanism throughout the whole deformation history marked by variations in the activityof the slip systems, which are attributed to variations in stress and strain rate partitioningwith regard to changing rheological properties of the deforming feldspars.Citation: Schulmann, K., J.-E. Martelat, S. Ulrich, O. Lexa, P. Štípská, and J. K. Becker (2008), Evolution of microstructure and melttopology in partially molten granitic mylonite: Implications for rheology of felsic middle crust, J. Geophys. Res., 113, B10406,doi:10.1029/2007JB005508.1. Introduction[2] Rheology of the continental crust is dominated byquartzo-feldspathic rocks, which are represented mainly bymetagranitoids, orthogneisses and felsic volcanics [Carterand Tsenn, 1987]. To date, the models of crustal rheologyuse laboratory derived laws described by constitutive equationsthat are established for minerals or monomineralicrocks such as quartzites and anorthosites [Ranalli, 1995].Most of laboratory experiments show that the quartz is1 Centre de Géochimie de la Surface, UMR7516, Université LouisPasteur, CNRS, Strasbourg, France.2 Laboratoire de Géodynamique des Chaînes Alpines, UMR5025,Université Joseph Fourier, Observatoire des Sciences de l’Univers deGrenoble, CNRS, Grenoble, France.3 Geophysical Institute, Czech Academy of Sciences, Prague, CzechRepublic.4 Institute of Petrology and Structural Geology, Charles University,Prague, Czech Republic.5 Institut für Geowissenschaften, Universität Tübingen, Tübingen,Germany.Copyright 2008 by the American Geophysical Union.0148-0227/08/2007JB005508$09.00weaker than plagioclase for the same homologous temperatures[Ranalli and Murphy, 1987; Schmid, 1982]. However,the natural quartzo-feldspathic rocks are mixtures withdifferent proportions of strong feldspars and weak quartzwith variable grain shapes and grain size distributions. Thedeformation of such natural rocks leads to strain partitioningbetween the different components and nonuniform deformation[Handy, 1990]. Handy [1994a] defined the loadbearingframework structure and interconnected weak layerstructure and proposed comprehensive empirical equationsthat determine the strength of polyphase composites. Handy[1994a] applied this concept to quartzo-feldspathic rocksand concluded that the proportion of weaker quartz controlsthe bulk rheology. The basis of these models is the coexistenceof two nonlinear viscous phases; the bulk rheologyis a consequence of the rock structure and the relativeproportions of the two mineral phases [Ji and Zhao, 1994].[3] Microstructural studies show that the progressiveorthogneiss deformation is associated with strain partitioningand variations in the deformation mechanisms offeldspars and quartz [Handy et al., 1999; Schulmann etal., 1996; Simpson, 1985]. For instance Gapais [1989] andB104062951of20

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Figure 1. Geological map of the central part of Bohemian Massif modified after Beneš [1964] andSynek and Oliverová [1993]. Black stars and capitalized letters refer to studied samples (B, D, H, M, R, S,T, and V). (a) Lower crustal rocks (granulites and eclogites), (b) the midcrustal monotonousmetasedimentary unit, (c) the midcrustal orthogneiss unit, (d) the kyanite micaschist unit, (e) theintrusives, (f) Cretaceous and Quaternary rocks, (g) the undifferenciated Lower Palaeozoic rocks, and(h) the Neo-Proterozoic metasediments.Schulmann et al. [1996] have shown that at amphibolitefacies the feldspars show evolution from dislocation creepto grain boundary sliding with increasing strain intensity inconjunction with variations in the activity of the quartz slipsystems. Therefore, the viscosity contrast between feldsparsand quartz was not constant but varied with increasingdegree of strain as the rheological role of individualminerals evolved as shown by a range of experimentaland natural studies [Handy, 1994b; Ji et al., 2000; Rybackiand Dresen, 2004; Stünitz and Fitz Gerald, 1993]. However,the aforementioned studies all neglect a possible roleof the melt on the deformation of the polyphase rocks.[4] The aim of this paper is to show, through detailedmicrostructural study and thermodynamical modeling, thecontribution of interstitial melt to the rheology of progressivelydeformed granites under midcrustal conditions. Weuse natural examples of a sequence of granite mylonites todocument the melt enhanced rheological inversion of disproportionatelystronger quartz compared to the weakfeldspars in midcrustal rocks. This work also shows thatwith increasing melt fraction the deformation mechanismsof feldspars varies from the grain boundary sliding accommodateddislocation creep to granular flow while the quartzis only deformed in the dislocation creep field.2. Geological Setting[5] The study area located in the central part of theBohemian Massif in the Czech Republic is known for theextreme deformation of porphyritic granites in a crustalscaleshear zone [Synek and Oliverová, 1993]. The porphyriticgranite mylonites studied are of Cambro-Ordovicianprotolith age and come from an orthogneiss-bearing, midcrustalunit that overlies kyanite micaschists in the west(Figure 1). In the east a similar rock assemblage occurs inan equivalent structural unit, although these orthogneissbodies exist within kyanite-sillimanite bearing micaschists.According to Schulmann et al. [2005] both units record aCarboniferous tectono-metamorphic history between 340and 325 Ma. The midcrustal unit is overlain by kyanitebearing migmatites and granulites of the orogenic lowercrustal unit that contain eclogite lenses with estimatedminimum pressures of 18–19 kbar and temperatures of800–900°C [Medaris et al., 1998]. The P-T conditions ofthe kyanite bearing micaschists in the footwall of themylonitic orthogneiss sheet have experienced temperaturesof 620–710°C at pressures of 6.5–9.5 kbar [Kachlík, 1999]similarly to 8–9 kbar and 610–660°C estimated in themicaschists in the eastern part of the studied area [Pitra andGuiraud, 1996].[6] The geological structure of the studied rocks wasdescribed by Synek and Oliverová [1993] who interpretedthe middle crustal orthogneiss-bearing unit and the overlyingorogenic lower crustal unit as a crustal nappe stackresulting from a Carboniferous deformation. On the basis ofthe structural position of the midcrustal orthogneiss and theregional metamorphic field gradient, the PT conditions ofmetamorphism and deformation of the studied rocks areestimated to be between 9 and 18 kbar and 650–850°C.More precise P-T estimations were not established becauseof the lack of pressure and temperature sensitive mineralassemblages that are necessary for standard thermobarometricmethods.3. Shape Analysis of Quartz and Feldspars[7] Undeformed porphyritic granitoids may serve as anexcellent example of multiphase mixtures of originallyspherical (ellipsoidal) clasts with constant phase fractions.When these rocks are subjected to deformation, mineralgrains reach different strain intensities, which can be easilyquantified using standard finite strain techniques [Ramsayand Huber, 1983]. Measurement of the shapes of naturallydeformed minerals in originally coarse-grained and porphyriticgranitoids may thus help to track the viscous behaviorof individual phases for different bulk strain intensities2of20296

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Figure 2. Macroscopic samples of deformed metagranite divided in three types according to thedeformation intensity and the macroscopic appearance. X, Y, and Z refer to the axes of the finite strainellipsoid. (a) Sample S1 is a weakly deformed metagranite with large centimeter-sized grains of quartzand feldspar representing Type I rock. (b, c) Sample M1 is an augen orthogneiss corresponding to Type IIrocks and intermediate strain intensity (equivalent to samples T1, T2, M2, and R4), (d, e) Sample V1 is abanded mylonite corresponding to Type III rock and the highest intensity of deformation (equivalent toother highly strained samples R5, R3, V2, and H1).[Treagus, 2002]. This allows constraining the degree ofstrain partitioning in rocks and the viscosity ratios betweenthe individual mineral phases.[8] The shape analysis of the feldspars and the quartzpolycrystalline aggregates was carried out on 17 sectionscut both perpendicular to the foliation and parallel to thestretching lineation and perpendicular to both the foliationand the lineation, i.e., parallel to the XZ and YZ sections ofthe finite strain ellipsoid (Figure 2). The K-feldspar cannotbe distinguished from plagioclase in highly deformed macroscopicsamples and therefore both minerals were groupedtogether for finite strain measurements. All studied samplesare composed on average of 60–70% feldspars, 35–25%quartz and up to 10% of biotite and muscovite. The almostconstant mineral composition for highly variable strainintensities and constant bulk rock chemistry shows the lackof chemical variations with strain (Table 1).[9] In our study we classified three major types ofdeformed orthogneiss according to the deformation intensitiesat the macroscopic scale: Type I is represented byweakly deformed metagranite (Figure 2a); Type II correspondsto augen orthogneiss with quartz porphyroclast(Figures 2b and 2c); Type III is a banded mylonite orthogneiss(Figures 2d and 2e).[10] Mineral shape data are plotted into a Flinn diagram[Flinn, 1965] (Figure 3). The analyses of feldspar polycrystallineaggregates and quartz of the individual samples areconnected by tie lines with the vertical ellipse representingthe bulk strain value of the whole rock. An importantfeature of all the studied samples is that feldspars showhigher strain intensities than quartz for any bulk strain(Figures 3a and 3b). In several samples of the so-calledDoubravčany pencil gneiss, the strain intensities cannot beidentified because the stretching of feldspar and quartzlayers exceeds the length of the samples (Figure 2d). Thestrain symmetry, represented by the K values of Flinn[1965], varies from prolate to oblate shapes (K = 2.7 to0.3). For quartz with prolate shapes, the correspondingfeldspar strain symmetry is close to plane strain, while foroblate quartz, the feldspars show the same shape or slightlymore oblate shapes (Figure 3b). Strain intensities areexpressed using Ramsay’s D value, which is an alternativeexpression of viscosity ratio values of Gay [1968a, 1968b] alsoused by Schulmann et al. [1996]. The inspection of the diagramin Figure 3b shows that the highest strain ratios are achievedbetween D fel /D qtz for Type III orthogneiss and some Type IIorthogneiss samples marked by high bulk strain intensities. Therelatively small ratio between D fel /D qtz suggests similar yield-3of20297

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Table 1. Bulk Rock Chemistry of Midcrustal OrthogneissesSample S1 M1 V1 D1SiO 2 72, 9 70, 9 73, 0 70, 2TiO 2 0, 3 0, 3 0, 2 0, 4Al 2 O 3 13, 3 14, 7 13, 7 14, 7Fe 2 O 3 1, 9 1, 4 1, 8 2, 6FeO 0, 9 1, 0 0, 7 0, 7MnO 0,0 0,1 0,0 0,0MgO 0,5 0,6 0,4 0,8CaO 1, 1 1, 4 0, 8 1, 0Na 2 O 2,3 2,6 2,5 1,8K 2 O 4,5 4,7 4,7 4,6P 2 O 5 0, 2 0, 2 0, 2 0, 2H 2 O- 0, 2 0, 3 0, 3 0, 2H 2 O + 1, 0 1, 2 1, 0 2, 1CO 2 0, 2 0, 2 0, 4 0, 2Total 99, 5 99, 5 99, 7 99, 6ing of both mineral phases for low bulk strains for severalsamples of weakly deformed Type II orthogneiss.4. Microstructure Development of DeformedMetagranitoids[11] The microstructural investigations covered the qualitativedescription of the rock and mineral structure usingoptical microscope and scanning electron microscope imaging.The microprobe work complements the scanning electronmicroscope (SEM) study to identify compositional variationsof the recrystallized feldspars (Tables 2a, 2b, and 2c).4.1. Type I: Metagranite and Weakly DeformedOrthogneiss[12] The metagranite samples show K-feldspar phenocrystsup to 10 cm in size surrounded by quartz blebs.Elongated plagioclase polycrystalline aggregates range from1 to 3 cm in size (Figure 2a). Light and SEM microscopiesshow that the plagioclase forming recrystallized aggregates(labeled Pl1 in Figure 4) corresponds to oligoclase An 15 .The Pl1 plagioclase has an average grain size of 5070 mm, straight boundaries and subequant shapes (Figure 4and Table 2c). The quartz blebs consist of large grains (400mm in size) with irregular and highly serrated boundaries.Biotite and muscovite are present as elongated recrystallizedaggregates. The plagioclase grains within alkaline feldsparphenocrysts are labeled Pl1b here. These plagioclases are 10to 50 mm in size, correspond to albite An 1–8 and are usuallyarranged into chains mostly parallel to (100) or locallyalong (010) planes of the host K-feldspar (Figures 4a and4b and Table 2c). The K-feldspars themselves show subgrainsof 80 mm in size. Subgrain boundaries are straightand meet in triple point junctions. At a smaller scale, thefeldspars show kinked domains along which the Pl1b chainsare rotated (Figure 4a). These kinked domains show internalstrain marked by irregular undulatory extinction. Microprobeanalyses show that K-feldspar cores exhibit ratherconstant composition revealing approximately 10 mol % ofalbite (Table 2a). Only around enclosed Pl1b chains thealkali feldspar shows a rapid decrease of the albite component(Figure 5).4.2. Type II: Orthogneiss With Quartz Augens[13] This rock is characterized by the presence of isolatedand elongated quartz augens surrounded by highly elongatedK-feldspar and plagioclase polycrystalline aggregates formingalmost monomineralic layers (Figure 6). Biotite andmuscovite are forming elongated monomineralic aggregateslocated at boundaries between feldspars and quartz orwithin plagioclase layers. Quartz augens are composed ofgrains 150–1000 mm in size with serrated boundaries.Plagioclase (Pl1) polycrystalline aggregates are composedof oligoclase An 10 – 20 subequant grains (50–150 mm inFigure 3. (a) Flinn diagram after Ramsay and Huber [1983] shows the shapes of deformation ellipsoidsrepresented as projections of X/Y axial ratios on the ordinate and Y/Z ratios on the abscissa in the graph.This diagram allows to visualize the shapes of the strain ellipsoid and the intensity of deformation in the2-D diagram. (b) Diagram showing strain intensity expressed as the D parameter on the abscissa againstshape K parameter on the ordinate; K = (X/Y-1)/(Y/Z-1) varies from 0 for oblate shapes, to 1 for planestrain shapes, and to infinity for prolate shapes, D = (((X/Y-1) 2 + ((Y/Z-1) 2 ) 1/2 . The values are obtainedby measuring 30 ellipses (harmonic sum) from XZ and YZ sections (principal planes of finite strainellipsoid). Squares show quartz, and circles show undifferentiated feldspar. Rock types are determinedaccording to the bulk strain intensity and macroscopic appearance (Figure 2) from weakly deformedmetagranite S1; intermediate strained augen orthogneiss samples T1, T2, M2, M1, and R4; and highlystrained banded orthogneiss samples H1, T5, V1, and V2.4of20298

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Table 2a. K-Feldspar Compositions aSample S1 S1 M1 M1 V1 V1Analysis 50 48 71 79 92 104SiO 2 65, 27 64, 07 64, 09 65, 05 64, 63 64, 96Al 2 O 3 18, 69 18, 42 18, 22 18, 56 18, 73 18, 71CaO 0, 00 0, 03 0, 04 0, 00 0, 03 0, 04Na 2 O 1,20 0,49 0,29 1,37 1,10 0,59K 2 O 15, 87 16, 46 16, 86 15, 52 15, 90 16, 48Total 101, 04 99, 49 99, 50 100, 54 100, 40 100, 82Si 2, 986 2, 985 2, 990 2, 988 2, 978 2, 984Al 1, 008 1, 011 1, 002 1, 005 1, 017 1, 013Ca 0, 000 0, 002 0, 002 0, 000 0, 002 0, 002Na 0, 106 0, 044 0, 026 0, 122 0, 098 0, 053K 0, 926 0, 978 1, 004 0, 910 0, 935 0, 966Total 5, 026 5, 020 5, 024 5, 024 5, 030 5, 017X Or 89, 69 95, 53 97, 27 88, 17 90, 36 94, 66X Ab 10, 31 4, 32 2, 54 11, 83 9, 50 5, 15X An 0, 00 0, 15 0, 19 0, 00 0, 14 0, 19a Structural formulae calculated on the basis of 8(O).size) with straight boundaries meeting at triple point junctions(Figures 5 and 6c and Table 2c). Plagioclase polycrystallineaggregates often contain biotite flakes parallel tothe foliation or interstitial quartz and new plagioclaseAn 1–10 (Pl2). K-feldspar polycrystalline aggregates arecomposed of slightly elongate to subequant grains withstraight boundaries forming a well developed triple pointnetwork lined by narrow films of pure albite An 1–10 (Pl2).Disintegration of the K-feldspar layers into elongate aggregatessurrounded by fine-grained Pl1 matrix was observedin some samples. This process is connected to the developmentof bands filled with recrystallized Pl1 grains up to150 mm wide oblique with respect to the long axis of thefeldspar polycrystalline aggregates (Figure 6b). Finally,elongated thin aggregates of Pl1 are smeared out in theK-feldspar rich matrix. Quartz also forms weakly elongatedpolymineralic aggregates characterized by highlyirregular boundaries with the surrounding feldspar matrix.Common feature are the lobes of quartz and cusps of theK-feldspar pointing in the direction perpendicular to thelong face of the aggregate and to the macroscopic foliation.4.3. Type III: Banded Mylonitic Orthogneiss andUltramylonite[14] Banded mylonitic orthogneiss is marked by thinquartz ribbons less than 1000 mm wide. They are surroundedby polymineralic layers of plagioclase, quartz andTable 2b. Isolated Plagioclase Pl1b and Pl2 Compositions aSample S1 T1 T1 M1 V1Analysis 52 119 115 77 105SiO 2 68, 72 69, 06 67, 94 68, 67 67, 69Al 2 O 3 19, 70 19, 74 20, 19 19, 78 20, 69CaO 0, 31 0, 10 0, 78 0, 22 1, 28Na 2 O 11, 84 12, 14 11, 27 11, 77 11, 19K 2 O 0,05 0,12 0,12 0,54 0,23Total 100, 62 101, 19 100, 30 101, 01 101, 09Si 2, 986 2, 986 2, 963 2, 980 2, 938Al 1, 009 1, 006 1, 038 1, 012 1, 058Ca 0, 014 0, 005 0, 037 0, 010 0, 060Na 0, 998 1, 018 0, 953 0, 991 0, 942K 0, 003 0, 007 0, 007 0, 030 0, 013Total 5, 010 5, 022 4, 998 5, 023 5, 010X Or 0, 28 0, 64 0, 67 2, 90 1, 26X Ab 98, 30 98, 91 95, 67 96, 11 92, 87X An 1, 42 0, 45 3, 66 0, 99 5, 87a Structural formulae calculated on the basis of 8(O).Table 2c. Large or Interconnected Plagioclase: Pl1 Compositions aSample S1 T1 M1 M1 V1Analysis 43 114 67 91 109SiO 2 64, 60 64, 70 65, 62 63, 75 65, 55Al 2 O 3 22, 72 22, 84 21, 43 23, 06 22, 22CaO 3, 48 3, 76 2, 36 4, 38 3, 07Na 2 O 9, 88 10, 03 11, 03 9, 48 10, 23K 2 O 0,23 0,13 0,10 0,27 0,22Total 100, 91 101, 46 100, 56 100, 97 101, 30Si 2, 827 2, 819 2, 878 2, 797 2, 854Al 1, 172 1, 173 1, 108 1, 192 1, 140Ca 0, 163 0, 176 0, 111 0, 206 0, 143Na 0, 838 0, 847 0, 938 0, 806 0, 864K 0, 013 0, 007 0, 006 0, 015 0, 012Total 5, 013 5, 022 5, 040 5, 017 5, 013X Or 1, 26 0, 70 0, 53 1, 47 1, 20X Ab 82, 65 82, 26 88, 95 78, 49 84, 75X An 16, 09 17, 04 10, 52 20, 04 14, 05a Structural formulae calculated on the basis of 8(O).K-feldspar (Figure 7b). Micas are generally dispersed in thematrix or form narrow layers parallel to the foliation(Figures 7c and 7d). Disintegrated relics of plagioclaseaggregates are composed of Pl1 grains An 14 – 20 (Figure 5and Table 2c) and of K-feldspar grains 50–100 mm in size(rarely 200 mm) and of irregular shapes. The interfacialboundaries with K-feldspar are generally straight and commonlylined by rims of Pl2 (An 4–12 ) or thin layers of quartz.Locally, relics of the K-feldspar layers that are a few grainswide (250 mm) occur with straight mutual boundaries linedby Pl2 films. The quartz commonly occurs in the form ofisolated grains with serrated boundaries and cusps pointingperpendicular and parallel to the macroscopic foliation. However,the most common are few hundred microns wide quartzribbons composed by elongate, 200–300 mm wide and up to1000 mm long grains with highly serrated boundaries.[15] A rather exceptional kind of Type III microstructureis represented by the exceptionally coarse-grained, myloniticorthogneiss sample B2. This rock type is characterizedby polyminerallic layers with poorly defined boundaries anda large average grain size of plagioclase and K-feldsparranging from 160 to 200 mm. The Pl1 grains in relictplagioclase aggregates show sutured boundaries and widePl2 rims while the interstitial Pl2 grains in the K-feldsparrich aggregates form wide films and cuspate pools. Theinterstitial quartz (100 mm in size) occurs in the form ofserrated grains in triple point junctions of K-feldspar andPl1 relict aggregates.4.4. Topology of Interstitial Phases in Plagioclase andK-Feldspar Aggregates[16] The most common interstitial phase at the grain scaleare thin films of Pl2 and quartz up to 50 mm long and 10 mmwide located along the mutual plagioclase boundaries, oftenat high angle to the foliation or in the form of cuspate poolsat triple point junctions. The mutual boundaries of Pl1aggregate grains are often lined by narrow rims of newplagioclase An 1–10 (Pl2). In Type II orthogneiss samplesshowing strongly elongated recrystallized K-feldspar andPl1 grains (samples T1, T2), the Pl2 films exhibit preferredorientation parallel to the direction of maximum stretching,i.e., along the crystal faces parallel to the foliation (Figure 8).In the M1, M2 samples of the orthogneiss Type II, the shapepreferred orientation of the recrystallized feldspar is low,5of20299

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406[18] Another important feature is the development of intragranular,wedge shaped fractures that are predominantlydeveloped in elongated feldspar grains with stronger SPOin samples T1 and T2 (Figure 8). These gently curvedfractures often originate at the foliation-parallel grain boundariesor triple point junctions, terminate in the interior of thegrains and are oriented at low angles (15–30°) to themaximum stretching. In Type III orthogneiss samples (e.g.,sample B2) the intragranular fractures occur as well beingoriented at high angle to the principal stretching direction.Figure 4. Type I microstructure (S1 sample). (a) Type Imetagranite marked by large alkaline feldspar clasts decomposedinto albite chains and K-feldspar and (b) scanningelectron microscope (SEM) image of alkaline feldspardecomposition. Rectangle indicates the position of the detailedSEM image. (c) The detail of the SEM image showing theshapes of new bulbous albite grains. (d) CorrespondingArcView Geographical Information Systems (GIS) image usedfor the quantitative mirostructural analysis. Scanning electronmicroscope images obtained in backscattered electron mode(Camscan microscope, Institute of Petrology and StructuralGeology Prague). K-feldspar is represented in light gray,biotite is represented as white laths, quartz and plagioclase areshown in dark gray, and rectangular white mica laths areshown in light gray. In ArcView GIS images the K-feldspar isrepresented in white, white mica and biotite are represented inblack, plagioclase is represented in light gray, and quartz isrepresented in dark gray. The ArcView GeographicalInformation System was used as an ideal environment fordigitizing mineral shapes [Lexa et al., 2005]. Mineralabbreviations correspond to those of Kretz [1983].which corroborates the development of thin albite-quartzfilms along the feldspar faces oriented along two maximawith respect to the macroscopic layering (Figure 8). Themean orientation of the Pl2 and quartz seams form an angleof 10–15° with respect to the foliation trace. Interstitialconvex quartz grains often occur at triple point junctionsbut locally line the feldspar boundaries as narrow films(Figures 6c and 6d).[17] The presence of interstitial phases is the most pronouncedin Type III orthogneiss (samples R3, V1 and B2 inparticular) where the recrystallized feldspar grains exhibitno shape preferred orientation. Here, wider cuspate-lobatepools and narrow films of Pl2 and quartz occur at a highangle to the macroscopic foliation. The B2 sample shows anextreme orientation of Pl2 seams which form an angle of upto 70° with the foliation trace.5. Quantitative Textural Analysis[19] The quantitative analysis of grain shapes and boundarieswas carried out in an ArcView GIS environment.Examples of analyzed samples are shown in Figures 4d,6d, and 7d. The statistical parameters involving grains sizedistributions, shape preferred orientation, degree of grainelongation and grain contact frequencies were performedusing the MATLAB PolyLX toolbox [Lexa et al., 2005] inorder to quantify the above described microstructural sequenceof rock types. The grain size distribution wasevaluated as an important parameter in deformed rocksbecause of its sensitivity to stress and temperature [Schmidet al., 1999]. The correct determination of grain size isessential in polyphase systems, where stress and strain ratepartitioning are expected [Handy, 1990]. Another importantparameter is the shape of recrystallized grains, which isstrongly dependent on the type of deformation mechanismsand grain growth history [Boullier and Guéguen, 1975;Schmid et al., 1987]. Grain contact distribution in the rockyields information about the evolution of spatial distributionof the different minerals in rocks with deformation ormelting [Lexa et al., 2005; Hasalová etal., 2008]. This isexpressed by the grain contact frequency (GCF) that statisticallyquantifies the deviation of the grain contact distributionfrom random [Kretz, 1969]. The aggregate distribution,marked by the dominant presence of like-like contacts,represents one end-member of the spatial distribution ofgrain boundaries that is resulting from the solid statedifferentiation process [McLellan, 1983]. The regular distributionis characterized by the predominance of unlikecontacts in rock and is considered as a second end-memberFigure 5. Ternary diagrams showing the composition offeldspars in different samples. Pl1 corresponds to isolatedplagioclase aggregate (see text for explanation).6of20300

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406examined for the interstitial and aggregate positions separately.The most striking feature is a progressive increase ofthe aggregate plagioclase grain size from 60 mm (medianvalue) in metagranite to 90 mm in augen orthogneiss to110 mm in banded orthogneiss, and to 170 mm forsample B2 (Figure 9a). We also observe a systematicallyhigher grain size of aggregate grains compared to theinterstitial plagioclase grains for all rock types. In addition,the median value of the aggregate plagioclase grainsis shifted toward the third quartile value, while in theinterstitial grains it is closer to the first quartile value.[22] In order to obtain information about the crystal nucleationand growth, the crystal size distribution (CSD) methodwas applied in the studied samples. The theory of CSD is awell-established technique in metallurgy, ceramics and chemicalengineering to reveal information about nucleation,growth rates and growth times of crystals [Randolph andLarson, 1971]. CSD in many metamorphic and igneous rocksshow a log linear relationship between the grain size L and thepopulation density N according to the equationN ¼ N 0 e L=GtFigure 6. Type II microstructure. (a) Micrograph ofsample M2 Type II microstructure characterized by largequartz ribbons and fine-grained feldspar matrix, (b) SEMimage of the M2 sample showing typical monomineralliclayering and plagioclase aggregate bridges, and (c) detailedSEM image of the T1 sample showing K-feldspar andplagioclase polycrystalline aggregate microstructure. Thecharacteristic feature is the high elongation of feldspargrains lined with Pl2 films, (d) corresponding ArcView GISdrawing of Type II microstructure of the T1 sample. TheSEM and ArcView GIS colors are the same as in Figure 4.where N 0 and Gt are constants and may be related to thenucleation density and growth rate of the crystals [Higgins,1998]. This technique was previously successfully appliedto the metamorphic rocks by Cashman and Ferry [1988],Hasalová etal.[2008], and Lexa et al. [2005]. CSD plots ofof the spatial distribution of the grain boundaries thatdevelops mostly because of solid state annealing [Flinn,1969], mechanical mixing [Kruse and Stünitz, 1999] orcrystallization of the interstitial melt [Dallain et al., 1999;Hasalová etal., 2008].5.1. Crystal Size Distribution[20] The grain size distributions of recrystallized quartz,K-feldspar and of plagioclase have been determined on thebasis of more than 500 measured grains per thin section foreach sample (except of about 200 measurements in coarsegrainedsample B2). Because of the ubiquitous lognormaldistribution of the grain populations, the median value wasconsidered to be the most reliable statistical value and thespread of the grain size distribution was evaluated as adifference between the third and first quartile instead of thestandard deviation.[21] The K-feldspar shows a similar grain size distributionfor all rock types and is characterized by the mediangrain size ranging from 60 to 180 mm (Figure 9a). Exceptsample B2 with exceptionally large grain size, the grain sizespread does not follow any systematic pattern. However, theinterstitial quartz shows an increasing median grain sizevalue from 20 to 50 mm in the Type II orthogneiss to 80–100 mm in Type III orthogneiss. The plagioclase wasFigure 7. Type III microstructure of banded orthogneiss.(a) Micrograph of sample D1 showing layers of quartzalternating with mixed layers of feldspars and quartz, (b)SEM image of the sample V1 showing quartz and feldsparmixing and poor definition of layer boundaries, (c) detailedSEM image of K-feldspar and plagioclase mixing in thematrix, and (d) corresponding ArcView GIS drawing. TheSEM and ArcView GIS colors are the same as in Figure 4.7of20301

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Figure 8. (a) Detail from the backscattered scanning electron image of sample T1 (location of Figure 8ais shown in SEM image in Figure 6c). The image shows the distribution of albite seams along boundariesof K-feldspar grains, character of compositional zoning of the plagioclase in the top left, and intragranularfractures filled with albite (I.F.). (b) Rose diagrams show preferred orientation of 100 interstitial quartzand Pl2 seams in K-feldspar aggregate. The mean direction and standard circular deviation are shown inthe bottom of each diagram. The thick horizontal line represents the orientation of lineation (X), and thethin line shows the orientation of the mean direction.all the samples constructed according to the method byPeterson [1996] exhibit linear correlations between thelogarithm of the population density (i.e., the number ofcrystals per size per volume) and the crystal size (Figure 9b).Applying the theory of CSD, such distributions could beparameterized by the zero size intercept N 0 (nucleationdensity) and slope Gt (growth rate multiplied by time).These two parameters are plotted in a N 0 -Gt diagram[Lexa et al., 2005] where the samples form a distincttrend. Trends in the grain size distributions using CSDmethod are visualized in Figure 9b. The crystal sizedistribution plot of the plagioclase aggregates is the mostpronounced and shows a systematic decrease of N 0 and anincrease of the G t values with increasing degree ofdeformation i.e., from Type I to Type III rocks.5.2. Grain Shapes and Shape Preferred Orientation(SPO)[23] The aspect ratio median value of plagioclase andK-feldspar varies between 1.5 to 3.1 and no systematicpattern related to the type of rocks and the degree ofdeformation is obvious (Figure 10). However, the SPO ofplagioclase and K-feldspar in most of Type II and IIIorthogneiss samples is higher compared to the Type Isample with exceptionally high SPO for samples T1 andT2. There is a difference between the aggregate plagioclase(Pl1) and the interstitial albite (Pl2) marked by a systematicallyhigher SPO for the former compared to the latter.5.3. Grain Contact Frequencies (GCF) and GrainBoundary Preferred Orientation (GBPO)[24] The combination of the GCF analysis with thestudies of the preferred orientation of the like-like (like–like contacts = boundaries of minerals of the same species)and unlike grain boundaries (GBPO) yields importantinformation about the organization of the grain boundarieswith respect to the deformation processes [Lexa et al.,2005]. So far the degree of deviation of the grain boundarydistributions from the random distribution has been evaluatedby plotting the observed/expected ratio of the like–like8of20302

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Figure 9. (a) Grain size statistics for the studied samples presented in box-and-whiskers diagrams.Horizontal axis corresponds to the ferret diameter of grain size in micrometers, and the thick barrepresents the median of grain size distribution. Vertical axis shows samples arranged according to thedegree of deformation from least deformed S1 to most intensely deformed sample V1. B2 represents anexceptionally coarse-grained mylonitic banded orthogneiss. For each sample the plagioclase grain sizestatistics for interstitial phases (I = Pl2 grains), and for recrystallized grains forming aggregates (A = Pl1grains) are shown. The number of grains is indicated. Grain size distributions of isolated quartz in thematrix and quartz grains forming centimetric ribbons or augen are also differentiated. This diagram showsprogressive coarsening of both Pl1 (aggregate) and Pl2 (interstitial) grains. (b) Plot of crystal sizedistribution (CSD) for plagioclases (I = Pl2 interstitial grain, A = Pl1 recrystallized grains in aggregates).N 0 ln(mm 4 ) values on vertical axis represent density of grains per volume, and dimensionless Gt valuesreflect the grain size frequency distribution. This diagrams show decrease of N 0 values and increase of Gtvalues which in classical CSD plots represent decreasing values of the intercept of the CSD curve withvertical axis associated with decreasing slope [Higgins, 1998]. In the CSD theory this evolution meansdecreasing nucleation rate and increasing growth rate contribution to the shape of the grain size frequencyhistogram [Lexa et al., 2005].contacts of the two major minerals against each other [e.g.,McLellan, 1983]. Lexa et al. [2005] proposed a newdiagram where the c valuec ¼ Observed pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiExpectedExpectedis plotted against the ratio of orientation tensor eigenvaluesthat represent the degree of GBPO. If the solidified melt isidentified in a rock the GBPO may yield information aboutthe types of channel networks as defined by Sawyer [2001]and quantified by Hasalová etal.[2008] and Závada et al.[2007].[25] In this work we use a method proposed by Lexa et al.[2005] where the grain contact frequencies are used toassess the character of the spatial distributions of grainboundaries. Figure 11a shows that the plagioclase like-likecontacts for the Type I metagranite plots in an intermediatepart of the diagram and is characterized by a rather smallaggregate distribution. This is due to the presence of the twoplagioclase populations resulting from the recrystallizationof the large plagioclase crystals (high aggregate distribution)and from the disintegration of the large alkalinefeldspar crystals into a mixture of albite and K-feldspar(Figure 4). With increasing deformation a clear evolution ofthe grain contact frequency toward strongly aggregateddistribution for weakly deformed samples of the Type IImicrostructure can be observed. This is probably due to thecoalescence of feldspar and plagioclase layers. At very highstrain intensities the plagioclase grain contact frequenciesevolve toward the random or regular types of distribution inType III microstructure. This type of evolution indicates analmost perfect mixing of the plagioclase with other mineralphases. The degree of GBPO of the like-like boundaries forplagioclase does not evolve with the above described trendbut remains rather constant for any degree of the finitestrain. The K-feldspar grain contact frequency shows asimilar behavior to the plagioclase but for some samples arather strong GBPO coupled with a decreasing grain contactfrequency was observed. The evolution of the unlikeplagioclase-K-feldspar grain boundaries exactly mirrorsthe evolution of the plagioclase like-like contact frequencybehavior (Figure 11b). For the observed constant grain sizeit is a geometrical necessity that an increase of the like-likecontacts causes a corresponding decrease of the unlikecontacts of feldspars.6. Crystallographic Preferred Orientation[26] The lattice-preferred orientation (LPO) of quartz,plagioclase and K-feldspar was determined using the electronbackscatter diffraction (EBSD) technique [Bascou et9of20303

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Figure 10. Plot of the grain shape preferred orientations (SPO)of plagioclase and K-feldspar in studied samples. The resultsare summarized by a box-type plot of axial ratios versuseigenvalue ratios of bulk shape preferred orientation forindividual phases. Individual boxes show median and firstand third quartile values. The whiskers represent a statisticalestimate of the range of data, while outliers are not shown.Vertical axis characterizes the shape of grains, while horizontalaxis represents the area-weighted degree of preferred orientation.Plagioclase (I = interstitial Pl2 grains, A= recrystallized Pl1 grainsforming aggregates). This diagram shows generally low aspectratios of plagioclase grains for all samples and exceptionallystrong SPO of some samples of Type II and III microstructures.al., 2001]. The lattice preferred orientation of quartz ispresented in the pole figures of the crystallographic directions< c > and < a > (Figure 12). In the case of plagioclaseand K-feldspar, the list of operative slip systems [Kruse etal., 2001; Tullis, 1983] has been used and measured data hasbeen plotted in the pole figures of these crystallographicplanes and directions. Pole figures of principal slip directionsand slip planes showing the best coincidence with themain axes of the finite strain ellipsoid of every sample arepresented (Figure 13). In the assessment of the maximaposition in the pole figures, it has been taken into accountthat samples M1, V1 reveal prolate shapes of the strainellipsoid. Hence the maxima of poles in the slip planes mightoccur along a girdle perpendicular to the slip direction.6.1. Quartz Augens and Ribbons CrystallographicPreferred Orientation (CPO)[27] The quartz c axis fabric for Type I microstructure(sample S1) shows a strong central maximum and weakersubmaxima close to the periphery of the diagram suggestingan ill-defined oblique cross girdle pattern with openingangles of around 45° (Figure 12). The fabric is a result ofthe dominant prism < a > and subordinate basal < a > slipsystems and a plane strain noncoaxial deformation. Type IImicrostructures (samples T2 and M1) show strong, central caxis maxima. There is also a tendency to form a singlegirdle of c axes distribution oriented oblique to the foliation(T2). This c axis pattern is commonly interpreted as a resultof a prism < a > slip activity and a noncoaxial deformation[Schmid and Casey, 1986]. The Type III microstructure(samples R3 and V1) is characterized by maxima locatedeither at the periphery of the diagram or in an intermediateposition. The maxima are organized either along singlegirdles oblique to the foliation (sample R3) or symmetricallyin case of sample V1. These c axis patterns developedfrom combined activity of the basal < a > and rhomb < a + c> slip systems with a minor contribution of the prism < a >slip during the noncoaxial deformation.6.2. Plagioclase and K-Feldspar CPO[28] Plagioclase CPO shows a weakening and an increasedactivity of secondary and tentative slip systemswith microstructural evolution from Type I to Type IIImicrostructures (Figure 13). A frequently described slipsystem (010)[100] [Kruhl, 1996; Martelat et al., 1999;Schulmann et al., 1996] has been observed only in the TypeI microstructures (S1) and shows an asymmetry of theposition of the maxima with respect to the foliation. Theplagioclase from the Type II and III microstructures showedless common slip systems that are supposed to be secondaryor tentative [Kruse et al., 2001]. In the Type II microstructure(T2 and M1), the CPO shows active slip systems withdissociated Burgers vectors [Montardi and Mainprice,1987; Olsen and Kohlstedt, 1985] namely 1/2[112](111)and 1/2[111](101) in the sample T2, and 1/2[112](110) and1/2[112](201) in the sample M1. Brief inspection ofboth pairs of slip systems shows that each pair occupies aclose position in the plagioclase crystal [Kruse et al., 2001,Figure 2] and it is very likely that they operated simultaneously.The Type III microstructures show very weak crystallographicpreferred orientation of the plagioclase and sampleR3 again reveals an activity on the 1/2[112](110) slip systemswith dissociated Burgers vector (Figure 13).[29] The K-feldspar phenocryst and albite neoblasts ofType I orthogneiss show the same crystallographic orientations(sample S1, Figure 13) that do not coincide with anyknown slip systems. Albite neoblasts originate along thestrings that are parallel to (100) planes. In the Types II andIII microstructures, the CPO of K-feldspar recrystallizedgrains reveal similar pattern indicative of the slip directionparallel to the 1/2[110] in all samples that operate on the(001), (111) slip planes (Figure 13). A less pronounced slipsystem [101](101) has been recognized in the augen gneisssample T2 from the Type II microstructure.7. Petrological Modeling[30] The structural position and shapes of interstitialphases, the character of mineral zoning and the compositionalvariations of the plagioclase (Pl2) in both Type II andIII microstructures indicate the presence of some kind offluid during the deformation [Fitz Gerald and Stünitz, 1993;Sawyer, 2001; Hasalová etal., 2008; Závada et al., 2007].Therefore, the pseudosection modeling was performed inorder to examine whether the deformation process in thestudied orthogneisses occurred in the presence of a melt oran aqueous fluid.10 of 20304

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Figure 11. Plot of the grain boundary frequencies. Plots of the deviations from random spatialdistribution versus the degree of grain boundary preferred orientation. The value of the deviation fromrandom spatial distribution is obtained by the contact frequency method [Kretz, 1969], and the lengthweighteddegree of preferred orientation is estimated as the ratio of eigenvalues of the bulk matrix ofinertia. The plot of the like-like and unlike boundaries is separated. (a) The diagram of feldspars like-likeboundaries shows a trend from the weak aggregate distribution of Type 1 microstructure, to the highlyaggregate distribution in Type II microstructures, and almost random distribution in highly mixed TypeIII microstructures. (b) The diagram of feldspar unlike boundaries mirrors the like-like evolutionarytrend.7.1. Modeling Method[31] The pseudosections were calculated in the system Na 2 O-CaO-K 2 O-FeO-MgO-Al 2 O 3 -SiO 2 -H 2 O (NCKFMASH). All theFe was treated as FeO and molar amounts of the consideredoxides were recalculated to 100 mol %. The calculationswere performed using THERMOCALC 3.25 [Powell et al.,1998] and the data set 5.5 [Holland and Powell, 1998]. Themixing models for the most solid solutions were takenfrom White et al. [2001] and the THERMOCALC documentation[Powell and Holland, 2004]. The feldspars areformulated using the model of Holland and Powell [2003],the paragonite-muscovite solution is after Coggon andHolland [2002]. The quartz, K-feldspar and plagioclaseare in excess and the pseudosections are contoured forliquid mole isopleths.7.2. Results[32] The first pseudosection is calculated with the amountof H 2 O, M(H 2 O) = 2.68 mol % in the whole rock composition,which is the amount of H 2 O tied in micas in theassemblage biotite-muscovite-plagioclase-K-feldspar-quartz(Bt-Ms-Pl-Kfs-Qtz) at 8 kbar and 600°C (Figure 14a). Thelimiting maximum pressure for the studied rock with theassemblage Bt-Ms-Pl-Kfs-Qtz is the appearance of garnet at8–9 kbar, the maximum temperature limit is the beginningof biotite dehydration melting marked by the appearance ofgarnet at temperatures of 680–700°C and the breakdown ofmuscovite at 680–700°C below 6 kbar that is not observed.The minimum pressure and temperature conditions cannotbe determined from the assemblage of the studied sample,therefore, we used the P-T path of Pitra and Guiraud [1996]that was determined in the neighboring metapelites. It ischaracterized by decompression from 8–9 kbar and 610–660°C to 4–5 kbar and 600–650°C. Path 1 (Figure 14a) ischaracterized first by heating followed by decompression.Melting in the assemblage Bt-Ms-Pl-Kfs-Qtz without externalH 2 O starts at c. 660°C and 8 kbars and the maximumamount of melt produced in the field of Bt-Ms-Pl-Kfs-Qtz-Liq is less then 1 mol %.[33] The evolution of assemblages that the rock compositionproduces with a varying amount of H 2 O(M(H 2 O) =0–4.78 mol %) on heating at 8 kbar is examined inFigure 14b. The evolution for the amount of H 2 O tied inmicas (=2.68 mol %) in a rock with the starting assemblageBt-Ms-Pl-Kfs-Qtz is indicated by path 1, and is equivalentto the isobaric heating path in Figure 14a. The assemblageBt-Ms-Pl-Kfs-Qtz starts to melt at c. 660°C, and producesless then 0.5% of melt at c. 685°C (path 1 in Figure 14b).The upper temperature limit is the appearance of garnet at c.690°C that is not observed in the studied rocks. The loweramount of H 2 O in the rock stabilizes the anhydrous phasessuch as the garnet or kyanite and higher amounts of H 2 Oindicates the presence of free aqueous fluid below 630°C.[34] If above the temperature of c. 630°C appears freeH 2 O in the rock that follows the path 1, the rock compositionis immediately drawn to the right side, and starts to11 of 20305

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406melt (for example path 2 in Figure 14a). This leucocraticmelt is generated by the water released by the dehydrationprocess in which water migrates upward into the melt-fertilerocks that are above the wet-solidus temperature [Scaillet etal., 1990; Thompson and Connolly, 1995]. The quantity ofproduced melt depends on the amount of available externalH 2 O. For example the total amount of H 2 O for path 2 isM(H 2 O) = 4.28 mol %, which is 1.60 mol % of H 2 O addedto 2.68 mol % H 2 O already present in the micas. Such H 2 Oaddition results in the production of 5 mol % of melt at660°C and 8 kbar.[35] The evolution during decompression was also studiedin P-M(H 2 O) sections at 660°C. Following a decompressionfrom 8 to 4.5 kbar, the rock with the original H 2 Otied in micas increases the amount of melt to 1 mol %. In arock that contains added H 2 O, the melt fraction increases by1–2.5%, for example for path 2 the total amount of melt at4.5 kbar is 7.5 mol % because of 1.6 mol % added H 2 Oonthe prograde path.8. Discussion[36] Here we discuss the anomalously weak K-feldsparand plagioclase compared to quartz in light of the strain dataand quantitative microstructural and textural analysis. Thequartz-feldspar rheology inversion is discussed as a result ofthe alkaline feldspar breakdown and the formation of a finegrainedfeldspar matrix. Finally, we propose a model ofmixing of the recrystallized plagioclase and K-feldspargrains with the silicate melt associated with grain boundarysliding controlled diffusion creep followed by a granularflow at higher melt proportions.8.1. Comparative Quartz and K-Feldspar Rheology[37] On the basis of the number of experimental data, thequartz can be considered to be significantly weaker comparedto the K-feldspar and plagioclase for a wide range oftemperatures [Jaoul et al., 1984; Kronenberg and Tullis,1984; Tullis, 1990]. This is supported by a number ofstudies of quartzo-feldspathic rocks deformed at mediumtohigh-temperature conditions [Gapais, 1989; Handy,1994a; Schulmann et al., 1996]. It is only at greenschistfacies conditions and in the presence of hydrous fluids,when feldspars become weaker than quartz via destabilizationand breakdown to mixture of retrograde fine-grainedreaction products as white mica, quartz (K-feldspar breakdown)and epidote, albite ± garnet (plagioclase breakdown)[Fitz Gerald and Stünitz, 1993; Handy, 1990; Stünitz andFitz Gerald, 1993]. Nevertheless, it is assumed that thestrength of feldspars is generally several orders of magnitudehigher than that of quartz [Handy, 1994a; Handy et al.,1999; Ranalli and Murphy, 1987].[38] The shapes of quartz and feldspar polycrystallineaggregates measured in this work clearly show that for12 of 20Figure 12. Electron backscatter diffraction in situ measurementson quartz ribbons. Pole diagrams showingcontoured crystallographic orientation (projected in lowerhemisphere equal area). Contoured at multiples of uniformdistribution (maximum 10 uniform). The foliation normal isN-S, and the stretching lineation is E-W. Black squarescorrespond to the pole of the mean orientation. The polefigures show evolution of active slip systems withincreasing deformation from activity of combined prism and basal < a > slip systems for weakly deformed Type Imicrostructures, via the activity of prism < a > slip systemin intermediate Type II microstructures to combined activityof basal < a > and rhomb < a + c > slip systems in highlystrained Type III microstructures.306

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Figure 13. Contoured pole figures of the most characteristic slip direction and slip plane of K-feldsparand plagioclase from the Type I to the Type III microstructures. Equal area projection, lower hemisphere.Contoured at interval 1.0 times of the uniform distribution. Foliation (full line) is horizontal, and lineationis in this plane in the E-W direction. N is the number of measured grains. Maximum densities (blacksquare in the pole figures) are marked below the pole figure. Plagioclase shows commonly reported slipsystem only for Type I microstructure and activity of secondary and tentative slip systems withdissociated Burgers vectors for Type II and III microstructures. The top right pole figures showdependency of CPO of the new exsolved albite grains Pl1b on the orientation of alkali feldspar host.K-feldspar reveals also slip systems with dissociated Burgers vectors for Types II and IIImicrostructures. See text for discussion.various deformation intensities the quartz aggregates areless deformed than the feldspar aggregates and the differenceof D values between feldspar and quartz aggregatesincreases with increasing bulk deformation. We suggest thatthe studied samples reveal a higher competency of quartzcompared to the feldspar polycrystalline aggregates for allexamined strain intensities at high-temperature conditions.The strength reversal between the quartz and feldspar inexperimentally deformed aplite was documented byDell’Angelo and Tullis [1996]. These authors showedthat at 700°C, the dispersed quartz grains in the apliteremain less deformed than the feldspar grains while athigher temperatures the interconnected quartz becomesweaker than feldspars.8.2. Structural Evolutionary Trends of PolyphaseMylonites[39] We discuss here a hypothesis of microstructuralevolutionary trend in which the Types I, II and III microstructuresrepresent different deformation stages along a13 of 20307

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Figure 14. (a) Pseudosection showing the evolution of assemblages that the rock composition produceswith varying amounts of H 2 O(M(H 2 O) = 0–3.88 mol %) on heating at 8 kbar. (b) The pseudosectioncalculated with the amount of H 2 O tied in micas, deduced from pseudosection (Figure 14a). Thepseudosection shows a prograde path between 7 and 8 kbar along which the melt appears above 640°Cinthe stability of Bt-Ms-Liq-Pl-Ksp-Qtz. For comparison, the contours of melt production in other fields arealso shown. Mineral abbreviations correspond to those of Kretz [1983].progressive deformation path. Theoretically, the loadbearingframework (LBF) minerals (LBF = interconnectedstrong phase) should shield a weak phase from viscousdeformation for weak phase fractions lower than 20%[Handy, 1990]. With increasing strain the weak phase startsto interconnect along localized shear bands, leading finallyto the development of a banded structure of alternatingmonomineralic layers [Jordan, 1988]. This corresponds toan evolution from high-viscosity contrast toward lowviscositycontrast interconnected weak layer (IWL) structures[Handy, 1994a].[40] The microstructure of the Type I orthogneisses showinterconnected recrystallized aggregates of plagioclase (representingthe weakest phase) surrounding strong clasts ofquartz and alkali feldspar. This kind of microstructurecorresponds to IWL microstructure with the deformationhighly localized into rheologically weak plagioclase and arelatively high volume (up to 60 vol. %) of strong fraction(Figure 15a). This observation is valid for the macroscopicscale, but different rheological behavior can be observed atthe scale of the individual feldspar clasts. The internalstructure of the alkali feldspar phenocrysts shows typicalcharacteristics for a LBF structure formed by K-feldspar andaggregates of recrystallized plagioclase representing anisolated weak phase [Eudier, 1962; Jordan, 1987; Tharp,1983]. A rather high volume of weak pockets (20–30%)indicates a low stability of the LBF structure for small strainintensities [Handy, 1994a]. This is consistent with theobserved onset of coalescence of plagioclase chains alongmicroshear bands in Figure 6b [Jordan, 1988].[41] The Type II microstructures are characterized by thecollapse of the internal LBF structure of alkali feldsparthrough the interconnection of recrystallized plagioclase andK-feldspar. This evolution is documented by the presence ofoblique ‘‘bridges’’ of recrystallized plagioclases crossingrecrystallized K-feldspar aggregates and connecting surroundingplagioclase matrix (Figure 6b). From a rheologicalpoint of view, the volume of the weak matrix increased from 30 to 60 vol. % through the addition of entirely recrystallizedplagioclase and K-feldspar from original phenocrystsof Type I rock to already recrystallized plagioclase ofType II orthogneiss (Figure 15b). This evolution leads tothe development of IWL structures at all scales marked bya low volume of strong phases (quartz) and a moderateviscosity contrast indicated by the elongated shapes ofquartz aggregates. An important feature of this stage is thecoalescence of K-feldspar and plagioclase leading to thedevelopment of almost monomineralic layers (Figure 6).[42] The Type III microstructure is characterized by thedestruction of feldspar monomineralic layering throughprogressive mixing of the feldspars and the subordinatesmall quartz grains forming a fine-grained matrix. The finalmicrostructure is represented by highly elongated quartzribbons surrounded by the homogeneously deformed feldspar-quartzmatrix. This type of rock microstructure correspondsto the IWL structures and is marked by a verylow-viscosity contrast between the quartz and the weakphases. The stronger quartz can be seen as a deformableinclusions in weak matrix (Figure 15c). In order to developsuch a microstructure the deformation mechanisms of feldsparshave to evolve toward similar efficiency.14 of 20308

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Figure 15. The evolutionary trend of the metagranite deformation. (a) Orthogneiss Type I showinginterconnected weak layer (IWL) structure with strain concentration in plagioclase. The K-feldspar showsinternal load-bearing framework structure with weak plagioclase strings, while quartz exhibits weakCPO, both indicating shielding of quartz and feldspar from viscous flow. (b) Orthogneiss Type II ismarked by almost uniform flow in plagioclase and K-feldspar monomineralic layers around the strongquartz-IWL structure with low viscosity contrast. The rock shows development of monomineraliclayering and intense yielding of quartz marked by strong CPO and activity of prism < a > slip.(c) Orthogneiss Type III shows a uniform flow of all mineral phases and mixing of feldspars associatedwith important crystal growth. The quartz texture exhibits weakening comparing to the previous stageand activity of the basal < a > slip.8.3. Mechanism of Alkali Feldspar Breakdown:Transition From Type I to II Rock[43] The key element controlling the rheological developmentof the bulk rock is the mechanism of decompositionof originally large alkali feldspar phenocrysts into chains ofplagioclase grains and adjacent K-feldspar hosts (Figure 4).EBSD measurements have shown that the texture of thenew plagioclase crystals in any structural position is almostidentical to that of the K-feldspar hosts (Figure 13), indicatingan orientation relationship between the host and theinclusion.[44] The above described features do not have an unambiguousexplanation. The coherent texture of the plagioclaseand host K-feldspar grain may reflect a heterogeneousnucleation process [Putnis et al., 2003; Ribbe, 1983]. TheTime-Temperature-Transformation diagrams for the feldsparexsolution during cooling suggest that a process of heterogeneousnucleation is the most likely mechanism because ofthe slow cooling rates (which are likely to occur in deepseated intrusions) [Putnis et al., 2003]. Our microstructuralanalyses show that the shape and size of the exsolutionpatterns is determined by the structure of the host namely by(100) and (010) planes of alkali feldspars. The experimentsof Putnis et al. [2003] show that the albite rich regionsoriginate as a monoclinic feldspar exsolution in the K-richhost. With falling temperatures the albite changes to a highalbitetriclinic structure. The strain generated at the lamellainterface leads to the segmentation of the lamella by thealbite twinning reducing the strain energy across the interfaceregion. It is possible that such twinned lamellae coarsenduring the subsequent annealing which may lead to thedevelopment of new albite grains. The process of coarseningmay have been enhanced by the thermally induceddeformation so that first subgrains and subsequently developednew grains originated from the progressively deformingtwinned lamellae. All these processes may result in thedevelopment of chains of albite grains subparallel to theformer exsolution domains. The process of the heterogeneousnucleation is supported by the compositional profileswhich show a decrease of albite components in the hostK-feldspar toward the albite boundary of rather constantcomposition similarly to the profiles published by Putniset al. [2003]. In theory the exsolution process should berelated either to the cooling history of crystallization ofthe granite or to the deformation-metamorphism processas suggested by White and Mawer [1986, 1988]. However,the spatial distribution of plagioclase chains in thefeldspar host and the crystallographic coherency suggestthat the transformation of the original alkali feldspar wasachieved by heterogeneous nucleation.8.4. Consequences of Feldspar Breakdown:Monomineralic Layering in Type II Rocks[45] The progressive textural evolution toward aggregatedistribution in the Type II microstructure is a typical featureof the high-grade deformation of granitoids [Gapais, 1989;Handy, 1994a; Schulmann et al., 1996]. The onset of granitedeformation is marked by the high stress concentrations inthe plagioclase grains due to the relative rheological inactivityof the quartz and K-feldspar at the onset of themetagranite deformation (Type I microstructure). The grainsize increase of the recrystallized plagioclase and K-feldsparassociated with the development of monomineralic layeringcan be interpreted in terms of stress relaxation [Hobbs,1981] coupled with a flow stress increase in the quartzaggregates indicated by the activity of the prism < a > slipsystem compared to the basal < a > in the Type I microstructure[Schmid and Casey, 1986]. Alternatively, theevolution of the quartz slip system can be attributed to theincreasing strain intensity in the Type II microstructure[Heilbronner and Tullis, 2006]. Rather unusual feldsparslip systems like once in Type II rocks (Figure 13) havebeen reported by Baratoux et al. [2005] from high-grademylonitic metagabbros and by Franěk et al. [2006] frompartially molten granulites. These authors interpreted theobserved slip systems as a result of a grain boundary slidingaccompanied with a crystal plastic deformation. All this15 of 20309

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406implies a progressive homogenization of the stress fieldduring this stage. Differences in the efficiency of thedeformation mechanisms lead to the development of monomineraliclayering precluding effective mechanical mixing.We emphasize that the involvement of feldspars in thehomogeneous flow increases the proportion of the weakmaterial up to 60%, which provided a significant drop of thebulk strain rate or stress assuming constant far field stressesor a bulk strain rate, respectively (Figure 15b). The effectiveweakening mechanism could be regarded in the presence offluid or melt phase at grain boundaries in Type II rocks whichis represented by the interstitial albitic Pl2 and quartz formingthin films parallel to the K-feldspar and Pl1 boundaries or assmall cuspate grains in triple point junctions.8.5. Reasons for Transition From Type II to Type IIIRock: Melt-Assisted Mineral Mixing[46] Type III microstructures correspond to a mixing offine-grained plagioclase and K-feldspar by two competitiveprocesses: a mechanical mixing accompanied by syndeformationalmass transfer either by diffusion [Baratouxet al., 2005; Kruse and Stünitz, 1999] or by melt/fluidredistribution along grain boundaries during the in situmelting [Závada et al., 2007; Hasalová et al., 2008] orinfiltration of hydrous fluids, respectively [Stünitz and FitzGerald, 1993]. The mechanical mixing process is supportedby the progressive thinning of the plagioclase aggregates ina K-feldspar matrix associated with the extreme stretchingof the rock at very high strains. This process is inevitablyassociated with a grain boundary sliding mechanism anddiffusion dominated creep. The growth of albite and quartzrims around the K-feldspar grains and the overall grain sizeincrease of both feldspars can result from the above mentionedmass transfer mechanisms.[47] The intensity of quartz CPO in Type III microstructuresdecreases compared to Type II microstructures. This isdue to a switch of the active slip system from prism < a >toward basal < a > and rhomb < a + c > slip systems in TypeIII microstructure (Figure 15c). These slip systems are notcommon in partially molten high-grade mylonites where theactivity of prism < c > glide is reported [e.g., Gapais andBarbarin, 1986; Martelat et al., 1999]. Our observations donot indicate a late reworking during a low-temperature eventand we therefore suggest that the activity of the ‘‘lowtemperature’’slip systems in high-grade mylonites resultsfrom strain rate or stress conditions which are currentlyunconstrained [Závada et al., 2007; Hasalová etal., 2008].Consequently, the deformation mechanism in quartz is thedislocation creep and the evolution of slip systems suggest adecrease of the flow stress or strain rate compared toprevious stages [Handy, 1990; Herwegh et al., 1997;Schmid and Casey, 1986]. However, the diffusional creepin the feldspar matrix supported by grain shapes andcrystallographic preferred orientations (Figure 13) showsthat the Type III microstructures represent a deformationstage marked by the overall flow stress drop associated withthe increase of homogeneity of the strain distribution. Thisoverall weakening is consistent with interstitial phasesforming a substantial proportion of the rock volume andshapes of pools or thick films of Pl2 and quartz grains thatare similar to the melt topology described by Sawyer [2001]or Marchildon and Brown [2003].8.6. Quantitative Microstructural and PetrologicalArguments for Syn-deformational Melting[48] The CSD curves (Figure 9b) show decreasing N 0 andincreasing G t values in relict plagioclase aggregates towardType III rocks. The trend from Type II to Type III rocks(Figure 9b) consistent with the rapid nucleation of interstitialPl2 is marked by the presence of thin isolated films inSEM images of Type II microstructures (Figure 6, samplesT1 and M2, and Figure 8). The decrease of N 0 and increaseof G t values in Type III microstructures correspond to agrowth of Pl2 pools and large new Pl2 rims in a K-feldspardominated matrix (Figure 7, samples V1 and D1). Thisevolution corrobarates the weakening of the aggregatedistribution in Type II microstructures toward an almostrandom distribution in Type III microstructures and anassociated development of important proportion of theinterstitial phases (Figure 11).[49] The pseudosection modeling indicates that the onlyfluid that could be present during the deformation of theorthogneiss at the estimated temperatures of 650–680°C isa silicate melt (Figure 14). Either the temperature increaseor the pressure drop along the path estimated by Pitra andGuiraud [1996] or Tajcmanová etal.[2006] for associatedmetapelites can be responsible for the dehydration meltingof the muscovite-biotite bearing orthogneiss assemblage.The amount of melting may even be increased by externalfluids introduced from the surrounding metapelites[Thompson and Connolly, 1995].[50] The nucleation dominated part of the deformationmeltinghistory in Type II microstructures is probablyassociated with the onset of the melting reaction andheterogeneous nucleation of the melt droplets at highenergytriple points or noncoherent grain boundaries. Thegrowth dominated process associated with the Type IIImicrostructures reflects a more advanced melting leadingto the development of large pools and the coalescence ofsmall nuclei along all boundaries in deformed aggregates.This process is best exemplified by the sample B2 whichresembles a migmatitic structure at the macroscopic scale.8.7. Grain Boundary Sliding Diffusional Creep andVariations in Melt Topology With Increasing MeltFraction[51] The studied microstructural sequence is characterizedby the increasing deformation intensity (Figure 3), theincreasing melt fraction (Figures 4, 6, and 7), grain size(Figure 9) and random grain contact distribution as well asthe low aspect ratio of grains (Figures 10 and 11). Thisevolutionary trend is accompanied by a change in orientationof the melt seams (Figures 8 and 16) from an orientationsubparallel to the foliation to seams oriented at highangle to the foliation.[52] Melt seams preferentially located along grain boundariesparallel to the foliation in Type II microstructures arecommonly reported in natural samples [Sawyer, 2001;Rosenberg and Berger, 2001] and in some experimentalor analog studies [Daines and Kohlstedt, 1997; Groebnerand Kohlstedt, 2006; Walte et al., 2005]. This geometry canbe observed in microstructures of the T1 and T2 samplesand is compatible with the combined activity of grainboundary sliding and crystal-plastic deformation mechanismsat very low melt fractions (

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406Figure 16. The evolution of (a) shear plane-parallel melt preferred orientation at low melt fractions and(b) shear plane-perpendicular melt preferred orientation at high melt fractions. At low melt fractions,shear is accommodated by the grain boundary sliding parallel to the shear direction. Since grain boundarysliding means a loss of the cohesion along these boundaries, they represent an easy path for melt whichaccumulates at these boundaries (bottom left, Figure 16a). At higher melt fractions, a switch from shearplane parallel to shear plane-perpendicular melt preferred orientation can be observed. This is probablybecause of stress concentrations at grain boundaries oriented perpendicular to simple shear andaccompanying dilation/cavitation of these boundaries (bottom right, Figure 16b). See text for furtherexplanations.orientation of melt seams can be explained by a loss ofcohesion at boundaries oriented parallel to the foliation.Because of shearing, the melt concentrated at triple pointssuffers an increasing overpressure. We suggest that thecombination of loss of cohesion and build up of meltoverpressure leads to the injection of melt between activelymoving grain boundaries oriented at a high angle to theprincipal compressive stress, i.e., parallel to the macroscopicfoliation (Figure 16a). Oblique intragranular fractures canbe interpreted as Riedel shears channeling melt from overpressured,dilatant grain boundaries accompanying movementsalong foliation-parallel grain boundaries [Rosenbergand Handy, 2000]. The M1 and M2 samples are characterizedby a general constrictional deformation, lower aspectratio of grains and low melt fraction (

B10406SCHULMANN ET AL.: RHEOLOGY OF PARTIALLY MOLTEN GNEISSESB10406suitably oriented planes in the aggregate [Hubbert andRubey, 1959]. As a result of melt overpressure, the feldsparaggregate dilates and the melt is accumulated in intergranularand intragranular pockets and fractures.[56] It is likely that the evolution in the orientation of meltpockets is caused by the increasing melt fraction and/or bythe change in differential stress magnitude and orientation[Cosgrove, 1997]. The implication of the increasing meltfraction and connectivity certainly influences the distributionof local compressive and shear stress magnitudesand orientations and the active deformation mechanism(Figure 16). In this context, it is not likely that the meltpreferred orientation is controlled by the variations indifferential stress [Gleason et al., 1999], hydrostatic annealing[Daines and Kohlstedt, 1997] or variations in thewetting angle [Walte et al., 2005]. In order to further explainthis relationship between melt seam orientations and meltfractions more analogs and possibly numerical modeling isneeded.[57] In Type II microstructures the diffusion creep ratesof both feldspars were effectively enhanced by the interstitialmelt phase wetting their boundaries and leading to arapid strength drop within the MCT field as proposed by[Rosenberg and Handy, 2005]. This is consistent with a lowamount of melt

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J. metamorphic Geol., 2008, 26, 29–53 doi:10.1111/j.1525-1314.2007.00743.xOrigin of migmatites by deformation-enhanced melt infiltrationof orthogneiss: a new model based on quantitativemicrostructural analysisP. HASALOVÁ, 1,2 K. SCHULMANN, 1 O. LEXA, 1,2 P. ŠTÍPSKÁ, 1 F. HROUDA, 2,3 S. ULRICH, 2,4J. HALODA 5 AND P. TÝCOVÁ 51 Université Louis Pasteur, CGS/EOST, UMR 7517, 1 rue Blessig, Strasbourg 67084, France (hasalovap@seznam.cz)2 Institute of Petrology and Structural Geology, Charles University, Albertov 6, 12843 Prague, Czech Republic3 AGICO, Ječná 29a, 621 00 Brno, Czech Republic4 Institute of Geophysics, Czech Academy of Sciences, Boční II/1401, 14131 Praha 4, Czech Republic5 Czech Geological Survey, Klárov 3, 118 21 Prague 1, Czech RepublicABSTRACTA detailed field study reveals a gradual transition from high-grade solid-state banded orthogneiss viastromatic migmatite and schlieren migmatite to irregular, foliation-parallel bodies of nebulitic migmatitewithin the eastern part of the Gfo¨ hl Unit (Moldanubian domain, Bohemian Massif). The orthogneiss tonebulitic migmatite sequence is characterized by progressive destruction of well-equilibrated bandedmicrostructure by crystallization of new interstitial phases (Kfs, Pl and Qtz) along feldspar boundariesand by resorption of relict feldspar and biotite. The grain size of all felsic phases decreases continuously,whereas the population density of new phases increases. The new phases preferentially nucleate alonghigh-energy like–like boundaries causing the development of a regular distribution of individual phases.This evolutionary trend is accompanied by a decrease in grain shape preferred orientation of all felsicphases. To explain these data, a new petrogenetic model is proposed for the origin of felsic migmatites bymelt infiltration from an external source into banded orthogneiss during deformation. In this model,infiltrating melt passes pervasively along grain boundaries through the whole-rock volume and changescompletely its macro- and microscopic appearance. It is suggested that the individual migmatite typesrepresent different degrees of equilibration between the host rock and migrating melt duringexhumation. The melt topology mimicked by feldspar in banded orthogneiss forms elongate pocketsoriented at a high angle to the compositional banding, indicating that the melt distribution wascontrolled by the deformation of the solid framework. The microstructure exhibits features compatiblewith a combination of dislocation creep and grain boundary sliding deformation mechanisms. Themigmatite microstructures developed by granular flow accompanied by melt-enhanced diffusion and/ormelt flow. However, an AMS study and quartz microfabrics suggest that the amount of melt present didnot exceed a critical threshold during the deformation to allow free movements of grains.Key words: crystal size distribution; melt infiltration; melt topology; migmatites; quantitative texturalanalysis.INTRODUCTIONMovement of a large volume of granitic melt is animportant factor in the compositional differentiationof the continental crust (Fyfe, 1973; Collins & Sawyer,1996; Brown & Rushmer, 2006) and the presence ofmelt in rocks profoundly influences their rheology(Arzi, 1978). The migration of melt through the crust iscontrolled by melt buoyancy and pressure gradientsresulting from the combination of gravity forces anddeformation (Wickham, 1987; Sawyer, 1994). Thereare three major mechanisms controlling melt migrationthrough the continental crust: (i) diapirism resulting inupward motion of low-density magma through higherdensity rocks (Chandrasekhar, 1961; Ramberg, 1981);(ii) dyking that describes melt migration by hydrofracturingof the host rock and transport of meltthrough narrow dykes (Lister & Kerr, 1991; Petford,1995); (iii) and migration of a melt through a networkof interconnected pores during deformation or compactionof solid matrix (McKenzie, 1984; Wickham,1987).Brown & Solar (1998a) and Weinberg & Searle(1998) proposed that during active deformation meltmoves by pervasive flow and it is essentially pumpedthrough the system parallel to the principal finiteelongation in the form of foliation-parallel veins.Based on a number of field studies, pervasive meltmigration at outcrop scale controlled by regionaldeformation has been suggested by various authors(Collins & Sawyer, 1996; Brown & Solar, 1998b;Vanderhaeghe, 1999; Marchildon & Brown, 2003).Ó 2007 Blackwell Publishing Ltd 29315

30 P. HASALOVÁ ET AL.These authors argued that magma intrudes pervasively,parallel to the main anisotropy represented byfoliation planes (John & Stu¨ nitz, 1997), fold hinges andinterboudin partitions (Brown, 1994; Brown et al.,1995). It is also commonly observed that vein-likeleucosomes are injected into extensional structuresprovided the magma pressure is high enough (Wickham,1987; Lucas & St.Onge, 1995) or parallel to axialsurfaces of folds (Vernon & Paterson, 2001).Microscopic studies of natural rocks show orientationsof former melt microstructures that are interpretedin terms of grain-scale channel networks(Sawyer, 2001). Melt migration pathways at the grainscale are commonly determined from distribution ofmelt films and pools (now glass) in experimentallyprepared samples or by distribution of minerals supposedto preserve the original melt topology in naturalrocks (Brown et al., 1999; Rosenberg & Riller, 2000;Rosenberg, 2001). The melt topology in experiments iscontrolled mainly by differential stress, confiningpressure and the amount of melt in the system(Rosenberg, 2001). At static conditions, the melttopology is characterized by equilibrium dihedral(wetting) angles at triple point junctions (Jurewicz &Watson, 1984; Laporte & Watson, 1995; Laporteet al., 1997; Cmı´ral et al., 1998; Walte et al., 2003) andthe mobility of the melt remains very low, even if themelt phase forms an interconnected network alongtriple-junction grain edges at dihedral angles lowerthan 60° (Laporte & Watson, 1995; Connolly et al.,1997).Experimental studies on rock analogues to investigategrain-scale melt flow under laboratory conditionsshow that during contemporaneous melting anddeformation melt connection allows the nucleation ofshear bands along which a melt is further segregated(Rosenberg & Handy, 2000, 2001; Barraud et al.,2001). Rosenberg (2001) reviewed the experimentaldata and concluded that the melt migration and meltflow direction are controlled by incremental shorteningand melt pressure gradients between source and areasof melt accumulation.There have only been a few attempts to quantifymelt distribution in rocks using methods of quantitativeand computer aided microstructural analysis(Dallain et al., 1999; Tanner, 1999; Marchildon &Brown, 2003). These studies commonly deal with graincontact frequency distributions, grain size evolutionand orientation of former melt films (Dougan, 1983;McLellan, 1983; Rosenberg & Riller, 2000). However,modern quantitative microstructural analysis mayprovide further important information about: (i)reorganization of the rock structure associated withmelt migration in terms of grain contact distributions(Lexa et al., 2005); (ii) characterization of dynamic orstatic conditions of melt movement through rocksusing analysis of grain boundaries and shape orientations(Rosenberg & Riller, 2000; Marchildon & Brown,2002); and (iii) cooling or heating histories of rocksusing crystal size distribution (CSD) theory (Higgins,1998; Berger & Roselle, 2001).In this work, a sequence of deformed felsic rocks isstudied, ranging from high-grade banded orthogneissto fine-grained isotropic migmatite both at macro- andmicroscale using structural, petrographic and quantitativemicrostructural analyses. It is shown that asequence of banded orthogneiss, stromatic, schlierenand nebulitic migmatites results from progressivedeformation in a crustal-scale shear zone in the presenceof melt. The microstructural and fabric modificationsconnected with disintegration of parentalbanded orthogneiss and development of random mineralmicrostructure are quantified. The relationships ofthe individual rocks types and the possible origin ofthis sequence are discussed in terms of deformationand migmatization of different protoliths, meltinfiltration from an external source or in situ melting ofthe same protolith during progressive deformation. Itis argued that banded orthogneiss and nebuliticmigmatites can be interpreted as end-members of acontinuous sequence resulting from melt infiltrationfrom an external source during deformation. Finally,the role of melt for activity of grain-scale deformationmechanisms and bulk rheological behaviour of crustalrocks during melt infiltration is discussed.GEOLOGICAL SETTINGThe Moldanubian zone represents the highest gradeunit of the Bohemian Massif and is interpreted as aninternal zone of the Variscan orogen developed duringthe Variscan convergence (Matte et al., 1990). TheMoldanubian zone is comprised essentially of highgradegneisses and migmatites containing relicts ofhigh-pressure felsic granulites, eclogites and peridotitesthat are intercalated with mid-crustal rocks (Fig. 1a).Schulmann et al. (2005) described the structural andmetamorphic evolution of high-grade crustal rocks ofthe so-called Gfo¨ hl Unit and of the adjacent middlecrustal units. For the mechanism of exhumation, theyproposed a model of vertical extrusion of orogeniclower crust and its lateral spreading in mid-crustallevels due to an indentation of the easterly Bruniapromontory. As a consequence of this process, thehigh-pressure rocks were thrust over adjacent middlecrustal units in conjunction with retrogression of originalmineral assemblages and partial melting of all therock types (Sˇtípska´ et al., 2004).Fig. 1. (a) Geological map of the eastern margin of the Bohemian Massif (modified after Schulmann et al., 2005) with the location ofthe study area (black rectangle). The upper right inset shows the general location of the Bohemian Massif within the EuropeanVariscides. (b) Schematic block diagram displaying the main structural features in the study area (modified after Schulmann et al.,2005). Dominant S 1 and S 2 fabrics with their orientations are shown. This block diagram is not vertically scaled.Ó 2007 Blackwell Publishing Ltd316

ORIGIN OF FELSIC MIGMATITES 31Ó 2007 Blackwell Publishing Ltd317

32 P. HASALOVÁ ET AL.The onset of the exhumation process is dated byzircon U–Pb ages of c. 340 Ma on felsic granulites,migmatites and mantle-derived syn-tectonicallyemplaced plutons (van Breemen et al., 1982; Kro¨ neret al., 1988; Holub et al., 1997; Schulmann et al.,2005). Tajčmanova´ et al. (2006) assigned metamorphicconditions of 840 °C at 18–19 kbar and 760–790 °C at10–13 kbar to relict steep granulitic fabrics whichoriginated by vertical extrusion of lower crust. Theseauthors also estimated the conditions of re-equilibrationof granulites associated with horizontal spreadingstage to 720–770 °C and 4–4.5 kbar. High pressurerocks of the Gfo¨ hl Unit are accompanied by largebodies of biotite–sillimanite Gfo¨ hl orthogneiss spatiallyassociated with K-feldspar–sillimanite paragneissesand leucocratic migmatites for which P–Tconditions of 7–10 kbar and 750 °C were estimated byRacek et al. (2006).The area of this study is located at the easternmosttermination of the Gfo¨ hl Unit (Fig. 1a). The main rocktype is represented by the Gfo¨ hl orthogneiss withprotolith ages 488 ± 6 Ma (U–Pb SHRIMP: Friedlet al., 2004) including small bodies of amphibolite,granulite, eclogite, ultrabasic rock and paragneiss. TheGfo¨ hl orthogneiss shows different stages of migmatizationcharacterized by the assemblage of Kfs +Pl + Qtz + Bt ± Grt ± Sill. This migmatizedorthogneiss complex is heterogeneously deformed bytop to the NE shearing along a large-scale, gentlydipping shear zone (Schulmann et al., 1994). Consequently,the northern margin of this complex is thrustover a footwall comprised of the Na´meˇsˇtÕ granulitebody and Neoproterozoic metagranites of the northeasterncontinental margin (Urban, 1992).STRUCTURAL EVOLUTIONMesoscopic structuresTwo major deformation events are recognized. Thefirst event (D 1 ) is represented by a steep, west-dippinghigh-grade foliation S 1 (Fig. 1b). This fabric is preservedin banded orthogneisses (type I), as an alternationof recrystallized monomineralic K-feldspar,plagioclase and quartz layers, separated by bands ofbiotite ± sillimanite (Fig. 2a). Lineation L 1 is locallymarked by alignment of biotite, sillimanite and byelongation of quartz and feldspar aggregates. Thisdeformation is attributed to an early stage of exhumationof the lower crust along a vertical channel(c)Type III(d)Type IV(b)Type II1 cm2 cm1cmd(a)Type Icba1cm1.0 mFig. 2. Schematic representation of the rock relationships at an outcrop scale and photographs of the individual rock types. (a) Bandedorthogneiss (type I); (b) stromatic migmatite (type II); (c) schlieren migmatite (type III); and (d) nebulitic migmatite (type IV). Theposition of this outcrop in the study area is shown in Fig. 1b.Ó 2007 Blackwell Publishing Ltd318

ORIGIN OF FELSIC MIGMATITES 33during horizontal shortening of the thickened orogenicroot (Schulmann et al., 2005).The second deformation (D 2 ) is associated withreworking and folding of the S 1 compositional layeringin banded orthogneiss, so that S 1 is only preservedlocally in elongate relict domains (Fig. 2a). The D 2shearing is attributed to horizontal flow of hot lowercrust in a zone up to 10 km wide at a mid-crustal levelabove the Brunia promontory over distances of severaltens of kilometres (Schulmann et al., 2005). Relicdomains with gently folded S 1 fabric are surrounded byhighly deformed zones with tightly folded S 1 fabric(Fig. 2). The composite S 1)2 fabric is characterized bya banded structure with diffuse boundaries betweenpolymineralic K-feldspar- and plagioclase-richdomains similar to a stromatic migmatite structure(type II) (Fig. 2b). Locally the S 1 fabric is completelytransposed and a new S 2 foliation is dipping gently tothe SW (Fig. 1b). A sub-horizontal, gently S–SWplunging L 2 lineation (Fig. 1b) is mostly defined bypreferred orientation of sillimanite.Detailed field observations reveal that with ongoingdeformation the type II rock gradually pass into moreisotropic rock (type III) composed of K-feldspar–quartz and plagioclase–quartz aggregates (Fig. 2c) andcontaining rootless folds of the deformed S 1 fabric.This rock type alternates with irregular bodies orelongate lenses of fine-grained isotropic felsic rock(type IV, Fig. 2d), which in this region traditionallyhas been described as a nebulitic migmatite (Matějovska´,1974). Such a structural sequence originatedthrough intense D 2 deformation superimposed on anolder steep anisotropy and was identified in outcropscale along several sections. These observations aresupported by the existence of macroscopically visibleleucosomes or granitic veins that are also parallel to S 2and form isolated elongate pockets and lock-up shearbands.This area has been extensively studied by Mateˇjovska´(1974) and Dudek et al. (1974) who used theclassical migmatite terminology of Mehnert (1971) forthe above-described rock types. These authors identifiedtype I rock as banded orthogneiss, rock type II asstromatic migmatite and rock type IV as nebuliticmigmatite. Rock type III resembles the schlierenmigmatite of Mehnert (1971). Because the Gfo¨ hl Unitis considered as one of the largest migmatitic terranesof the Variscan belt, the traditional migmatite terminologywas adapted to these rocks.MICROSTRUCTURAL OBSERVATIONSThe microstructural characteristics including grainsize, grain shape and grain boundary geometry werestudied in each of the four rock types and inK-feldspar- and plagioclase-rich domains. Thin sectionswere cut perpendicular to the foliation andparallel to L 2 lineation (XZ section). To discriminateK-feldspar from plagioclase, the thin sections werestained according to the method of Bailey & Stevens(1960).Type I: banded orthogneissThis rock type is a fine-grained orthogneiss with 0.25-to 2.0-mm-thick layers of recrystallized plagioclase(30 modal%), K-feldspar (40 modal%) and quartz(20 modal%), separated by discrete layers of biotite(10 modal%) commonly associated with minor sillimaniteand garnet (Table 1, Fig. 3a).K-feldspar forms completely recrystallized aggregates(0.2–0.8 mm grain size) with straight grainboundaries locally meeting in triple point junctions at120° (Fig. 4a). Numerous rounded inclusions of quartz(0.05 mm) occur preferentially at triple points, alongplanar boundaries or in cores of feldspar (Fig. 4a).Plagioclase (An 10)20 ) is present in K-feldspar aggregatesas small interstitial grains or forms thin filmspreferentially tracing those K-feldspar boundaries thatare oriented at a high angle to the foliation (Fig. 5a).Rarely, tiny interstitial biotite is present in the K-feldspar-rich bands.Plagioclase aggregates (0.2–0.5 mm) are composedof an equidimensional polygonal mosaic with straightboundaries, and minor interstitial quartz and biotite(Fig. 4b). The plagioclase grains show abundanttwinning and form a foam-like texture with a perfecttriple point network of grain boundaries. Plagioclaseexhibits normal zoning with homogeneous oligoclasecores (An 24)28 ) and more sodic (An 10)18 ), clear,2to10lm-thick rims at boundaries with K-feldspar.Plagioclase grain size continuously decreases from thecentre of an aggregate towards its borders. Quartzoccurs as small (0.01–0.05 mm) rounded inclusions orinterstitial grains, whereas K-feldspar exhibits characteristiccuspate shapes (Fig. 5b). Tiny biotite grains(0.1–0.5 mm in length; X Fe ¼ 0.42–0.48, Ti ¼ 0.2–0.27 p.f.u.) commonly occur along the plagioclaseboundaries that are sub-parallel to the foliation(Fig. 3a).Quartz ribbons 0.3–1.0 mm wide are composed ofelongate grains with straight grain boundaries perpendicularto the ribbon margin (Fig. 3a). Quartz–feldspar boundaries are gently curved, with cusps thatpoint from feldspar to quartz. Biotite-rich layerscommonly show decussate microstructure, which is atextural equivalent of the foam-like texture of the felsicminerals (Vernon, 1976). Contacts between biotiteandplagioclase-rich layers are marked by numerous(

34 P. HASALOVÁ ET AL.Table 1. Representative data for the quantitative textural analysis.Banded orthogneiss Stromatic migmatite Schlieren migmatite Nebulitic migmatiteKfs domain P1 domain Kfs domain P1 domain Kfs domain P1 domainGrain size – Feret diameter (mm)Median Kfs 0.430 0.121 0.345 0.065 0.172 0.138 0.137P1 0.134 0.224 0.086 0.225 0.094 0.119 0.110Qtz 0.079 0.076 0.071 0.079 0.070 0.076 0.074Ql Kfs 0.120 0.065 0.194 0.042 0.103 0.082 0.085P1 0.103 0.095 0.055 0.158 0.061 0.084 0.074Qtz 0.046 0.051 0.044 0.050 0.046 0.047 0.039Q3 Kfs 0.630 0.170 0.556 0.101 0.263 0.211 0.237P1 0.257 0.373 0.161 0.350 0.172 0.164 0.160Qtz 0.105 0.114 0.119 0.127 0.105 0.121 0.127Q3 ) Q1 Kfs 0.510 0.105 0.362 0.059 0.161 0.129 0.152P1 0.154 0.278 0.107 0.192 0.111 0.080 0.086Qtz 0.059 0.063 0.075 0.077 0.060 0.074 0.089Crystal size distribution (CSD)N0. (mm )4 ) Kfs 0.0037 – 0.00487 – 0.053 – 0.2124P1 – 0.0303 – 0.0733 – 0.06812 0.1857Qtz 1.008 2.334 1.6448 3.73 2.093 0.6585 2.286Gt Kfs 0.347 – 0.286 – 0.148 – 0.1127P1 – 0.15731 – 0.1269 – 0.11 0.0689Qtz 0.0736 0.0569 0.0669 0.0547 0.0644 0.0813 0.0623Shape preferred orientation (SPO)Eigenvalue ratio (Rg) Kfs 1.48 1.42 1.42 1.32 1.21 1.13 1.15P1 1.17 1.42 1.1 1.23 1.21 1.14 1.13Qtz 1.25 1.47 1 1.24 1.1 1.33 1.32Aspect ratio (median) Kfs 1.66 1.69 1.6 1.5 1.59 1.6 1.55P1 1.51 1.6 1.59 1.44 1.65 1.5 1.61Qtz 1.5 1.5 1.46 1.5 1.5 1.5 1.49Bt 2.14 2.7 2 2.2 2.2 2.35 2.2Grain boundary preferred orientation (GBPO)Eigenvalue ratio (Rb) Kfs–Kfs 1.34 – 1.25 – 1.06 – 1.5Kfs–P1 1.15 1.18 1.09 1.13 1.14 1.13 1.12Kfs–Qtz 1.17 – 1.15 – 1.17 – 1.17P1–P1 – 1.18 – 1.15 – 1.15 1.36Modal proportion (%)Kfs 70–80 10 70–80 5 50 20–25 30P1 10 60 10 60 20 40 30Qtz 10–20 20 10–20 25 30 30 30Bt

ORIGIN OF FELSIC MIGMATITES 35(a)Type Iplagioclase (Fig. 4c). Biotite (10–15 modal%; X Fe ¼0.76–0.79, Ti ¼ 0.18–0.19 p.f.u.) is homogeneouslydispersed and is most prevalent in the plagioclase–quartz domains. Atoll-shaped garnet (0.05–0.25 mm insize; X Fe ¼ 0.96–0.97) appears inside the felsic aggregates,rather than along contacts with biotite.Type IV: nebulitic migmatiteThis type of rock is composed of almost equal amountsof plagioclase, K-feldspar and quartz, and containsminor biotite (X Fe ¼ 0.91–0.93, Ti ¼ 0.01–0.04 p.f.u.),sillimanite and garnet (X Fe ¼ 0.98–1.00) (Fig. 3d),with a weakly developed preferred orientation of thebiotite and sillimanite; modes are given in Table 1. K-feldspar (0.10–0.25 mm in size) occurs in the form ofirregular grains embayed with quartz and plagioclase.Commonly, the intensity of quartz and plagioclaselobes correlates well with highly cuspate irregularforms of corroded relics of K-feldspar (Fig. 4e). Similarly,the relics of irregular plagioclase (0.05–0.15 mmin size; An 6)10 in the core and An 0)4 at the rim) showcuspate boundaries, but with curvature less pronouncedthan that of the corroded relics of K-feldspargrains. An important feature is the presence of newplagioclase (An 0)1 )–K-feldspar intergrowths embayingcorroded relics of K-feldspar grains (Fig. 4f). Quartz(0.04–0.07 mm) with highly lobate boundaries is uniformlydistributed in the rock. Biotite of low aspectratio shows highly corroded cuspate forms filled withquartz, K-feldspar and plagioclase.(b)(c)1 mm1 mmType II1 mmType IIISummary of modal changesModal composition of the feldspar aggregates in thetype II migmatite does not change significantly comparedwith the type I orthogneiss. However, the typeIII migmatite is characterized by an important increasein quartz content in feldspar domains (up to 30 modal%)associated with a slight increase in interstitialplagioclase in K-feldspar-rich domains and K-feldsparin plagioclase-rich domains. The proportions of thefelsic minerals are equal in the type IV migmatite.(d)KfsPlQtzBt,Sill,Grt1 mmType IVFig. 3. Representative digitalized microstructures (XZ sections)for individual textural types (note differences in scales whenmaking comparisons). (a) Banded orthogneiss (type I) with distinctmonomineralic layers composed of a polygonal mosaic ofwell-equilibrated plagioclase, K-feldspar and quartz polycrystallineribbons separated by discrete layers of biotite ± sillimanite± garnet (sample PH60/B). (b) Stromatic migmatite (type II)composed of K-feldspar-rich, plagioclase-rich and quartz-richaggregates separated by relicts of biotite ± sillimanite-rich layers(sample PH60/A). (c) Schlieren migmatite (type III) showingalternation of K-feldspar- and plagioclase-rich domains interpretedto correspond to an original spatial distribution (K-feldspardomain is shown, sample PH90). (d) Isotropic nebuliticmigmatite without any gneissosity (type IV) composed of equalamounts of K-feldspar, plagioclase and quartz (sample PH59/D).Ó 2007 Blackwell Publishing Ltd321

36 P. HASALOVÁ ET AL.Evidence of meltingSawyer (1999, 2001) summarized criteria for recognitionof former melt at grain scale in metamorphicrocks. The three most important features are: (i) mineralpseudomorphs after thin melt films along crystalfaces, a feature typically observed in melting experimentsunder dynamic conditions (Jin et al., 1994); (ii)rounded and corroded reactant minerals embayed bysurrounding mineral pseudomorphs after melt (Bu¨ sch(a)(b)KfsKfsQPlBt0.5 mm0.2 mm(c)Qtz(d)PlPlKfsKfsBtPlKfsQtzKfsQtzKfsPl0.5 mmQtz100 µm(e)(f)PlQtzKfsKfsQtzQtzKfs-PlintergrowthPl100 µmBtPlKfsPl0.2 mmQtz Pl Myrmekites Irregular embayments of relict feldsparÓ 2007 Blackwell Publishing Ltd322

ORIGIN OF FELSIC MIGMATITES 37(a)Pl(An25)(b)QtzQtzQtzQtzAn 10-20An 10-20An 10-20QtzQtzKfsPl(An25)PlKfsQtzKfsKfsQtzQtzQtz0.5 mm100 µm(c)Pl core(An 12-16 )QtzKfs(d)PlKfs(An 15 )(An 0-4)Pl rim(An 1-4 )KfsPl rim(An 1-4)Pl(An 12-16 )(An 1-4 )QtzPlrelict grain(An 12-16)100 µmKfsPlQtzPl(An15)QtzPlQtzKfsQtz200 µmFig. 5. SEM backscatter images showing the inferred former melt topology (note differences in scales when making comparisons). (a)Type I banded orthogneiss: interstitial plagioclase (An 10)20 ), representing the plagioclase component crystallized from the anatecticmelt (grey arrow), tracing the K-feldspar boundaries sub-perpendicular to the foliation (sample PH60/B). Black arrows show smallrounded quartz grains crystallized along feldspar boundaries. (b) Type I banded orthogneiss: inferred former melt pools with cuspatemargins in a plagioclase band (sample PH60/B). The former melt has crystallized to K-feldspar (cuspate melt pools), plagioclase(growing on the old plagioclase grains) and quartz (forming small rounded grains along the feldspar boundaries (black arrow)). (c)Type III schlieren migmatite: more developed interstitial plagioclase (grey arrow) with normal zoning (core ¼ An 12)16 ; rim ¼ An 1)4 )and distinct albite rims (An 1)4 ) on relict feldspar grains (white arrow) (sample PH90). The interstitial plagioclase is not in opticalcontinuity with any residual plagioclase grains adjacent to it and does not show any preferred orientation, in contrast to plagioclase intypes I and II. (d) Type III schlieren migmatite: new plagioclase inferred to have crystallized from melt (growing on an old plagioclasegrain in the form of the discrete albite rims (white arrow)) and quartz grains that resorb relict K-feldspar grains (sample PH14/D).Fig. 4. Photomicrographs showing characteristic textures of the rock sequence (note differences in scales when making comparisons).(a) Type I banded orthogneiss: recrystallized K-feldspar aggregate with straight grain boundaries and numerous smaller roundedquartz grains (white triangles) along the boundaries or in the cores of feldspar (sample PH60/B). (b) Type I banded orthogneiss: welldevelopedplagioclase polygonal foam-like texture with straight grain boundaries, interstitial quartz (white triangles) and biotite(sample PH60/B). (c) Type III schlieren migmatite: typical microstructure with irregularly shaped feldspar and quartz grains withhighly lobate boundaries. Myrmekitic aggregates commonly develop along the K-feldspar boundaries (black arrow). New smallinterstitial plagioclase (grey triangles), K-feldspar and quartz (white triangles) grains trace almost all the relict feldspar boundaries.Interstitial quartz forms preferentially rounded shapes different from plagioclase which forms thin elongated grains/films coatingK-feldspar boundaries (sample PH90). Such a microstructure is typical also for the type IV. (d) Type III schlieren migmatite: irregularcuspate K-feldspar grain embayed with newly crystallized quartz and plagioclase (sample PH90). (e) Type IV nebulitic migmatite:corroded relics of K-feldspar grains (sample PH59/D). (f) Type III nebulitic migmatite: plagioclase-K-feldspar intergrowths embayingrelict K-feldspar grain (sample PH14/D). White arrows in (d), (e) and (f) point to irregular embayments of relict K-feldspar originatedthrough resorption of old K-feldspar grains by newly crystallized material.Ó 2007 Blackwell Publishing Ltd323

38 P. HASALOVÁ ET AL.PlQtz70% 30%Banded orthogneiss (Type I)Stromatic migmatite (Type II)Schlieren migmatite (Type III)Nebulitic migmatite (Type IV)50%30%et al., 1974); and (iii) cuspate and lobate areas inferredto represent pools of crystallized melt (Jurewicz &Watson, 1984).The former presence of melt at grain scale was inferredfrom the following microstructures (Figs 4 & 5).(i) Plagioclase films between adjacent K-feldspargrains, inferred to represent a plagioclase componentcrystallized from melt (Fig. 5a, c). This plagioclase ischaracterized by more albitic composition and by differenttopology compared with original grains. (ii) Pl–Kfs–Kfs and Kfs–Kfs–Pl dihedral angles commonlylower than 30° (Fig. 5a, c), as observed in granitic meltcrystallized under experimental conditions (e.g.Laporte et al., 1997). (iii) Cuspate K-feldspar pools inplagioclase aggregates (Fig. 5b), inferred to represent aK-feldspar component crystallized from melt (Jurewicz& Watson, 1984; Sawyer, 1999, 2001). (iv) Normalzoning of plagioclase from An 10)30 to An 0)15 (Sawyer,1998; Marchildon & Brown, 2001) lining K-feldsparboundaries (Fig. 5c, d). An important feature is thepreferential orientation of plagioclase films coating K-feldspar boundaries in type I orthogneiss and type IImigmatite sub-perpendicular to the foliation (Fig. 5a),in contrast to the types III and IV migmatites, wherethese films are wider and do not show any opticallyvisible preferred orientation. Bulbous myrmekite(Fig. 4d) and new highly irregular lobate grains thatovergrow partially resorbed corroded feldspar grains(e.g. Fig. 4c) are similar to microstructures describedas typical of minerals reacting with melt (Mehnertet al., 1973; Bu¨ sch et al., 1974; McLellan, 1983).QUANTITATIVE TEXTURAL ANALYSISKfsFig. 6. Modal changes in both plagioclase (open symbols) andK-feldspar (closed symbols) aggregates in different rock typesplotted in a quartz–plagioclase–K-feldspar triangle. Arroweddashed lines indicate evolutionary trend from type I bandedorthogneiss to type IV nebulitic migmatite.The quantitative analysis of texture is based onstatistical evaluation of grain size distributions(Kretz, 1966, 1994; Ashworth, 1976; Ashworth &McLellan, 1985; Cashman & Ferry, 1988; Cashman& Marsh, 1988; Higgins, 1998; Berger & Roselle,2001), spatial distribution of minerals and GBPOs(Panozzo, 1983), and grain contact frequencies(Flinn, 1969; Kretz, 1969; McLellan, 1983; Kruse &Stu¨ nitz, 1999). In simple chemical systems, thesetextural parameters are more sensitive to changes ofphysical conditions than compositional characteristics.This is due to the high activation energies ofchemical reactions needed to produce new crystalgrowth compared with the small amount of latticestrain energy and grain boundary energy required todrive recrystallization processes (Spry, 1969; Stu¨ nitz,1998).In this study, the textures of three samples wereanalysed from each rock type, and in each samplemore than 1000 grains were evaluated in thin section.Due to significant textural variations, the individualK-feldspar-rich and plagioclase-rich domains wereanalysed separately. Maps of grains with full topologywere manually traced into the ESRI ArcViewDesktop GIS environment and grain boundaries weregenerated using the ArcView PolyLX extension (Lexa,2003). The ÔshapefilesÕ of individual digitalized thinsections are attached in Appendix S1 (Supplementarymaterial). Analysis of grain size, CSD, grain shapepreferred orientation (SPO), grain boundary preferredorientation (GBPO) and grain contact frequencieswere obtained using the MATLAB TM PolyLX toolbox(Lexa, 2003; http://petrol.natur.cuni.cz/ ondro/). Thegrain sizes of the minerals were evaluated in termsof Feret diameter (diameter of a circle having thesame area as the grain). Two methods were used todetermine the grain SPO: (1) mean directions usingcircular statistics; and (2) eigenvalue analysis ofScheidegger’s bulk orientation tensor calculated fromindividual long axes weighted by grain size (Lexaet al., 2005), where degree of SPO is expressed as theeigenvalues ratio Rg. GBPO was assessed by similartechniques, but the bulk orientation tensor is formedfrom the decomposed grain boundaries betweenchosen phases (Lexa et al., 2005) and the degree ofGBPO is expressed as the eigenvalues ratio Rb. Graincontact frequency, used to examine statistical deviationfrom a random spatial distribution of contactrelations between the individual minerals, was evaluatedin a manner similar to the method of Kretz(1969, 1994), except that contact frequencies wereobtained directly from grain map topologies insteadof using line intercepts.Results of the quantitative microstructural analysesshow an evolutionary trend from the banded orthogneiss,through the migmatite types II and III to thenebulitic migmatite. Therefore, in the following sectionsthe rock types are discussed as a sequence inwhich the type I orthogneiss and type IV nebuliticmigmatite are considered to be end-members of acontinuous microstructural evolution.Ó 2007 Blackwell Publishing Ltd324

ORIGIN OF FELSIC MIGMATITES 39Grain size analysisThe CSD is an important tool to estimate residencetime of magmas in magma chambers, cooling rates inrapidly quenched lavas, as well as to quantify texturesrelated to phenocrysts accumulation and fractionation(Cashman & Marsh, 1988; Marsh, 1988; Higgins,1998). In metamorphic petrology, CSD is used toobtain quantitative information concerning crystalnucleation and growth rates and nucleation densityand/or annealing (Randolph & Larson, 1971; Cashman& Ferry, 1988; Carlson, 1989; Waters & Lovegrove,2002). Hickey & Bell (1996) proposed thatduring dynamic recrystallization decreasing strain rateto temperature ratio (_e/T) leads to decrease in the ratioof nucleation and growth rate (N/G) and developmentof coarser grain size, whereas increasing _e/T leads toincreasing N/G and therefore to grain size decrease.This hypothesis is well documented in experimentalstudies with steel alloys (Sakai & Jonas, 1984) supportedby Azpiroz & Ferna´ndez (2003) and Lexa et al.(2005) in naturally deformed rocks. These authorsevaluate the role of recrystallization mechanisms on N/G ratio of the CSD. The CSD is commonly used informerly partially molten rocks to evaluate combinedprocess of resorption and grain size decrease in reactingphases in mesosome and nucleation and graingrowth and coarsening of minerals crystallizing inleucosomes (Dougan, 1983; McLellan, 1983; Ashworth& McLellan, 1985; Dallain et al., 1999). Because theprocesses controlling grain size distributions in thecrystallization of partially molten rocks are complexand interpretations uncertain, CSD has only rarelybeen used to describe textural evolution of migmatites(Berger & Roselle, 2001). In this work, the CSDmethods are used as a practical approach to parameterizegrain size frequency histograms and visualizetheir trends in a simple manner.Grain size statistics were evaluated for the four rocktypes for plagioclase, K-feldspar and quartz and theresults are presented in the form of average grain size,expressed as a median value of the Feret diameter, andgrain size range expressed as the difference between thethird and first quartiles instead of standard deviationbecause of the log-normal distribution of measureddata (Fig. 7a, Table 1). The results are also summarizedas CSD curves (plot of logarithms of populationdensity against crystal size) that were constructed usingthe method of Peterson (1996); values of the zero-sizeintercept (N 0 – population density interpreted as theratio of nucleation rate to growth rate) and negativeinverse of slope (Gt interpreted as a function of growthrate) of the linear parts of the CSD curves are plottedin Fig. 7b, c.Both plagioclase and K-feldspar in the type Iorthogneiss are characterized by log-normal grain sizedistribution exhibiting average grain size of 0.2 and0.2–0.5 mm respectively. Interstitial quartz yields significantlysmaller average grain size of 0.05–0.1 mm inboth domains. Quartz grains from polycrystalline ribbonswere not evaluated statistically but their grain sizeof 0.5–2.0 mm was estimated using an optical microscope.The grain size of new interstitial plagioclase inthe K-feldspar aggregates is close to 0.1 mm. The grainsize distributions from type I orthogneiss to type II,type III and type IV migmatites are characterized bythe following features. The average grain size of plagioclaseand K-feldspar decreases compared with typeI orthogneiss (Fig. 7a, Table 1). This is accompaniedby a continuous decrease in grain size range for bothfeldspars. The interstitial quartz grain size remainsfairly constant throughout all the stages of texturalevolution, ranging between 0.05 and 0.1 mm, beinglarger in the K-feldspar than in the plagioclasedomains (Fig. 7a). The grain size of minor plagioclasein the K-feldspar domains shows a bimodal distributionthat is attributed to the presence of small newlynucleated grains (0.06–0.1 mm) and to larger plagioclasegrains (0.2 mm) already present in the feldsparaggregates.The CSD of plagioclase indicate continuous increasein N 0 (nucleation density) values coupled with adecrease in Gt (growth rate) values from type Iorthogneiss towards type IV migmatite (Fig. 7b,Table 1). By contrast, K-feldspar shows a decrease inGt values from type I orthogneiss to type III migmatitewithout significant increase in N 0 values, which remainvery low. From type III to type IV migmatite a dramaticincrease in N 0 values is observed for K-feldsparat almost constant Gt values (Fig. 7c). This evolutionis clearly shown by steepening of the slopes of the CSDcurves accompanied by increase in their upper interceptwith the ordinate axis (insets in Fig. 7b, c).Grain shapes and grain shape preferred orientationGrain shape or grain aspect ratio together with grainSPO analyses provide important information aboutdeformation during or after leucosome formation(Mehnert, 1971; McLellan, 1983) or about degree ofinheritance of original anisotropy (Ashworth, 1979).Measurements of preferred orientations of inferredmelt-filled grain boundaries in rocks give insights intoprocesses of melt draining and melt transfer (Rosenberg& Handy, 2000, 2001; Sawyer, 2001).Grain shape and SPO statistics were evaluated in allthe textural types for plagioclase, K-feldspar, quartzand biotite. The results of SPO statistic are summarizedin a boxplot-type diagram, where the axial ratiosof the individual minerals are plotted against bulk SPO(Lexa et al., 2005) for the corresponding minerals(Fig. 8).Aspect ratios for both K-feldspar and plagioclaseshow small median values ranging from 1.5 to 1.7throughout the whole microstructural sequence(Fig. 8, Table 1). Quartz exhibits slightly smaller andstable aspect ratio close to 1.5. An important feature isthe continuous decrease in SPO of K-feldspar andÓ 2007 Blackwell Publishing Ltd325

40 P. HASALOVÁ ET AL.(a)(b)(c)Fig. 7. Grain size statistics and CSD evolution for the rock sequence. (a) Calculated average grain size (median value of the Feretdiameter) and range (difference of third and first quartiles) for plagioclase, K-feldspar and quartz. (b,c) Plots of crystal size distributionparameters N 0 (corresponding to the nucleation density per size per volume) and Gt (non-dimensional value dependent on the growthrate) with examples of linearized CSD curves (upper right insets) used for Gt and N 0 estimates. (b) Plagioclase, (c) K-feldspar. The CSDcurves show single lines of four representative samples corresponding to the individual rock types.plagioclase from type I orthogneiss to type IV nebuliticmigmatite (Fig. 8, Table 1). Rose diagrams for therock types I, II and III show that K-feldspar andplagioclase (Fig. 8b) have weakly inclined SPO withrespect to the aggregate elongation direction at anangle of 15°–30°. Biotite shows a high aspect ratio fortype I orthogneiss (Table 1) and strong SPO parallelwith mesoscopic foliation for the types I, II and IIImigmatites. In the type IV migmatite, biotite aspectratio and preferred orientation are lower and the latterparameter shows bimodal distribution with one maximumsub-parallel to the main foliation and a secondone almost perpendicular to it. Interstitial K-feldspar,plagioclase and quartz exhibit always small aspectratio and weakly developed SPO maxima at an angleof 40°–60° to the foliation for types I, II and III. Theexception is type IV migmatite, where, in similarfashion to biotite, the interstitial plagioclase shows twomaxima, one sub-parallel and one perpendicular to thefoliation.Grain contact frequency analysis and grain boundarypreferred orientationThe grain contact frequency method (Kretz, 1969)allows an examination of the statistical deviationfrom the hypothesis of random distribution of phasesin rocks. In random distribution, the number ofÓ 2007 Blackwell Publishing Ltd326

ORIGIN OF FELSIC MIGMATITES 41(a)(b)Fig. 8. Plot of grain shape preferred orientation (SPO) of K-feldspar (a) and plagioclase (b). The results are summarized in a box plotof aspect ratios (characterizing the shape of grains) v. eigenvalue ratios (showing the degree of preferred orientation). Individual boxesshow median, and first and third quartiles of the aspect ratio. The whiskers represent a statistical estimate of the data range whereoutliers are not plotted. Representative rose diagrams for individual rock types show maxima orientation in respect to the aggregateelongation direction. The degree of shading corresponds to the individual rock types.contacts of given phases depend only on the totalnumber of grains of each phase present. There aretwo possible deviations from random distribution: (i)aggregate distribution, where grains of the samephase tend to occur in aggregates in which contactsbetween grains of the same phase (like–like contacts)predominate; and (ii) regular distribution, where thegrains tend to occur in a regular (chessboard-like)pattern in which contacts between different phases(unlike contacts) are more common. McLellan (1983)reviewed processes responsible for different types ofgrain distributions. A random distribution shouldtheoretically develop during rapid quenching of graniticmelt, whereas regular distribution commonly isinterpreted as resulting from extensive solid-stateannealing under very high temperatures (Flinn, 1969;Vernon, 1976; McLellan, 1983; Lexa et al., 2005).These interpretations are based on the assumption ofreducing surface energy (Seng, 1936; DeVore, 1959)by elimination of high-energy contacts (commonlynon-coherent like–like contacts) either by reduction ofgrain boundary area or by nucleation and growth ofnew phases along such a boundary (Kim & Rohrer,2004). In addition, Kruse & Stu¨ nitz (1999) and Baratouxet al. (2005) proposed that the regular distributionwas induced by mechanical mixing andheterogeneous nucleation. According to Vernon(1976) and McLellan (1983), an aggregate distributionresults from a solid-state differentiation associatedmostly with dynamic recrystallization where developmentof monomineralic layers results from unevenefficiency of deformation mechanisms simultaneouslyoperating in different phases (Jordan, 1988).Grain contact frequency and the GBPO were evaluatedfor K-feldspar, plagioclase and quartz in all thetextural types over the full digitized area of individualK-feldspar and plagioclase domains. The results arepresented in Fig. 9, where the v-valuev ¼ Observed pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiExpectedExpectedor deviation from the random distribution is plottedagainst the ratio of eigenvalues of the orientationtensor or the degree of GBPO (Lexa et al., 2005).Values of expected frequencies are estimated usingLafeber’s method of testing for randomness (Lafeber,1963; Kretz, 1969). This diagram offers a simple visualevaluation of the relationship between degree of deviationfrom expected random distribution of graincontacts and GBPOs of like–like and unlike boundaries.The type I orthogneiss is characterized by a relativelysmall proportion of like–like K-feldspar andplagioclase contacts indicating a weak regular distribution(slightly negative v-values for like–like andpositive v-values for unlike contacts; Fig. 9), despitea macroscopically banded texture in which a strongaggregate distribution should be observed. This featureis attributed to a great proportion of minorinterstitial grains (Qtz, Bt, Kfs and Pl) lining theK-feldspar and the plagioclase boundaries. Additionally,the number of like–like K-feldspar andplagioclase contacts continuously decreases from typeI orthogneiss to type IV migmatite, whereas thenumber of Pl–Kfs, Kfs–Qtz and Pl–Qtz unlikeÓ 2007 Blackwell Publishing Ltd327

42 P. HASALOVÁ ET AL.Fig. 9. Grain boundary statistics plotted as the deviation from a random spatial distribution (grain contact frequency) v. degree ofgrain boundary preferred orientation (GBPO). For details see text. The degree of shading corresponds to the individual rock types.contact continuously increases (negative like–like v-values and positive unlike v-values) (Fig. 9). This isin a good accordance with the increasing amount ofinterstitial phases towards the type IV migmatite.Quartz exhibits the same strong regular distributionfrom the type I orthogneiss to the type IV migmatitein both feldspar domains.In K-feldspar-rich aggregates, the degree of GBPOof the K-feldspar like–like boundaries slightly decreasesfrom type I orthogneiss to type III migmatite,whereas type IV migmatite is characterized by an increasein the degree of K-feldspar like–like GBPO(Fig. 9, Table 1). The GBPO of plagioclase–plagioclaseboundaries in the plagioclase-rich aggregates issimilar to the evolution of K-feldspar like–likeboundaries. The GBPO of the K-feldspar–quartzboundaries as well as those of K-feldspar–plagioclaseboundaries are weak, and decrease throughout thetextural evolution (Fig. 9, Table 1).MINERAL FABRICIn rocks deformed in the presence of melt, the texturesof quartz and feldspar can be used to evaluate thedeformation mechanisms of the solid fraction as wellas the deformation of crystallizing intragranular melt(Za´vada et al., 2007). The mineral fabrics of ferromagnesianphases can be indirectly assessed usinganisotropy of magnetic susceptibility (AMS). TheAMS method has been recently used to determine thedegree of susceptibility, shape of the fabric ellipsoidand relative contribution of ferro- and para-magneticminerals to the bulk fabric in migmatites (Ferre´ et al.,2003, 2004).Anisotropy of magnetic susceptibilityTypes III and IV migmatites are macroscopically closeto isotropic, so that the mineral alignment defined bythe orientation of dispersed biotite is poorly defined(Fig. 2c, d). To better characterize the fabric, the AMSmethod was used to determine the internal fabric ofthese rocks. Oriented samples were collected using aportable drill at four sampling sites covering a sectionacross the well-defined structural sequence. The AMSdata were statistically evaluated using the Anisoftsoftware package (Jelínek, 1978; Hrouda et al., 1990).The low values of mean susceptibility (

ORIGIN OF FELSIC MIGMATITES 43(a)(b)Fig. 10. Plots to show the anisotropy of magnetic susceptibility (AMS). (a) P¢–T plot, where the P¢ parameter represents the degree ofmagnetic anisotropy and T is a shape parameter that describes the shape of the ellipsoid of magnetic susceptibility. T can take eitherpositive values (T > 0), characteristic for a planar fabric, or negative values (T < 0), typical for a linear fabric. Dashed ellipses showtwo distinct datasets. For comparison, data obtained by Schulmann K., Edel J.-B., Hasalova´ P., Lexa O., Jezˇek J. & Cosgrove J. W.(unpublished data) and Bouchez (1997) are shown. (b) Magnetic foliation (circles), plotted as the minimal susceptibility direction (K 3 ),perpendicular to the magnetic foliation, and magnetic lineation (squares), plotted as the maximal susceptibility direction (K 1 ).Lattice preferred orientationTo understand the deformation behaviour of individualphases, we measured and evaluated statistically thelattice preferred orientation (LPO) of aggregate grains(Pl, Kfs and Qtz) and grains apparently crystallizedfrom melt (Pl, Kfs and Qtz) separately (Fig. 12g, h).The LPO of quartz, plagioclase and K-feldspar weremeasured on a scanning electron microscope Cam-Scan3200 in the Czech Geological Survey using theelectron back-scattered diffraction technique (EBSD)and HKL technology (Adams et al., 1993; Bascouet al., 2001). Diffraction patterns were acquired at20 kV of accelerating voltage, 5 nA of probe currentand working distance of 33 mm from the thin sectionprepared from the structural XZ plane. The procedurewas carried out manually due to small differences indiffraction patterns. The chemistry and orientation ofindividual grains was controlled using a forescatterdetector with combination of orientation and chemicalcontrast. Thus, each individual grain is represented byonly one orientation measurement. The resulting polefigures are presented as lower hemisphere equal-areaprojections in which the trace of foliation is orientedalong the equator and the stretching lineation is in theE–W direction.Old quartz grains in ribbons of the type I orthogneissshow c-axes distributed in weak sub-maxima arrangedalong weakly developed small circles close tothe S 1 foliation trace. The most intense sub-maximaare developed close to the lineation direction. This typeof c-axis pattern may indicate preferential prism Æcæslip-system activity and dominantly coaxial deformation.The c-axes of large quartz grains in types II, IIIand IV migmatites reveal strong maxima either parallelto the S 2 foliation pole or close to the centre of thediagram. These c-axis patterns indicate mainly activityof basal Æaæ or rhomb Æa + cæ slip-systems and lessfrequently prism Æaæ slip (Fig. 11a). Towards types IIIand IV migmatites, the LPO of the matrix quartz becameless well developed, preserving activity of thesame slip-systems as in the previous microstructuraltypes (Fig. 11a). New quartz grains crystallized frommelt in type I orthogneiss, and type II migmatite showvery weak LPO and nearly random distribution of allquartz axes (Fig. 11b). Whereas old grains show progressiveweakening of the LPO from type II to type IVmigmatite, the new and randomly crystallized grainstends to develop weak crystal preferred orientationduring the same microstructural evolution from type IIto type IV migmatite (Fig. 11). It is difficult to distinguishold from new quartz grains in the type IV rockand therefore the LPO of quartz in this microstructurelinks LPO evolution between old and new grains in thefinal microstructural type.K-feldspar and plagioclase commonly show weakLPO in all rock types regardless the origin of grains.K-feldspar shows crystallographic patterns which arecompatible with dominant activity of the 1/2 [110](001)slip system (Willaime & Gandais, 1977; Willaime et al.,1979) (Fig. 12a, c). Contribution of other slip systemsas [100](010) (Fig. 12b) and [100](001) (Fig. 12d) hasalso been identified in both relict K-feldspar grains andin K-feldspar grains apparently crystallized from meltrespectively.Distribution of the main lattice directions of plagioclaserevealed slip parallel either to 1/2[1 10] on (001)and (11 1) planes (Fig. 12e) or to 1/2 [110] on (001) and(1 1 1) planes (Fig. 12f) for all types of rocks and bothaggregate grains and plagioclase inferred to havecrystallized from melt (Fig. 12g) (Olsen & Kohlstedt,1984). The textures of plagioclase inferred to haveÓ 2007 Blackwell Publishing Ltd329

44 P. HASALOVÁ ET AL.(a)(b)Fig. 11. Characteristic c-axes preferred orientations of (a) old/relict quartz grains and (b) new quartz grains crystallized from areas ofinferred former melt for all rock types. The c-axis patterns of old/relict quartz grains in type I banded orthogneiss indicate prism Æcæslip system activity whereas in type II, III and IV migmatites basal Æaæ or rhomb Æa + cæ slip systems are dominant with minorprism Æaæ slip. New quartz grains inferred to have crystallized from melt in type I banded orthogneiss to type IV nebulitic migmatiteshow very weak LPO and nearly isotropic distribution of all quartz axes. Equal area projections, lower hemisphere, contoured atinterval of 0.5 times uniform distribution. Foliation is horizontal and lineation is in this plane in the E–W direction. N is the number ofmeasured grains. Maximum densities are marked on the bottom right of each pole figure. The dashed line represents the lowest contourlevel and the grey circle corresponds to the minimum density value.crystallized from melt are commonly weak with theexception of strong LPO of plagioclase in the type Iorthogneiss (Fig. 12f). Such slip-systems are supposedto be secondary and active if grains are in unsuitable(hard) orientation to the dominant slip-system[100](010) (Kruse et al., 2001).DISCUSSIONThis study presents a detailed microstructural andquantitative textural analysis of four types of migmatiticrocks identified in one of the largest (5000 km 2 )migmatitic complex of the eastern Variscan belt. Therock types are interpreted as representing a texturalsequence from banded orthogneiss via stromatic andschlieren migmatites to nebulitic migmatite. The possiblemechanisms that could account for the origin ofthis rock sequence involve: (i) genetically unrelatedmigmatites that have originated from distinct protoliths;(ii) variable degree of in situ partial melting of asingle protolith or different protoliths; and (iii) meltinfiltration from an external source through solid rockin which banded orthogneiss and nebulitic migmatiterepresent genetically linked end-members. Thesehypotheses are discussed further below.Spatial relationships of individual migmatite types withinthe shear zoneThe structural sequence described in this work indicatesan intimate relationship between types I to IIImigmatites and nebulitic type IV migmatite sheets thatcan be interpreted in terms of a shear zone, which wasexploited by rising magma (Brown et al., 1995; Collins& Sawyer, 1996; Brown & Solar, 1998b). We haveshown that the D 2 flat fabrics that cross-cut the steepfoliation S 1 developed at high-temperature solid-stateconditions (Fig. 1). Tajcˇmanova´ et al. (2006) andRacek et al. (2006) described a similar sequence ofsuperposed fabrics in lower crustal rocks several tensof kilometres to the north and south of the studied arearespectively. These authors proposed that the flat D 2deformation fabrics originated due to thrusting oforogenic lower crust over middle crustal units along alarge-scale retrograde shear zone. In agreement withthese authors, we suggest that the D 2 fabrics developedin a thrust related crustal-scale shear zone reportedalready by Urban (1992), Schulmann et al. (1994) andredefined later by Schulmann et al. (2005). The maindifference between other regions is in the degree of D 2reworking, which is so high in the studied area thatÓ 2007 Blackwell Publishing Ltd330

ORIGIN OF FELSIC MIGMATITES 45Fig. 12. Characteristic LPO patterns of K-feldspar (a–d) and plagioclase (e,f). Both feldspars commonly show weak LPO in all rocktypes regardless of the origin of the grains. An exception is the strong LPO of new plagioclase grains in the type I banded orthogneiss(f). K-feldspar usually shows activity of 1/2[110](001) (a,c; type I), but also of [100](010) (b; type IV) and [100](001) (d; type IV) slipsystems. Plagioclase reveals activity of secondary slip systems such as 1/2[1 10] on (001) and (11 1) or 1/2 [1 10] on (001) and (1 1 1)(e,f; type I). Equal area projections, lower hemisphere, contoured at intervals of 0.5 times uniform distribution. Foliation is horizontaland lineation is in this plane in the E–W direction. N is the number of measured grains. Maximum densities are marked on thebottom right of each pole figure. The dashed line represents the lowest contour level and the grey circle corresponds to the minimumdensity value. (f,g) BSE images depicting the microstructural appearance of examples of the measured phases (sample PH60/B).Original plagioclase (g) and K-feldspar (h) aggregates with newly crystallized quartz (white arrow), plagioclase (black arrows) andK-feldspar (grey arrows) are shown.Ó 2007 Blackwell Publishing Ltd331

46 P. HASALOVÁ ET AL.steep D 1 fabrics are preserved only as rare relics shownin Fig. 2.Our structural observations are compatible withprogressive transposition within the ductile shear zoneranging from type I banded orthogneiss to highlyreworked type III schlieren migmatite. This interpretationis based on progressive folding of early steeporthogneiss fabric, development of isoclinal folds andcomplete fabric transposition and development of typeIII migmatite in a ductile shear zone (Fig. 2; Turner &Weiss, 1963). In such a context, the elongated bodies oftype IV nebulitic migmatite can be seen as veins ofisotropic granite penetrating parallel to the main S 2mylonitic anisotropy (e.g. Cosgrove, 1997; Brown &Solar, 1998a). Alternatively, the type IV nebuliticmigmatites could be interpreted in terms of injectedmelt into hot country rocks (called also magmawedging, Weinberg & Searle, 1998) preventing magmafreezing during D 2 shearing. Finally, the nebuliticmigmatite can be also regarded as the most extremeend-member of the structural sequence, i.e. completelydisintegrated parental orthogneiss.In summary, the type I orthogneiss to type III migmatiteshow intimate spatial relationships suggestingthat they have originated from the same protolith andthat they are genetically linked. However, the macroscopicobservations alone cannot distinguish the originof type IV migmatite and further arguments arerequired.Microstructural and petrological arguments for melt–rockinteraction during exhumationWe suggest, in agreement with Sawyer (1999, 2001),that the position and topology of new plagioclase,quartz and K-feldspar grains in type I orthogneiss andtype II migmatite may be interpreted in terms of meltproducts crystallized along boundaries of the feldsparin individual aggregates (Fig. 5a–c). The main differencebetween type I orthogneiss and type II migmatiteis a more albitic composition of plagioclase and agreater modal content of new phases in the latter.Types III and IV migmatites show development ofhighly corroded shapes of K-feldspar, plagioclase andbiotite (Fig. 4c–f). This indicates that all rock typesexhibit features compatible with the presence of meltand its interaction with the solid rock. Additionally,the degree of melt–rock interaction is inferred toincrease from type I orthogneiss towards type IVnebulitic migmatite (Fig. 4).The structural sequence exhibits a distinct trend inmodal composition of originally monomineralic layersthat are progressively converted into polymineralicaggregates of granitic composition (Fig. 6). The compositionalpaths show evolutionary trends from type Iorthogneiss to type III migmatite, interpreted as beingassociated with crystallization of melt, culminating inthe type IV migmatite, which has equal amounts ofplagioclase, K-feldspar and quartz (Fig. 6).Plagioclase shows systematic decrease in anorthitecontent for both original plagioclase grains (An 30 toAn 25 ), their rims and inter-granular aggregates (An 20to An 10 ) towards type IV nebulitic migmatite. Bothgarnet and biotite exhibit systematically increasing X Fetowards type IV migmatite (from 0.7 to 1.00 and from0.4 to 0.9 respectively) coupled with decrease in Ticontent in biotite (from 0.2 to 0.04 p.f.u.). The mineralcompositional data suggest systematic equilibration ofgarnet, biotite and plagioclase compositions in thestability field of sillimanite with decreasing temperature.The full petrological data and P–T estimates formelt–rock interaction are presented in a companionpaper (Hasalova´ et al., 2008a). Here, we quote the P–Testimates based on thermodynamic modelling usingTHERMOCALC (Powell et al., 1998) to point out thedecrease in temperature from 790 to 850 °C at 7.5 kbarfor type I orthogneiss to 690–770 °C at 4.5 kbar fortype IV nebulitic migmatite.Taken together, the microstructural and petrologicaldata show paradoxically an increasing degree ofapparent melt–rock interaction coupled with decreasingequilibration temperature with textural evolutionfrom type I to type IV rock type. The microstructuraland modal composition data do not exclude eitherpartial melting or infiltration of melt from an externalsource. However, the systematic modification ofchemical composition of minerals across the migmatitesequence is in contradiction with the model of in situpartial melting. Namely, as suggested by many fieldand experimental studies, the composition of plagioclasewould be shifted towards more anorthitic contentsand the X Fe of garnet and biotite would decreaseduring partial melting process (Le Breton & Thompson,1988; Vielzeuf & Holloway, 1988; Gardien et al.,1995; Greenfield et al., 1998; Dallain et al., 1999).Interpretation of quantitative microstructural dataMicrostructural studies of partially molten rock haverevealed systematic changes in grain size, grain SPOand spatial distribution of individual phases duringincreasing degree of partial melting (e.g. Vernon, 1976;McLellan, 1983; Dallain et al., 1999). Here thesetrends are compared with our quantitative microstructuraldata, and the alternative origins that mayresult in the observed microstructural sequence arediscussed.Interpretation of crystal size distributionsThe most significant result of this study is the systematicdecrease in average grain size (Fig. 7a) andsystematic increase in a population density (nucleationrate) associated with possible decrease in growth ratefor all feldspar from type I orthogneiss to type IVmigmatite (Fig. 7b, c).Results from migmatitic terranes show that the CSDassociated with partial melting is characterized byÓ 2007 Blackwell Publishing Ltd332

ORIGIN OF FELSIC MIGMATITES 47production of coarse-grained felsic mineral aggregatesresulting from increase in temperature (e.g. Dougan,1983; McLellan, 1983). This process is commonly followedby textural coarsening (Ashworth & McLellan,1985; Dallain et al., 1999; Berger & Roselle, 2001)explained by two competing approaches: the Lifshitz–Slyozov–Wagner (LSW) model (Lifshitz & Slyozov,1961), and the communicating neighbour theory (CNof DeHoff, 1991). Higgins (1998) showed that texturalcoarsening results in progressive decrease in N 0 valueand decrease in the slope of the CSD curve; he interpretedthis trend as a result of rapid undercoolingduring solidification of magma followed by reducedundercooling, suppression of nucleation and texturalcoarsening. However, our textural sequence exhibitsthe opposite trend in evolution of CSD curves, which isinterpreted as indicating that in situ partial melting andtextural coarsening are not responsible for the origin ofobserved CSDs.The observed CSD trend may be explained by one ofthree different mechanisms: (1) solid-state deformationunder decreasing temperature and or increasing strainrate (Hickey & Bell, 1996; Azpiroz & Ferna´ndez, 2003;Lexa et al., 2005); (2) a different degree of reactionoverstepping (Waters & Lovegrove, 2002; Moazzen &Modjarrad, 2005); and (3) a different degree of undercooling(Marsh, 1988).The grain size for dynamically recrystallized grainsin a power-law creep regime is a function of differentialstress (Twiss, 1977). Such grains are characterized bystrong shape and LPO and commonly solid-state differentiation(Baratoux et al., 2005; Lexa et al., 2005).However, this microstructural study does not revealany features in quartz, plagioclase and K-feldspar ofrock types II, III and IV which may indicate a dynamicrecrystallization processes operating under decreasingtemperature. Differences in the degree of reactionoverstepping have been documented in contact aureoles,but may be rejected in this case due to theregional nature of the metamorphism. However, therole of different degrees of undercooling relating to anoverall decrease in equilibration temperature cannot beexcluded.Our data indicate that the sequence of rock typesreflects the progressive resorption of residual grainsand crystallization of new grains from melt in intergranularspaces. Moreover, the trend of CSD curvessuggest a progressive increase in nucleation rate anddecrease in growth rate from type I orthogneiss totype IV nebulitic migmatite. This trend could beexplained by an increase in undercooling consistentwith the decreasing equilibration temperature wereport.The CSD trend is compatible with crystallization ofmelt in a progressively exhuming and rapidly coolingsystem. This is in accordance with exceptionally highcooling rates up to several hundred degrees celsius permillion years estimated for nearby granulites byTajcˇmanova´ et al. (2006).Interpretation of spatial distributions of phasesThe quantitative analysis of spatial distributions ofindividual phases shows that the intensification ofregular distribution (increasing amount of unlikecontacts; Fig. 9) correlates with an increasing degree ofhost rock–melt equilibration. The process of meltcrystallization leads to new mineral growth on thesurfaces of residual grains. This is responsible for theincrease in unlike grain boundaries, which commonlyretain melt–solid geometries. Our case study showsthat the development of a regular distribution of felsicphases is not related to solid-state annealing, as supposedby some authors (Flinn, 1969; McLellan, 1983;Lexa et al., 2005), but to the process of crystallizationof melt, consistent with precipitation of the minorphase on triple points in granular polygonal aggregatesto achieve lower total interfacial energy (Spry, 1969;Vernon, 1974). This process was documented by Dallainet al. (1999), who showed that the predominanceof unlike contacts in polycrystalline aggregates originatedthrough wetting of grain boundaries by fluids ormelt, and subsequent precipitation of other phases onlike–like contacts. However, we cannot exclude thepossibility that a regular distribution reported fromgranulites and high-grade gneisses (Flinn, 1969; Kretz,1994) results from solid-state annealing of rocks wheremelt crystallized. Therefore, the regular distributiondeveloped during melt crystallization may be inheritedand perhaps further accentuated during later thermaland textural re-equilibration.Origin of microstructural and compositional trendsThe sequence from type I orthogneiss to type IVmigmatite exhibit continuous trends in all quantitativeparameters (Table 1). The grain size decreases(Fig. 7a) and there is a progressive development of aregular distribution of all felsic phases (Fig. 9), whichis linked with mineral compositional trends indicatingtemperature decrease. These clear evolutionary trendsare incompatible with a process of partial melting ofdifferent protoliths. Partial melting of the same protolithmay develop continuous trends, but these shouldshow increase in grain size of individual felsic phases(Dallain et al., 1999) and different mineral compositionalevolution (e.g. Gardien et al., 1995; Greenfieldet al., 1998). Additionally, we show that the degree ofregular distribution for K-feldspar- and plagioclasedominatedaggregates evolves in the same mannerthroughout the microstructural sequence (Fig. 9).However, Dallain et al. (1999) reported significantlymore advanced regular distribution of plagioclasecomparedwith K-feldspar-rich aggregates in themicrostructural sequence originated by partial melting.These authors proposed that this microstructuralcontrast originated due to melting process preferentiallyoperating in mica–plagioclase rich aggregates,whereas the K-feldspar-rich aggregates were moreÓ 2007 Blackwell Publishing Ltd333

48 P. HASALOVÁ ET AL.refractory. In the present case, Hasalova´ et al. (2008b)report continuous trends in whole-rock geochemistryand mineral compositions for the sequence of rocktypes, but different Nd isotopic composition for thetype I orthogneiss compared with the rest of the sequence,which precludes of in situ anatexis in a closedsystem.Melt infiltration modelThe discrepancies between the evolutionary trends wereport and generally accepted trends for anatecticterranes require an appropriate explanation that isconsistent with the structural, quantitative microstructuraland mineral compositional data. As a possibleexplanation, we introduce the concept of meltinfiltration from an external source, where melt passespervasively along grain boundaries through the wholerockvolume and changes macroscopic (Fig. 2) andmicroscopic (Fig. 3) appearance of the rock. Thisprocess is characterized by resorption of old phases,nucleation of new phases along high-energy like–likegrain boundaries and modification of mineral andwhole-rock compositions. These gradual changes areaccompanied by grain size reduction (Fig. 7) andprogressive disintegration of former aggregate (layered)distribution of original phases (Fig. 9). We suggestthat the individual migmatite types representdifferent degrees of equilibration between the host rockand migrating melt. It should be emphasized, that allthese processes occur along a retrograde path duringexhumation of the Gfo¨ hl Unit. We are aware that adecrease in P–T conditions during melt infiltration is afundamental and limiting factor for the model proposed.The amount of melt and its connectivity are criticalparameters controlling melt mobility and the rheologicalbehaviour of melt-present rocks. To constrainthese parameters both AMS and EBSD were used.Using AMS, it is possible to distinguish between solidstatedominated deformation mechanisms in themelanosome and free rigid body particle rotation in theleucosome (e.g. Ferré et al., 2003). On the other hand,using the EBSD technique enables us to distinguishdeformation mechanisms in the solid framework andto constrain the mechanical role of melt during thedeformation.AMS fabric origin: solid framework or melt controlleddeformationThe AMS study shows that the magnetic anisotropy isdominated by biotite. The oblate shape of magneticellipsoid and high degree of anisotropy of type I orthogneissand type II migmatite (Fig. 10a) are consistentwith strong preferred orientation of biotite and thefact that biotite has a intrinsically oblate shape of thesingle-grain magnetic ellipsoid (Zapletal, 1990; Martı´n-Hernandéz & Hirth, 2003). The type III and IV migmatitesreveal partly resorbed biotite flakes uniformlydispersed in the rock marked by slightly weaker degreeof magnetic anisotropy and less oblate fabric ellipsoidcompared with types I and II migmatite (Fig. 10a).This contrasts with common granites and diatexitesfrom other migmatitic terranes which show significantlylower values of degree of anisotropy and highlyvariable shapes of AMS ellipsoids (Fig. 10a; Bouchez,1997; Ferre´ et al., 2003).Numerous natural studies supported by numericalmodelling indicate that the magnetic susceptibility inviscously flowing magmas is characterized by a verylow degree of anisotropy, pulsatory fabrics and dominantlya plane strain AMS ellipsoid shape (Blumenfeld& Bouchez, 1988; Hrouda et al., 1994; Arbaretet al., 2000). A comparison of the AMS fabrics withthose of diatexites and results of numerical modelsindicate that the intensity of the AMS fabric of typesIII and IV migmatites does not originated from freelyrotated biotite in viscously flowing melt. On the contrary,we argue that the AMS fabric in all types ofmigmatites resembles fabrics usually acquired throughsolid-state deformation of a load-bearing framework,similar to melanosomes in migmatites (Ferre´ et al.,2003). To understand the mechanisms responsible fordevelopment of such fabrics the grain-scale deformationmechanisms and melt behaviour in individual rocktypes is discussed.Deformation mechanismsExperimental studies of low melt fraction rocks deformedunder high differential stress show that matrixminerals deform by grain boundary migrationaccommodated dislocation creep (DellÕ Angelo et al.,1987; Walte et al., 2005). Strong shape and GBPO offeldspar (Figs 8 & 9) as well as LPO of residual quartzgrains (Fig. 11a) in the type I orthogneiss may beinterpreted in terms of plastic deformation consistentwith a dislocation creep deformation mechanism(Rosenberg & Berger, 2001). However, the weak LPOof residual grains of both feldspars (Fig. 12a, e) in thetype I orthogneiss suggests a contribution of grainboundary sliding during the development of themicrostructure. In other words, the type I microstructurecorresponds to a transient microstructure interms of decreasing activity of dislocation creep andenhancement of diffusion controlled processes.Decrease in SPO and GBPO and constantly weakLPO in feldspar of type II migmatite (Figs 8 & 9) maybe interpreted as a result of melt-enhanced diffusioncreep (Garlick & Gromet, 2004). However, the largequartz grains reveal intense activity of basal Æaæ slipsuggesting important plastic yielding of this mineral(Fig. 11a). Elongate pockets inferred to represent formermelt oriented at a high angle to the stretchinglineation in the type I orthogneiss (Fig. 5a) and type IImigmatite indicate that the melt distribution wascontrolled by the deformation. This is supported by theÓ 2007 Blackwell Publishing Ltd334

ORIGIN OF FELSIC MIGMATITES 49strong LPO of interstitial plagioclase (Fig. 12f).Rosenberg & Riller (2000) reported that pockets withinquartz aggregates inferred to have been former meltare oriented at high angle to the foliation plane, possiblyclose to r 1 . Melt distribution in our samples issimilar to their results and also to experiments at highdifferential stresses (DellÕ Angelo & Tullis, 1988) andhigh confining pressures. In these experiments, meltaccumulated in pockets along faces of the grains subparallelto the main compressional stress direction r 1(DellÕ Angelo & Tullis, 1988; Daines & Kohlstedt,1997). Such a melt topology is also termed ÔdynamicwettingÕ (Jin et al., 1994).In types III and IV migmatites both residual andnew grains of K-feldspar and plagioclase exhibit lowSPO, GBPO and LPO (Figs 8, 9 & 12), indicatingabsence of dislocation creep, in contrast to quartz,which exhibits relatively strong crystallographic preferredorientation (Fig. 11a). The topology of formermelt is poorly constrained in both types III and IVmigmatite but, the SPO of the minor phases interpretedto have crystallized from melt shows a bimodaldistribution sub-perpendicular and sub-parallel to theS 2 foliation. These observations are neither compatiblewith high differential stress nor low differential stressexperiments, in which the melt occurs primarily intriple point junctions without any SPO (DellÕ Angeloet al., 1987; Gleason et al., 1999). However, in somenatural samples, former melt pockets are preferentiallylocated along grain boundaries parallel to the foliation(John & Stu¨ nitz, 1997; Sawyer, 1999; Rosenberg &Berger, 2001), indicating that the orientation of meltpocket in nature is not always in agreement withexperimental studies (Rosenberg, 2001). In this study,the melt pocket orientation sub-parallel to the foliationmay indicate low differential stress and high fluid/meltpressure as suggested by Cosgrove (1997).We conclude that during evolution from type Ibanded orthogneiss to type IV nebulitic migmatite meltwetted a majority of grain contacts. The AMS studyand quartz microfabrics in types II to IV migmatitessuggest that the melt fraction did not exceed the criticalamount to allow free relative movement of grainswithout interference, i.e. the melt fraction is below thecritical threshold (e.g. RCMP of Arzi, 1978; RPT ofVigneresse et al., 1996). Rosenberg & Handy (2005)argued that melt fractions of only / ¼ 0.07 (meltconnectivity threshold, MCT) will enable the formationof interconnected networks of melt under dynamicconditions which will lead to a substantial strengthdrop. These authors suggested that weakening at theMCT probably involves localized, inter- and intragranularmicrocracking, as well as limited rigid bodyrotation of grains, without an important contributionof dislocation creep and diffusion processes at grainboundaries. However, we do not observe any strainlocalization associated with brittle failure and thereforeit is suggested that the deformation has to beaccommodated by mechanisms operating homogeneouslyacross significant rocks volumes. Materialscience experiments (Mabuchi et al., 1997) show thatweakening due to melt-enhanced grain boundary slidingat low melt fraction is an efficient mechanismallowing homogeneous deformation. We suggest thatdeformation of both feldspars and quartz in the type IIto type IV migmatites occurred by melt-enhanced grainboundary sliding with a contribution to the overalldeformation by dislocation creep. These characteristicsare compatible with granular flow as described byPaterson (2001) accompanied by melt-enhanced diffusionand/or direct melt flow.CONCLUSIONSBased on a detail field and microstructural study, wedistinguish four types of gneiss/migmatite in the Gfo¨ hlgneiss complex: (i) banded orthogneiss (type I), withdistinct layers of recrystallized plagioclase, K-feldsparand quartz separated by layers of biotite; (ii) stromaticmigmatite (type II), composed of plagioclase and K-feldspar aggregates with subordinate quartz andirregular quartz aggregates – the boundaries betweenindividual aggregates are ill defined and rather diffuse;(iii) schlieren migmatite (type III), which consists ofplagioclase–quartz- and K-feldspar–quartz-enricheddomains with a foliation marked only by preferredorientation of biotite and sillimanite dispersed in therock; and, (iv) nebulitic migmatite (type IV), with norelicts of gneissosity. It is demonstrated that this is acontinuous sequence developed by melt-presentdeformation, in which the type I banded orthogneissesand type IV nebulitic migmatites are end-members.The progressive disintegration of the bandedmicrostructure and the development of nebuliticmigmatite is characterized by several systematic texturalchanges. The grain size of all felsic phasescontinuously decrease whereas the population densityof precipitated phases increases. The new phasespreferentially nucleate along high-energy like–likeboundaries, causing the development of a regular distributionof individual phases. Simultaneously, themodal proportions of felsic phases evolve towards aÔgranite minimumÕ composition. Further, this evolutionarytrend is accompanied by a decrease in grainSPO of all felsic phases. To explain these textural andcompositional changes we introduce a model of meltinfiltration from an external source in which melt isargued to pass pervasively along grain boundariesthrough the whole-rock volume. It is suggested that theindividual migmatite types represent different degreesof equilibration between the host rock and migratingmelt during the retrograde metamorphic evolution.The inferred melt topology in type I orthogneissexhibits elongated pockets of melt oriented at a highangle to the compositional banding, indicating that themelt distribution was controlled by deformation thesolid framework. Here, the microstructure exhibitsfeatures compatible with a combination of dislocationÓ 2007 Blackwell Publishing Ltd335

50 P. HASALOVÁ ET AL.creep and grain boundary sliding deformation mechanisms.The types II–IV microstructures developed bygranular flow accompanied by melt-enhanced diffusionand/or melt flow. However, the amount of melt presentnever exceeded a critical threshold during the deformationto allow free rotation of biotite grains.The model of melt infiltration based on structuraland microstructural observation is supported by thermodynamic(Hasalova´ et al., 2008a) and geochemicalmodelling (Hasalova´ et al., 2008b). Although our dataseem to be consistent with such a model, there are stilla number of issues to be resolved (e.g. time-scale of theprocess, the character of the melt and the grain-scaledeformation mechanisms enabling pervasive flow ofviscous melt). Nevertheless, our model has profoundconsequences for the petrogenesis of migmatites, therheology of anatectic regions during syn-orogenicexhumation and melt transport in the crust.ACKNOWLEDGEMENTSThis work was financially supported by the GrantAgency of the Czech Academy of Science (Grants No.IAA311140 and No. 205/04/2065) by an internalCNRS funding (CGS/EOST) and by the FrenchNational Agency (No. 06-1148784). Visits by P.Hasalova´ to ULP Strasbourg were funded by theFrench Government Foundation (BGF). We gratefullyacknowledge A. Langrova´ from the Institute ofGeology at the Czech Academy of Sciences for operatingthe microprobe. We also thank the reviewersC. Rosenberg, R. Weinberg and H. Stu¨ nitz for theirconstructive comments and suggestions for improvingthis paper, and the Journal editor M. Brown for hiscareful handling of the manuscript. R. Powell andM. Brown are thanked for helpful discussions.REFERENCESAdams, B. L., Wright, S. I. & Kunze, K., 1993. 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ORIGIN OF FELSIC MIGMATITES 53eastern margin of the Bohemian Massif (Thaya Dome).Journal of Structural Geology, 16, 355–370.Schulmann, K., Kro¨ ner, A., Hegner, E. et al., 2005. Chronologicalconstraints on the pre-orogenic history, burial andexhumation of deep-seated rocks along the eastern margin ofthe Variscan orogen, Bohemian Massif, Czech Republic.American Journal of Science, 305, 407–448.Seng, H., 1936. Die Migmatitfrage und der Mechanismus parakristallinerPrägung. Geologische Rundschau, 27, 471–492.Spry, A. H., 1969. Metamorphic Textures. Pergamon Press,Oxford.Sˇtı´pska´, P., Schulmann, K. & Kro¨ ner, A., 2004. Vertical extrusionand middle crustal spreading of omphacite granulite: amodel of syn-convergent exhumation (Bohemian Massif,Czech Republic). Journal of Metamorphic Geology, 22, 179–198.Stu¨ nitz, H., 1998. Syndeformational recrystallization – dynamicor compositionally induced? Contributions to Mineralogy andPetrology, 131, 219–236.Tajcˇmanova´, L., Konopa´sek, J. & Schulmann, K., 2006. Thermalevolution of the orogenic lower crust during exhumation withina thickened Moldanubian root of the Variscan belt of CentralEurope. Journal of Metamorphic Geology, 24, 119–134.Tanner, D. C., 1999. The scale-invariant nature of migmatitefrom the Oberpfalz, NE Bavaria and its significance for melttransport. Tectonophysics, 302, 297–305.Turner, F. J. & Weiss, L., 1963. Structural Analysis of MetamorphicTectonites. McGraw Hill, New York.Twiss, R. J., 1977. Theory and applicability of a recrystallizedgrain size paleopiezometer. Pure and Applied Geophysics, 115,227–244.Urban, M., 1992. Kinematics of the Variscan thrusting in theEastern Moldanubicum (Bohemian Massif, Czechoslovakia):evidence from the Na´meˇsˇtÕ granulite massif. Tectonophysics,201, 371–391.Vanderhaeghe, O., 1999. Pervasive melt migration from migmatitesto leucogranite in the Shuswap metamorphic corecomplex, Canada: control of regional deformation. Tectonophysics,312, 35–55.Vernon, R. H., 1974. Controls of mylonitic compositional layeringduring non-cataclastic ductile deformation. GeologicalMagazine, 111, 121–123.Vernon, R. H., 1976. Metamorphic Processes: Reactions andMicrostructure Development. Wiley, New York.Vernon, R. H. & Paterson, S. R., 2001. Axial-surface leucosomesin anatectic migmatites. Tectonophysics, 335, 183–192.Vielzeuf, D. & Holloway, J. R., 1988. Experimental determinationof the fluid-absent melting relations in the pelitic systemconsequencesfor crustal differentiation. Contributions toMineralogy and Petrology, 98, 257–276.Vigneresse, J.-L., Barbey, P. & Cuney, M., 1996. Rheologicaltransitions during partial melting and crystallization withapplication to felsic magma segregation and transfer. Journalof Petrology, 37, 1579–1600.Walte, N. P., Bons, P. D., Passchier, C. W. & Koehn, D., 2003.Disequilibrium melt distribution during static recrystallization.Geology, 31, 1009–1012.Walte, N. P., Bons, P. D. & Passchier, C. W., 2005. Deformationof melt-bearing systems – insight from in situ grain-scaleanalogue experiments. Journal of Structural Geology, 27,1666–1679.Waters, D. J. & Lovegrove, D. P., 2002. Assessing the extent ofdisequilibrium and overstepping of prograde metamorphicreactions in metapelites from the Bushveld aureole. Journal ofMetamorphic Geology, 20, 135–149.Weinberg, R. F. & Searle, M. P., 1998. The Pangong InjectionComplex, Indian Karakoram: a case of pervasive granite flowthrough hot viscous crust. Journal of Geological Society, 155,883–891.Wickham, S. M., 1987. The segregation and emplacement ofgranitic magmas. Journal of the Geological Society of London,144, 281–297.Willaime, C. & Gandais, M., 1977. Electron microscope study ofplastic defects in experimentally deformed alkali feldspars.Bulletin de la Societe Francaise de Mineralogie et de Cristallographie,100, 263–271.Willaime, C., Christie, J. M. & Kovacs, M.-P., 1979. Experimentaldeformation of K-feldspar single crystals. Bulletin deMineralogie, 102, 168–177.Zapletal, K., 1990. Low-field susceptibility anisotropy of somebiotite crystals. Physics of the Earth and Planetary Interiors,63, 85–97.Za´vada, P., Schulmann, K., Konopa´sek, J., Ulrich, S. & Lexa,O., 2007. Extreme ductility of feldspar aggregates – meltenhanced grain boundary sliding (GBS) and creep failure:rheological implications for felsic lower crust. Journal ofGeophysical Research, doi: 10.1029/2006JB004820.2007Received 14 February 2006; revision accepted 25 September 2007.SUPPLEMENTARY MATERIALThe following material is available online at http://www.blackwell-synergy.com:Appendix S1 Digitalized thin sections (separatedplagioclase and K-feldspar domains) that have beenused for the quantitative analyses.Ó 2007 Blackwell Publishing Ltd339

J. metamorphic Geol., 2011, 29, 103–130 doi:10.1111/j.1525-1314.2010.00911.xOrigin of felsic granulite microstructure by heterogeneousdecomposition of alkali feldspar and extreme weakening oforogenic lower crust during the Variscan orogenyJ. FRANĚK, 1,2 K. SCHULMANN, 2 O. LEXA, 1,3 S. ULRICH, 4 P. ŠTÍPSKÁ, 2 J. HALODA 1 AND P. TÝCOVÁ 11 Czech Geological Survey, Klárov 3, 118 21 Prague, Czech Republic (honzaf2@seznam.cz)2 Institut de Physique du Globe de Strasbourg, IPGS – UMR 7516, CNRS et Université de Strasbourg (EOST), 1 Rue Blessig,67084 Strasbourg, France3 Institute of Petrology and Structural Geology, Charles University, Albertov 6, 128 43 Prague, Czech Republic4 Institute of Geophysics, Czech Academy of Sciences, Boční II⁄ 1401, 141 31 Prague, Czech RepublicABSTRACTThis study answers the question of origin and evolution of a granulitic microstructure typicallydeveloped in felsic granulites of the European Variscan belt. It shows that the precursor of the Variscanfelsic granulites was a high-pressure alkali feldspar-rich coarse-grained layered orthogneiss. Its S1subhorizontal layering is defined by the alignment of alkali feldspar porphyroclasts alternating withmonomineralic bands of quartz and bands rich in plagioclase and garnet. The alkali feldsparporphyroclasts contain inclusions of quartz, garnet, kyanite, biotite and rutile, reflecting peak P–Tconditions of 1.6–1.8 GPa and 850 °C during S1 formation. Superimposed steep folds and steepcleavage, S2, are associated with recrystallization of alkali feldspar, plagioclase and quartz, and garnetchemistry modifications that correspond to 0.9–1.0 GPa and 800 °C. During exhumation, involving0.8 GPa decompression and cooling, the probably perthitic alkali feldspar underwent an unusualprocess of heterogeneous decomposition along irregular reaction fronts forming a fine-grained matrixcomposed of plagioclase and K-feldspar grains. Regular grain distributions in the matrix, nucleationdominatedcrystal size distribution and preservation of lattice orientation of the parental perthitecrystals are all explained by a discontinuous precipitation process. This heterogeneous decomposition ofalkali feldspar solid solution is controlled by chemically and strain induced grain-boundary migration.During exhumation and decompression, the fine-grained matrix underwent viscous deformation,forming the typical microstructure of the Variscan granulites. Random phase distributions, minorcoarsening and feldspar textures are interpreted as a result of strain softening due to diffusion creepaccommodatedgrain-boundary sliding. Subordinate large quartz ribbons were rheologically strongerthan the feldspar-dominated matrix due to the activity of different deformational mechanisms. Finally,in mid-crustal levels, the subvertical structure was overprinted by a perpendicular steep fabric associatedwith the growth of sillimanite, heterogeneous hydration and local partial melting, development ofaggregate phase distributions and significant coarsening. This evolution is accompanied with thedevelopment of a strong lattice preferred orientation of quartz, K-feldspar and plagioclase, reflecting aswitch to dislocation creep mechanism and a general hardening of the granulites under amphibolitefacies conditions.Key words: felsic granulites; Moldanubian domain; quantitative microstructural analysis; quartz andfeldspar rheology; Variscan belt.INTRODUCTIONFelsic high-pressure (HP) granulites represent the mostabundant deep crustal rocks exposed in the PalaeozoicVariscan belt in Europe and form large accumulationsof the orogenic lower crust, e.g. in the Vosges Massif(Gayk & Kleinschrodt, 2000) and the BohemianMassif (OÕBrien & Carswell, 1993), where the classicexample is the Saxony Granulite Massif (Behr, 1961).These rocks are characterized by a fine-grainedrecrystallized matrix of K-feldspar and plagioclasecontaining kyanite, garnet and quartz ribbons. Thesehighly deformed rocks have attracted the interestof structural geologists for decades (e.g. Behr, 1961;Lister & Dornsiepen, 1982), who have studied thequartz textures and considered the granulite microstructurea testimony to deep crustal flow. However,despite the generally accepted opinion that granulitesrepresent deep-seated tectonites reflecting deformationin the deep crust, the parental rocks as well as physicalconditions and mechanism of deformation remainmatters of discussion. Therefore, the key issue of thiswork is understanding the formation of the granulitemicrostructures and related deformation mechanismsÓ 2010 Blackwell Publishing Ltd 103341

104 J. FRANĚK ET AL.in terms of thermal conditions and deformationalprocesses. Unravelling the complex deformationalbehaviour of the felsic granulites should assist broaderconsiderations of rheology and mechanical behaviourof orogenic lower crust during various stages ofdeformation and exhumation.The deformation mechanisms governing rheology oflower crustal rocks are often hard to determine due tochanges of the microstructures during exhumation,recrystallization in later deformations or retrogression(e.g. Rutter & Brodie, 1992). At granulite faciesconditions, the ductile strain is usually accommodatedby dislocation creep, or diffusion creep that may becomplemented by grain-boundary sliding (GBS)(Martelat et al., 1999; Garlick & Gromet, 2004),each mechanism having variable importance. Grainboundarydiffusion or the presence of silicate meltsmay favour the diffusion creep or GBS according tonew results of Schulmann et al. (2008) and Za´vadaet al. (2007). The GBS may then operate at the expenseof other mechanisms and promote granular flow,particularly in fine-grained rocks (Za´vada et al., 2007;Schulmann et al., 2008).Feldspar, the main constituent of the felsic granulites,is an essential component of the EarthÕs crust.Nevertheless, its deformational behaviour is not yetfully understood, mainly due to complex solid solutionmixing and variation in crystallographic structure withcooling (e.g. Ribbe, 1983; Putnis et al., 2003; Abartet al., 2009). A common process is exsolution withcooling, which modifies the rheological behaviour ofalkali feldspar and may lead to drastic weakening ofthe orogenic lower crust (Schulmann et al., 2008).Feldspar recrystallization, driven mainly by chemicaldisequilibrium of the Or–Ab–An solid solution,has been studied in natural examples (Stu¨ nitz, 1998;Putnis, 2002) or experimentally (Stu¨ nitz & Tullis,2001), but these studies have focused mainly onmedium-temperature water-assisted processes below500 °C. At these conditions the original chemicallyunstable feldspar undergoes dissolution and precipitatesas two separate feldspars of different composition.Such results cannot be easily extrapolated to thegranulite water under-saturated HT conditions of atleast 850 °C (OÕBrien & Ro¨ tzler, 2003; Sˇtı´pska´ &Powell, 2005), where the available fluid is representedby silicate melt, rather than water.In order to address the above-mentioned aspects, wepresent a microstructural analysis of exceptionallypreserved samples of lower-crustal felsic granulites froman 8.5 · 2.5 km domain (Franeˇk et al., 2006, 2011) withwell-preserved granulite facies fabrics. These rocks formpart of the Blansky´ les Granulite Massif (BLG) inthe southern Bohemian Moldanubian domain, whichbelongs to the Variscan collisional chain in centralEurope. The combined microstructural and petrologicalanalysis shows evidence of a complex evolution ofalkali feldspar rheology during granulite formation andexhumation. Changes in ductility are ascribed tochemically and deformationally driven recrystallization,variations in grain size as well as changingtemperature and the amount of interstitial melt.Geological settingThe Bohemian Massif (Fig. 1a,b) represents the easternexposure of the Variscan orogen in Europe. Duringthe Variscan orogenesis (380–300 Ma), involvingSaxothuringian oceanic subduction and subsequentcontinental underthrusting, a 300-km wide orogenicchain evolved. From the NW to the SE, the followingtectonic sequence is developed (Schulmann et al.,2009): the Saxothuringian domain represented byNeoproterozoic basement covered by Palaeozoicsedimentary rocks, the Tepla´ suture zone and thesupra-crustal Tepla´–Barrandian Unit. Further tothe SE, the arc-related granitoid plutons separate theTepla´–Barrandian folded sedimentary rocks from thehigh-grade Moldanubian Zone, which shows widespreadanatexis and contains slices of lower-crustal andmantle rocks. This pervasively deformed root domain isfurther to the east bounded by the Brunia microplate(e.g. Schulmann et al., 2005), which is only marginallyaffected by Variscan tectonometamorphic processes.The Moldanubian Zone consists of middle- andlower-crustal segments, offering an excellent opportunityto examine evidence of the exhumation processesoperating in a collisional setting. The exhumed lowercrust, designated as the Gfo¨ hl Unit (Fuchs, 1976), isrepresented by felsic granulites and anatectic gneissesthat enclose small bodies of mafic granulites, mantlerocks and eclogites. The mid-crustal paragneiss-dominatedlevel has been divided according to the prevailinglithology into the Monotonous Group, with only limitedcontent of intercalations, such as amphibolites orquartzites, and the Varied Group, bearing a largeproportion of intercalated amphibolites, quartzitesand marbles (Fuchs, 1976; Matte et al., 1990). Thestudied BLG is the largest granulite body in SouthernBohemia, and belongs to the Gfo¨ hl Unit, which islocated in a complex stack between the Monotonousand Varied groups, being accompanied by severalneighbouring granulite bodies (Fig. 2).Previous studies of South Bohemian granulitesP–T estimates (Fig. 3) of peak metamorphic conditionshave been calculated by various authors using eitherconventional thermobarometry yielding 1000 °C ⁄1.6 GPa (e.g. Vra´na, 1989; OÕBrien & Seifert, 1992;Carswell & OÕBrien, 1993; Cooke, 2000), thermodynamicmodelling in THERMOCALC software that yieldsa maximum of 850 °C ⁄ 1.6–1.8 GPa (Sˇtípska´ & Powell,2005) or TWEEQU yielding 970–1000 °C ⁄ 1.6–1.7 GPa(Kro¨ ner et al., 2000). The conditions for the amphibolitefacies overprint are estimated to 700–800 °C and0.5–0.8 GPa (Kro¨ ner et al., 2000; Sˇtı´pska´ & Powell,2005; Verner et al., 2007).Ó 2010 Blackwell Publishing Ltd342

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 105(a)(b)Fig. 1. (a) Position of the Bohemian Massif in the Variscan chain in Europe. (b) Simplified geological architecture of the BohemianMassif. Rectangle marks the study area depicted in Fig. 2. Modified after Franke (2000).Radiometric ages ascribed to protolith crystallizationof the South Bohemian felsic granulites give between469 ± 4 and 357 ± 2 Ma using the U–Pb method onzircon (Wendt et al., 1994; Kro¨ ner et al., 2000). Subsequentpeak HP metamorphism took place between351 ± 6 Ma (Wendt et al., 1994) and 341 ± 3Ma(Kro¨ ner et al., 2000). Retrogression under amphibolitefacies conditions proceeded immediately after exhumationto mid-crustal levels, between 340 ± 3 and338 ± 3 Ma (both U–Pb on zircon, Kro¨ ner et al.,2000). Cooling below 500 °C is constrained by the331 ± 1 Ma 40 Ar– 39 Ar age of hornblende from anamphibolite adjacent to the BLG (Kosˇler et al., 1999).Protolith to the felsic granulites is still debated, butthe majority of studies consider the protolith to havebeen a granitic igneous rock (e.g. Fiala et al., 1987;Jakesˇ, 1997; Kotkova´ & Harley, 1999; Finger et al.,2003; Janousˇek et al., 2004, 2006; Tropper et al., 2005;Janousˇek & Holub, 2007). Janousˇek et al. (2004) suggestedan Ordovician granitic protolith with a modelage of c. 450 Ma, which is supported by the radiometricU–Pb zircon ages of Kro¨ ner et al. (2000) andFriedl et al. (2003).Franeˇk et al. (2006) reported a succession of threeductile fabrics in the BLG (Fig. 2). The oldestsubhorizontal fabric S1 is preserved in an 8.5 km ·2.5 km elliptical area of the BLG and macroscopicallyis defined by an alternation of up to 1-cm thickwhite bands formed by recrystallized alkali feldspar,up to 1-cm thick quartz-rich bands and 1- to 3-mmthick darker plagioclase–garnet-dominated bands(Fig. 4a,b). Large perthite porphyroclasts, up to acentimetre in diameter, are preserved in the feldsparrichbands. The S1 foliation is folded by metre-scalepassive F2 folds with steeply inclined axial planesand pervasively developed penetrative cleavage S2,which is macroscopically characterized by strongsubhorizontal elongation of quartz ribbons andbiotite aggregates. Within the limbs, the S1 compositionallayering is stretched and progressively rotatedtowards the N–S striking steep S2 cleavage, whichcontains the subhorizontal L2 stretching and minerallineation. In strongly reworked areas, the only relictsof the S1 fabric are highly attenuated remnants of theS1 compositional layering parallel to S2. Macroscopically,the D2 structures are best seen in the formÓ 2010 Blackwell Publishing Ltd343

106 J. FRANĚK ET AL.Fig. 2. Structural map of the Blansky´ les Granulite (modified after Franeˇk et al., 2006). Stereographic projections depict contoureddensities of foliation poles and dots mark corresponding lineations: S2 – 55 foliations and 24 lineations; S3 – west 51 foliationsand 35 lineations; S3 – east 51 foliations and 54 lineations. Note that in all cases the foliation poles are distributed along great circle(s)whereas the lineations plunge subparallel to p-axes corresponding to these great circles.Ó 2010 Blackwell Publishing Ltd344

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 107Fig. 3. P–T estimates from Moldanubiangranulites accompanied by relevantradiometric ages suggest almost isothermalrapid exhumation from peak metamorphicconditions to mid-crustal levels. Resultsfrom south Bohemian bodies emphasized bythick lines. For references, see fig. 3 inFraneˇk et al. (2011).of the elongated quartz ribbons and biotite aggregatesthat show a strong shape-preferred orientation. Boththe S1 and S2 foliations show a mineral assemblageof Grt–Ky–Bt–Kfs–Pl–Qtz, indicating HP granulitefacies conditions. Mafic boudins enclosed in S2 arecrosscut by Grt–Ky–Kfs–Pl–Qtz felsic dykes. At themarginal part of the elliptical structural relict domain,the S2 foliation is folded by outcrop-scale F3 foldswith the formation of axial gneissosity S3 (Fig. 4c).This gneissosity penetratively transposes all previousfabrics in the majority of the BLG and showssynkinematic retrogressive breakdown of garnet tobiotite and kyanite to sillimanite (Fig. 5f), constrainingthe deformation to mid-crustal conditions. Thedetailed tectonic history and exhumation in a formsimilar to a forced diapir are described in Franeˇket al. (2011) and Lexa et al. (2011).PETROLOGYAnalytical proceduresThe minerals were analysed using the scanning electronmicroscopes Tescan VEGA\\ XMU at 15 kV and0.4 nA at the Strasbourg University and CamScan CS3200 at 15 kV and 3 nA at the LAREM laboratory ofthe Czech Geological Survey in Prague. Compositionalmaps of the feldspar were complemented by spotor area analyses of representative places. Mineralformulae and end-member proportions were calculatedusing the NORM software (Ulmer, 1986). Microstructuraltypes are qualitatively described accordingto the deformational fabrics (S1–S3) and subsequentlythey were quantitatively evaluated using theprogram PolyLX (Lexa et al., 2005) and by electronÓ 2010 Blackwell Publishing Ltd345

108 J. FRANĚK ET AL.(a)perthite porphyroclasts were oriented with respect tothe penetrative S2 foliation and L2 lineation, representingXY, XZ and YZ sections.(b)(c)Petrography of the granulites with the S1 fabricRocks at outcrop H296 are white-grey, fine-grainedgranulites composed of alkali feldspar, quartz,plagioclase and garnet (0.2 mm), with minor biotite,kyanite (0.3 mm) and porphyroclasts (up to 17 mm)of perthitic alkali feldspar. The rocks record evidenceof the complete granulite facies structural evolutiondescribed above. Microscopically, S1 contains discontinuousbands or lenses dominated by plagioclase thatcontain numerous garnet, kyanite, some quartz andbiotite (Fig. 5b,d). The almost monomineralic S1quartz bands are recrystallized into elongated S2ribbons and only quartz accumulation into stripesindicates the original S1 layering (Fig. 5c). Less elongatedquartz grains are rarely preserved in pressureshadows of perthite porphyroclasts. Large perthiteporphyroclasts with numerous lensoidal to lamellaroligoclase exsolutions are recrystallized at their grainboundaries to a mixture of small K-feldspar grains(0.063 mm) with rare perthitic exsolution lamellaeand oligoclase grains (0.047 mm). The feldspardominatedbands with rare garnet are predominantlycomposed of this K-feldspar–plagioclase mixture withminor quartz (0.055 mm), ascribed to the D2recrystallization process (Figs 5a & 6a,b). In therecrystallized matrix composed of K-feldspar, plagioclaseand quartz, minor garnet, kyanite, biotite, rutile,ilmenite, zircon, monazite and apatite also occur(Fig. 6g,h).The perthitic porphyroclasts (up to 17 mm across)contain inclusions of quartz, garnet and kyanite, andmore rarely biotite, rutile, ilmenite, zircon, monazite,apatite and Fe-sulphide (Fig. 6a–f). Quartz inclusions(up to 1 mm across) commonly consist of a singlecrystal with oval or euhedral shape. Garnet (up to1.2 mm across) enclosed in perthite is euhedral and iscommonly surrounded by a thin corona of plagioclase.Subhedral kyanite inclusions are also separated fromperthite by a thick plagioclase corona. Biotite inclusionsin perthite have short prismatic habits, and are inplaces partially retrogressed to chlorite.Fig. 4. Field photographs of (a and b) passive F2 folds depictingpenetrative development of S2 axial cleavage across the foldedS1 compositional layering. (c) F3 folds reworking at amphibolitefacies conditions the granulite facies S2 mylonite.back-scattered diffraction (EBSD). To study the transitionfrom the S1 fabric into the S2 cleavage andthe associated P–T path, the rocks were collected fromone locality within the elliptical relict structuraldomain (outcrop H296; Fig. 2; 48°51¢52.343¢¢N, 14°19¢14.135¢¢E). Sixty thin sections containing 350 largeMineral chemistryTo specify the P–T path for the S1 and S2 fabrics, onesample (H296-S1A) with the S1 layering affected bythe S2 cleavage, as described above, was analysed indetail. It contains garnet, kyanite, perthitic K-feldspar,plagioclase, quartz, biotite, rutile, ilmenite, apatite andzircon. The composition of the minerals in the individualbands is similar. Large perthite grains includequartz, garnet, kyanite, rutile, ilmenite, apatite andzircon. Garnet included in large perthite and garnetfrom the matrix are zoned from core to rim withÓ 2010 Blackwell Publishing Ltd346

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 109(a)(b)(c)(d)(e)(f)Fig. 5. BSE images characterizing the three examined fabrics. (a) S1 compositional layering. (b) S1 rotated subparallel to S2 in afold limb; note quartz shapes in perthite and in the matrix. (c) Quartz-rich S1 band recrystallized to S2 ribbons. (d) Grt–Ky–Pl-rich S1band affected by the D2 deformation. (e) Penetratively developed S2 granulitic mylonite; note plagioclase corona around kyanite.(f) Amphibolite facies S3 fabric; note biotite with sillimanite growing around garnet and a relict kyanite.Ó 2010 Blackwell Publishing Ltd347

110 J. FRANĚK ET AL.(a)(b)(c) (d) (f)(e)(g)(h)Fig. 6. BSE images of S1 layering. (a–f) Inclusions in the large perthite grains reflect peak metamorphic assemblage. A microchemicalprofile across garnet in (f) is presented in Fig. 7. (g) Grt–Ky–Pl-rich S1 band containing biotite, which grows parallel to S2. (h)K-feldspar-rich band enclosing kyanite with typical well-developed plagioclase corona.decreasing grossular, increasing pyrope and almandine,and flat X Fe (Alm 0.52Þ0.61 Grs 0.23Þ0.04 Prp 0.22Þ0.28Sps 0.00–0.01 ; X Fe =Fe⁄ (Fe + Mg) = 0.70, Fig. 7b).The X Fe of biotite in the matrix ranges from 0.33 to 0.35with Ti = 0.21)0.26 (pfu based on 22 oxygen).Recrystallized K-feldspar and plagioclase are zonedfrom core to rim, with Or 77 fi 97 Ab 23 fi 03 An 0.0 andOr 02 fi 01 Ab 81 fi 77 An 18 fi 21 , respectively. Kyanite,Ó 2010 Blackwell Publishing Ltd348

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 111(a)(b)Fig. 7. (a) Ternary plot of feldspar compositions acquired from microchemical analyses of perthite recrystallization and S2 fabric.Similar bulk compositions of parental perthites and shielded matrix are contrasted against areal analyses of deformed matrix. Zoningof plagioclase and to a lesser extent also of K-feldspar are highlighted in the description of data clusters. (b) Microchemical profileacross garnet enclosed in large perthite showing a wide flat grossular-rich core, gradual compositional zoning at both rims and almostconstant X Fe across the whole crystal.both included in perthite and in the matrix, issurrounded by a plagioclase corona.The studied sample H296-S1A lacks biotite inclusionsin the perthite porphyroclasts, and because biotiteand garnet inclusions in perthite are generally rare,the other five samples from the same macroscopicallyhomogeneous outcrop were analysed to test whetherthere is a difference in composition between biotite andgarnet inclusions in perthite and those in the matrix.The biotite X Fe ranges from 0.23 to 0.40 andTi = 0.15)0.34 pfu, and there is no clear differencebetween biotite inclusions and biotite in the matrix(Table 1). Garnet in other thin sections shows the samezoning as that from the sample H296-S1A, but additionallya decrease in X Fe to 0.68 at some garnet rimswas detected.The composition of the large unzoned perthitegrains was acquired in several samples from the sameoutcrop by areal analyses performed by scanning ofregions containing 100 plagioclase exsolution lamellae.It corresponds on average to 68.2% Or, 27.3% Aband 4.5% An (Fig. 7a). The K-feldspar domains ofperthite porphyroclasts contain on average 87.1% Or,12.3% Ab and 0.6% An. The relict perthite porphyroclastsreveal two generations of plagioclase exsolutionlamellae (e.g. Fig. 6c). The coarser exsolution lamellae,presumably older, exhibit elongated braided shapesand are 0.8% Or, 76.7% Ab and 22.5% An in composition.The second micro- to crypto-perthitic generationof exsolution lamellae was not analysed becauseof its small thickness. Detailed feldspar compositionsstudied in relation to the recrystallization mechanismsfor individual deformation episodes are given below.P–T path for S1 and S2 fabricsA pseudosection was calculated using THERMOCALC 3.30(Powell & Holland, 1988) and DATASET 5.5 (Holland &Powell, 1998; November 2003 upgrade), in the systemNCKFMASHTO (Na 2 O–CaO–K 2 O–FeO–MgO–Al 2 O 3 –SiO 2 –H 2 O–TiO 2 –O) with biotite and meltmodels from White et al. (2007), garnet from Dieneret al. (2008), ilmenite from White et al. (2000), feldsparfrom Holland & Powell (2003), white mica fromCoggon & Holland (2002) and cordierite from theTHERMOCALC documentation (Powell & Holland, 2004).Mineral composition isopleths x(g, bi) = Fe ⁄ (Fe +Mg), z(g) = Ca ⁄ (Ca + Fe + Mg) and t(bi) =XTi(M1) were plotted for garnet and biotite. Theanalysed composition of the sample H296-S1A (inwt% SiO 2 = 71.98, TiO 2 = 0.42, Al 2 O 3 = 13.53,FeO = 2.1, MnO = 0.03, MgO = 0.73, CaO = 1.93,Ó 2010 Blackwell Publishing Ltd349

112 J. FRANĚK ET AL.Table 1. Representative microchemical analyses of principal minerals constituting the sample H296-S1 used for calculation of the P–Tpseudosection in Fig. 8.Mineral Grt Grt Bt Kfs Kfs Pl Pl PlSample H296-S1A H296-S1A H296-S1A H296-S1A H296-S1A H296-S1A H296-S1A H296-S1AAnalysis 1 g-c 1 g-r 43 44 45 46 5 47Position core matrix rim matrix matrix core rim core rim very rimSiO 2 38.67 38.79 39.40 65.54 64.95 63.46 63.02 67.40TiO2 na na 4.37 0.00 0.00 0.00 0.00 0.00Al 2 O 3 21.51 21.48 16.88 18.66 18.34 22.66 23.06 20.41FeO 24.93 29.93 12.37 0.00 0.00 0.35 0.00 0.00MnO 0.40 0.44 )0.13 0.00 0.00 0.00 0.00 0.00MgO 5.96 7.99 12.95 0.00 0.00 0.00 0.00 0.00CaO 8.49 1.41 0.00 0.00 0.00 3.76 4.47 2.09Na 2 O na na 0.34 2.48 0.32 9.41 9.17 10.31K 2 O na na 9.73 12.87 15.80 0.16 0.32 0.22Total 99.95 100.05 95.91 99.55 99.42 99.79 100.04 100.43Si 3.00 3.01 3.00 3.01 3.03 2.81 2.79 2.96Ti 0.00 0.00 0.25 0.00 0.00 0.00 0.00 0.00Al 1.97 1.97 1.51 1.01 1.01 1.18 1.20 1.06Fe 3+ 0.04 0.00 0.00 0.00 0.00 0.01 0.00 0.00Fe 2+ 1.58 1.94 0.79 0.00 0.00 0.00 0.00 0.00Mn 0.03 0.03 )0.01 0.00 0.00 0.00 0.00 0.00Mg 0.69 0.93 1.47 0.00 0.00 0.00 0.00 0.00Ca 0.70 0.12 0.00 0.00 0.00 0.18 0.21 0.10Na 0.00 0.00 0.05 0.22 0.03 0.81 0.79 0.88K 0.00 0.00 0.94 0.75 0.94 0.01 0.02 0.01Total 8.00 8.00 8.00 5.00 5.00 5.00 5.01 5.00Prp ⁄ An 0.23 0.31 0.00 0.00 0.18 0.21 0.10Alm ⁄ Ab 0.53 0.64 0.23 0.03 0.81 0.77 0.89Grs ⁄ Or 0.23 0.04 0.77 0.97 0.01 0.02 0.01Sps 0.01 0.01XFe (Fe 2+ ) 0.70 0.68 0.35Na 2 O = 2.76, K 2 O = 4.01, P 2 O 5 = 0.15, H 2 O ) =0.22, H 2 O + = 0.56, CO 2 = 0.03) was modified formodelling by adding 1 mol.% of kyanite to enablea small amount of aluminosilicate to be stable atthe estimated peak metamorphic conditions, as it isobserved in the thin section. The amount of H 2 O=0.33 mol.% was chosen after the construction ofT–M H2O sections, as it allows the garnet X Fe = 0.70 tobe stable in the pseudosection (not shown; see Hasalova´et al., 2008a for approach). The major features of thepseudosection involve melt being stable above 810 °C,biotite stable up to 1.6 GPa and 860 °C, ilmenite stablebelow 1.26 GPa, cordierite stable below 0.7 GPa, anda muscovite-out line heading from 750 °C to 810 °Cat 1.35 GPa and then continuing to 880 °C and2.0 GPa (Fig. 8).The textural features indicate that perthite andminerals included in perthite (kyanite, garnet, biotite,rutile and quartz) belong to the early assemblage,whereas recrystallized plagioclase and K-feldspar,matrix biotite, rutile, ilmenite and quartz belong to alater assemblage. Because of intense D2 recrystallization,the original plagioclase composition is notknown, but the recrystallized monomineralic plagioclasebands and lenses indicate the stability of plagioclasewith alkali feldspar within the S1 structure.Chloritized biotite and probably ilmenite in perthiteporphyroclasts are considered as resulting fromre-equilibration. The flat profile of garnet X Fe isinterpreted as having equilibrated during the laterstage of metamorphism, even for the garnet included inperthite porphyroclasts, whereas strong grossularzoning is consistent with the core composition havingbeen preserved from the early metamorphic stage, andthe rim having equilibrated during the later stage.In the pseudosection, the grossular compositionalisopleth of garnet z(g) = 23 occurs outside the biotitestability field (Fig. 8), whereas biotite also belongs tothe observed peak assemblage. However, the biotitepresentfield is only 0.15 GPa and 30 °C apart from thecompositional isopleth z(g) = 23, and the peak istherefore estimated as the area between these lines,to 1.6)1.8 GPa and 850)880 °C (area 1 in Fig. 8).The X Fe of the garnet (=70) and the grossular contentat the garnet rim (grs = 4) indicate conditions of0.9)1.0 GPa and 825 °C for the second stage (area2 in Fig. 8). An attempt was made to correlate thebiotite composition with the two metamorphic stages,but even if the compositions fit approximately thecalculated values in the pseudosection, it is likelythat the measured range (X Fe = 0.29)0.36, and Ti =0.19)0.26 pfu) reflects partial re-equilibration ondecompression, as does the X Fe of garnet, and thereforecannot be used for more precise estimation ofP)T conditions. The feldspar from the pseudosectioncalculations is not used for thermobarometricÓ 2010 Blackwell Publishing Ltd350

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 113Fig. 8. P–T pseudosection calculated in THERMOCALC software for a felsic granulite sample H296-S1. The indicated P–T path reflectscoarse-grained alkali feldspar growth at peak conditions followed by decompression during S2 fabric development.estimations, because recent models use significantextrapolations from relatively low pressure and temperaturewith respect to the conditions of formation ofthe ternary feldspar of interest (see discussion inSˇtípska´ & Powell, 2005). The lack of a clear differencebetween the chemistry of biotite inclusions and matrixbiotite in the other studied thin sections is interpretedas being due to partial chemical re-equilibration ondecompression, even within perthite porphyroclasts.A decrease in X Fe to 0.68 at some garnet rims in theother studied thin sections (Fig. 7b) may be explainedby further decompression in liquid-absent conditionswhere x(g) decreases, or by heating. Plagioclase coronasaround kyanite reflect decompression (Tajcˇmanova´et al., 2007; Sˇtípska´ et al., 2010).In summary, the peak P–T conditions 1.6)1.8 GPaand 850)880 °C are inferred for the stage when alkalifeldspar porphyroclasts were stable with garnet, kyanite,biotite, rutile, plagioclase and quartz, presumablywithin the S1 layering. The S2 fabric started to developÓ 2010 Blackwell Publishing Ltd351

114 J. FRANĚK ET AL.at these P–T conditions and re-equilibrated to theassemblage corresponding to the garnet rim composition,recrystallized K-feldspar, plagioclase, biotite,rutile, ilmenite and quartz, whereas kyanite separatedby plagioclase from the equilibrated matrix wasmetastable (see also Sˇtípska´ et al., 2010). The P–Testimate for the second stage is 0.9)1.0 GPa and825 °C, showing 0.8 GPa of decompression associatedwith cooling within the S2 fabric.Granular microstructuresProgressive recrystallization of the original coarsegrainedmicrostructure into the fine-grained granularmatrix is recorded in finger-like granular matrixdomains preserved in large perthite crystals shieldedfrom D2 deformation (Fig. 9a,b,e). In this context, theS2 matrix originated by growth and coalescence of thegranular domains preserved in the perthites (Fig. 9a,b)and the perthite grains, as well as the plagioclaseaggregates, represent precursors of the fine-grainedgranular matrix (Fig. 9c,d). A penetratively developedS2 mylonitic foliation is characterized by a granularmatrix composed of fine-grained K-feldspar, plagioclaseand quartz, and elongated coarse-grained ÔplatyÕquartz ribbons (Figs 5e & 9e). The S2–S3 transitionis marked by a substantial change of the granulitemylonitic microstructure towards orthogneiss-likerocks mainly by phase redistribution and grain coarsening(Franeˇk et al., 2006).Type I: microstructure of granular domains within perthitesDevelopment of the granular microstructure startedwith the formation and growth of finger-like, finegraineddomains inside the perthite porphyroclasts(Type I microstructure, Fig. 9a,b). These domainscommonly show a lack of preferred orientation in allthe studied orthogonal sections; only in XZ sectionsare they locally elongated subparallel to the X directionof the D2 strain. The Type I microstructuredeveloped dominantly inside perthite porphyroclastsor along two adjacent perthite crystals and garnet–perthite grain boundaries (Fig. 6f), whereas alongquartz–perthite or kyanite–perthite boundaries itdevelops rarely.The K-feldspar grains in the Type I microstructureshow mainly oval shapes of small axial ratio comparedwith plagioclase. The plagioclase grains range fromcircular at K-feldspar triple junctions to highly elongatedwhere it coats boundaries between either newK-feldspar crystals or between relict perthite and newK-feldspar (Figs 9b,f & 10a). At the recrystallizationfront, the phase boundaries of K-feldspar and plagioclaseare curved and irregular, whereas in the recrystallizedgranular matrix the feldspar boundaries arestraighter (Fig. 10a,b).Areal microchemical analyses of the Type I microstructure(regions containing 100 grains) yield anaverage composition of 67.3% Or, 29.1% Ab and3.6% An, similar to that of neighbouring parentalperthite grains. On average, the K-feldspar compositionis 87.6% Or, 11.8% Ab and 0.6% An, and plagioclasecores consist of 1.5% Or, 78.1% Ab and20.4% An. Compositional maps reveal a sharp reactionfront separating parental perthite from the newlyformed Type I microstructure (Fig. 10a). The boundaryis easily distinguishable by crypto-perthitic exsolutions,which are abundant in the parental perthitebut absent from K-feldspar of the Type I microstructure.New K-feldspar grains reveal weak gradual zoningfrom Or 82 in cores to Or 93 at grain boundaries. Thematrix plagioclase grains show strong zoning markedby An 18–23 cores surrounded by An 11–12 rims. Locally,irregular patches of pure albite occur at the edges ofplagioclase grains.At the edges of the perthite grains, the new matrix isweakly deformed, showing still the Type I microstructure.Nevertheless, increasing elongation of quartzinclusions with weakly recrystallized Type I microstructure(Fig. 9a) documents the influence of D2deformation. This Transitional type II microstructurediffers by the occurrence of quartz grains in the matrix,which suggests exchange of chemical components withthe perthite surroundings. The feldspar composition ofthis transitional type also shows significant differences(Fig. 7a). Area analyses yield an average feldsparcomposition of 46.5% Or, 42.0% Ab and 11.5% An,which is significantly different from both the parentalperthites and the shielded recrystallized domains. TheK-feldspar grains are 90.9% Or, 8.1% Ab and 1.0%An, and the plagioclase is on average 1.6% Or, 76.5%Ab and 21.9% An.The compositional map acquired at the outer edgeof a decomposing perthite also reflects the chemicalchanges. Plagioclase zoning exhibits a distinct pattern,in which An 24 homogeneous grains contain onlyisolated remnants of An 11–13 outer domains. Weakgradual zoning of K-feldspar, in which the rims aredepleted in albite, resembles K-feldspar in the matrixshielded inside perthite.Plagioclase aggregates at this stage are completelyrecrystallized to an equi-dimensional mosaic markedby equigranular grains with straight boundaries,commonly meeting in triple point junctions (Fig. 9d).Minor quartz and K-feldspar locally occur at triplejunctions. Quartz forms weakly elongated aggregatessurrounded by the feldspar matrix.Type II: microstructure of penetrative fabric S2The Type II microstructure is linked to D2 deformationand it is the dominant microstructuraltype throughout the 8.5-km wide relict granulitefacies domain. The Type II microstructure is definedby a feldspar-dominated granular matrix enclosinglarge quartz ribbons. The feldspar aggregate is amixture of plagioclase and K-feldspar, forming anÓ 2010 Blackwell Publishing Ltd352

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 115(a)(b)(c)(d)(e)(f)Fig. 9. BSE images documenting recrystallization of large perthite and formation of the fine-grained granular matrix. (a and b)Heterogeneous recrystallization of the parental perthite at a sharp reaction front in domains partly parallel to trace of the S2 fabric.Note the elongation of a quartz inclusion released from the perthite documenting D2 strain. (c and d) Recrystallized Pl–Grt-richband with interstitial quartz and cuspate geometry of several K-feldspar grains. (e) Recrystallization of a perthite grain partly exposedto the D2 strain (top of the porphyroclast) and partly in a pressure shadow (right side). (f) Detail of a shape-preferred orientationof plagioclase in a newly formed granular matrix adjacent to a parental perthite.Ó 2010 Blackwell Publishing Ltd353

116 J. FRANĚK ET AL.(a)(b)Fig. 10. Compositional maps of Ca and K from (a) matrix shielded inside parental perthite and (b) progressively deformed matrix inthe S2 fabric. For Ca, the K-feldspar and quartz grains are masked by white and black colour to emphasize the Ca zoning inplagioclase. Similarly, the plagioclase and quartz grains are masked in the K distribution maps to demonstrate the K zoning inK-feldspar. The BSE images are offered for better orientation in phase distribution.Ó 2010 Blackwell Publishing Ltd354

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 117equi-dimensional mosaic (Fig. 5e). Transposition of S1compositional layering into the S2 fabric via rotationand attenuation suggests intense deformation duringD2 (Fig. 9a). The quartz bands inherited from S1disintegrated to form the ribbons parallel to S2 andwere further dynamically recrystallized along ribbonboundaries, leading to significant reduction of quartzgrain size (Fig. 11a). Numerous new small quartzgrains nucleated in triple junctions of feldspar grains inthe matrix. The garnet grains became dispersed in thematrix and kyanite crystals deformed by kinking.Isolated biotite flakes lie parallel to S2 and, togetherwith quartz ribbons, define a strong L > S fabric.Areal analysis of a Type II matrix region covering100 feldspar grains yields an average feldspar compositionof 56.4% Or, 35.9% Ab and 7.7% An(Fig. 7a). The average composition of the K-feldspar is84.7% Or, 13.5% Ab and 1.8% An, and the plagioclaseis composed of 1.8% Or, 74.6% Ab and 23.6%An.Two compositional maps acquired from the feldspar-dominatedmatrix (one depicted in Fig. 10b)reveal typical mild zoning of K-feldspar, reflectingthe loss of Na around grain boundaries. Plagioclaseexhibits uniform An 23 composition with a weakincrease of 2–3% anorthite content towards rims.Locally, thin films of albite occur along plagioclase–K-feldspar and K-feldspar–K-feldspar boundaries;they are significantly thinner than rims in the Type Imicrostructure.Type III: microstructure of S3 amphibolite facies fabricThe S3 fabrics prevailing in the felsic granulites of theBLG reveal highly variable degrees of retrogressionof the previous mineral assemblage. Contemporaneoushydration of granulites heterogeneously increasestowards the boundaries of the granulite massif, whereit is accompanied by widespread partial melting andsegregation of the melt into mm–dm thick bandsparallel to S3 (Kodym, 1972; Franeˇk et al., 2006). Thequartz, K-feldspar and plagioclase show irregulargrain boundaries and develop into monomineralicaggregates (Fig. 12), whereas the quartz ribbons typicalfor the S2 entirely disappear. The aspect ratios ofquartz grains significantly decrease, compared with theS2 fabric, in conjunction with coarsening of feldsparmosaic (Figs 5f & 11b). The abundant biotite flakeslie parallel to the foliation planes and sillimanite insome instances defines the lineation. The resulting S3microstructure resembles a common Bt + Sil ± Grtorthogneiss more than a retrograde granulite.In order to eliminate the effect of strain localizationinto melt bands or biotite stripes that developed duringD3, only macroscopically homogeneous S3 sampleswith dispersed biotite were studied. The petrology andmicrochemistry of a steep fabric from a neighbouringKrˇisˇťanov Granulite Massif, analogous to the S3 in theBLG, are given in Verner et al. (2007).Quantitative microstructural analysisThe previous section defined three types of microstructuresdeveloped from the coarse-grained precursorof the felsic granulites: Type I corresponds togranular microstructure inside coarse perthite and atransitional type in the vicinity of coarse perthitegrains; Type II microstructure corresponds to thepervasively developed granulitic S2 fabric; and TypeIII microstructure corresponds to the S3 fabric developedunder amphibolite facies conditions. Accordingto the principle of fabric superposition (Wilson, 1961),the succession of microstructural types is seen as anevolutionary microstructural trend reflecting deformationstages along the exhumation path of thegranulites. The coarse-grained orthogneiss precursormicrostructure cannot be sufficiently describedusing quantitative methods. In order to comparethe three microstructural stages, we have quantifiedseveral parameters characterizing the microstructuresusing the PolyLX toolbox (Lexa et al., 2005) forthe MATLAB TMsoftware package and CSD-correctionprogram (e.g. Higgins, 1998). Four thin sections orientedparallel to the lineation and perpendicular to thefoliation (XZ sections) from representative samples ofthe S2 (956 & 1252 grains) and S3 fabrics (848 & 886grains) were digitalized in an ArcView GIS environment(Fig. 11a,b) and analysed by the PolyLX andCSD-correction software. In order to obtain statisticallysignificant results, we have merged four areas ofthe Type I microstructure yielding 801 grains. Beforemerging, the individual areas were rotated with respectto the trace of S2. In addition, the transitional matrixarea at the boundary between the perthite and D2matrix was digitalized, covering 868 grains.Different microstructural types have been quantifiedin plots of bulk grain-boundary preferred orientation(GBPO) against contact frequencies (Fig. 12) andslope of linear regression from crystal size distribution(CSD) analysis v. regression intercept (Fig. 13). Thegrain size as well as other quantitative characteristicsare statistically evaluated in Table 2, more detaileddescription of the quantitative data is given inAppendix S1.The contact frequency method (Kretz, 1969)compares the observed (O) count of contacts betweentwo minerals with the value expected (E) for a perfectlyrandom distribution. The calculation involves modalproportions of phases, only a scatter in grain size maycause a limited uncertainty of the results. The boundariesare designated as Ôlike–likeÕ for contacts betweengrains of one phase or ÔunlikeÕ for the case of boundarybetween two different phases. The resulting v 2 -value(O ) E) ⁄ sqrt(E) is a measure of the deviation ofmineral spatial distribution from random. It is positivefor like–like contacts and negative for unlike contactsif the corresponding microstructure tends to formmonomineralic aggregates (Lexa et al., 2005). On theother hand, the unlike contacts exhibit positive values,Ó 2010 Blackwell Publishing Ltd355

118 J. FRANĚK ET AL.(a)(b)Fig. 11. Microstructural character of XZ sections (X is parallel to lineation, Z normal to foliation) of the two fabrics quantified inthe felsic granulites: (a) the granulite facies S2 and (b) the amphibolite facies S3. The left column combines optical image undercrossed polarizers with its vectorized form used for statistical analysis in Fig. 12. The right column presents BSE images pointing outphase distribution.Ó 2010 Blackwell Publishing Ltd356

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 119Fig. 12. Quantitative microstructuralanalysis of the degree of grain boundarypreferredorientation v. contact frequenciesfor like and unlike boundaries of feldspars.Numbers at individual data points refer tothe evolution stages defined in the sectionGranular microstructures.Fig. 13. Crystal size distribution curves for both feldspars. Values of slope and upper intercept of linear regression are plotted togetherin the inset showing an evolutionary trend for both the feldspars. Numbers at individual data points refer to the evolution stagesdefined in the section Granular microstructures.whereas the like–like contacts are negative where thephases tend to be regularly distributed.Crystal size distributions measured in metamorphicrocks yield quantitative information about crystalnucleation and growth rates, growth times and thedegree of overstepping of reactions during metamorphism(Cashman & Ferry, 1988). The CSDs aredescribed as the population density function of thecumulative number of crystals per unit volume perlinear crystal size. The data are plotted in log-normalspace and a linear regression of points is the characteristic.The intercept values (N0) are proportional tothe nucleation density of the examined phase at theonset of nucleation (size 0), and the slope (Gt) is proportionalto the modification of original grain-sizedistribution by a variety of coarsening processes suchÓ 2010 Blackwell Publishing Ltd357

120 J. FRANĚK ET AL.Table 2. Quantitative characteristics of the microstructuralevolution.EAD (lm) Bulk SPO AxialratioQ1 Median Q3as Ostwald ripening or Ôcommunicating neighboursÕtheory (Marsh, 1988). Although the individual absolutevalues are meaningless without knowledge ofkinetic parameters, they can be used to compare individualstages of the textural maturation history in asingle rock type (Lexa et al., 2005; Hasalova´ et al.,2008b; Schulmann et al., 2008). The SPO and GBPOare calculated as ratios of eigenvalues of the orientationtensors of long axes of minerals and grainboundaries, respectively (Lexa et al., 2005), and theirvalues usually depend on types of active recrystallizationand ⁄ or deformation mechanisms. The aspectratio, SPO and GBPO are relatively high for dynamicallyrecrystallized grains deformed by dislocationcreep (Kruse et al., 2001; Ulrich et al., 2002; Baratouxet al., 2005), whereas low values commonly characterizeGBS-controlled diffusion creep associated withmutual rotation of individual grains (Boullier &Gueguen, 1975; Behrmann & Mainprice, 1987).Grain contact frequency evolutionCSDinterceptCSDslope1-shielded granular Kfs 33.7 56.3 84.4 1.21 1.53 11.57 )16.2matrixPl 26.8 42.3 59.4 1.22 1.51 12.78 )22.62T-weakly deformed Qtz 18.0 26.6 40.9 2.90 1.67matrixKfs 44.8 67.1 90.8 1.40 1.49 11.57 )15.7Pl 29.2 40.3 59.0 1.20 1.53 13.54 )26.02-S2 granulitic matrix Qtz 14.7 22.5 45.2 2.12 1.49Kfs 56.5 83.9 126.7 1.30 1.43 10.40 )11.6Pl 46.3 64.7 90.6 1.14 1.48 11.73 )16.73-S3 retrograde fabric Qtz 39.5 59.8 104.2 1.50 1.46Kfs 42.7 67.4 109.2 1.40 1.54 10.07 )11.4Pl 40.9 67.2 111.5 1.27 1.43 10.50 )12.5Equal area diameter (EAD) statistical characteristics depicting K-feldspar, plagioclase andquartz grain-size evolution from the perthite recrystallization process to amphibolite faciessyntectonic retrogression. Q1 and Q3 refer to 1st and 3rd quartiles. The bulk SPO iscalculated as the ratio of eigenvalues of bulk orientation tensor according to Lexa et al.(2005), the axial ratio values represent modus of their lognormal distribution.The grain contact frequency method yields a welldefinedevolutionary trend of like–like contacts forboth Kfs and Pl. The two minerals reveal fairly similarand highly negative values of (O ) E) ⁄ sqrt(E) ratio()6) for Type I microstructure (Fig. 12), suggestinghighly regular spatial distribution of both feldspars.Both feldspars also reveal distinct GBPO (1.07–1.15).The Type II microstructure shows an important increaseof like–like grain contact frequency in Fig. 12,reaching zero, which suggests an almost perfectlyrandom distribution of both feldspars, coupled withalmost complete loss of GBPO. Finally, the Type IIImicrostructure reveals a further increase of(O ) E) ⁄ sqrt(E) values, suggesting a progressivelydeveloping aggregate distribution (higher for plagioclasecompared to K-feldspar) associated with anincrease of GBPO values. The evolution of K-feldspar–plagioclase grain boundary frequencies mirrors thetrend of like–like contacts, but the GBPO remainsmore developed for all stages, compared with like–likeboundaries, with peaks in Types I and III microstructures.The progressive evolution of rock structurefrom a regular distribution of grains through randomto an aggregate distribution indicates a process ofsolid-state differentiation and development of minerallayering. The fluctuation of GBPO suggests destructionof the original preferred orientation of grainboundaries by D2, whereas D3 was responsible for thedevelopment of a new GBPO.Grain-size evolutionThe grain-size statistics (Table 2, Appendix S1) has arather constant and low value (55–70 lm) for recrystallizedK-feldspar of the Type I microstructure andtransitional microstructures adjacent to perthite crystals,and a slightly larger grain size for the Types II andIII microstructures (84 & 67 lm, respectively). Similarly,fine-grained plagioclase has a small and constantgrain size for the Type I microstructure (42 lm) and alarger grain size for Types II and III (66 lm). Thequartz grain size first decreases from the transitionalmicrostructure (26 lm) to the Type II microstructure(23 lm), and increases in the Type III microstructure(60 lm). In general, the K-feldspar recrystallized grainsize is the largest of all the studied minerals (70 lm).The CSD analyses reveal a two stage grain-sizeevolution for both plagioclase and K-feldspar(Fig. 13). Both minerals show decrease of N0 valuesand increase of Gt values from the Type I to Types IIand III microstructures. For plagioclase the evolutionis marked by gradual decrease in the N0 value, coupledwith an increase in the Gt value from the Types II toIII microstructures. Importantly, although the evolutionof the plagioclase and K-feldspar grain size startsfrom very different positions in the Gt–N0 plot (higherN0 and smaller Gt in plagioclase, compared withK-feldspar), they end up with nearly the same values ofboth parameters for the Type III microstructure. Bothminerals reach the greatest difference in grain-sizeparameters for the Type II microstructure, for whichplagioclase shows significantly higher N0 values andsmaller Gt values compared with K-feldspar.Shape-preferred orientation and grain elongationThe aspect ratio remains constant for both K-feldsparand plagioclase during the whole deformationsequence and the bulk shape-preferred orientation(SPO) of plagioclase remains similarly low, whereasK-feldspar SPO is slightly higher for Types II andIII microstructures (Table 2). Quartz reaches thehighest degree of SPO (mean value 2–3) for Type II,together with a relatively high aspect ratio. The TypeIII microstructure is characterized by a decrease of thequartz SPO in conjunction with a significant decreasein the aspect ratio (Table 2).Ó 2010 Blackwell Publishing Ltd358

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 121Lattice-preferred orientationLattice-preferred orientation (LPO) measurements ofquartz, K-feldspar and plagioclase grains from all thedescribed microstructural types were additionallycarried out in order to evaluate operative deformationmechanisms. A strong LPO usually originates duringplastic deformation via dislocation creep, whereas thediffusion-accommodated GBS generally weakensthe LPO of deforming grains (e.g. Jiang et al., 2000).The slip systems accommodating dislocation creep arewell known for quartz (e.g. Schmid & Casey, 1986) andplagioclase (e.g. Tullis, 1983; Montardi & Mainprice,1987; Kruse et al., 2001; Stu¨ nitz et al., 2003), but lessknown for K-feldspar (Tullis, 1983). To ensure highquality of measurements, the crystal lattice orientationswere collected manually via EBSD in theLAREM laboratory of the Czech Geological Surveyand at the Institute of Petrology and StructuralGeology at Charles University in Prague. LPO data foreach sample (XZ thin section) were plotted separatelyas non-polar projections on a lower hemisphere andimportant slip planes and directions for quartz,K-feldspar and plagioclase were projected.Type I microstructureAt first, the EBSD study was carried out in domainsconsisting of perthite and the newly formed Type Imatrix enclosed within the perthite, therefore protectedfrom D2 and D3 deformation. The Type I K-feldsparand plagioclase grains reveal roughly the same latticeorientation as the K-feldspar host and plagioclaseexsolutions in parental perthite. When passing outsidethe protected matrix into the Transitional type IImicrostructure, a continuous increase of LPO scatteringwith respect to the parental perthite orientationoccurs in the matrix without any tendency to develop anew LPO (Fig. 14).Type II microstructureThe pole figures of matrix quartz show incomplete highangleType II cross-girdle (Lister & Price, 1978) of c-axes(Fig. 15a). High opening angle and central maximumare consistent with dominance of prism and slip systems, being variably accompanied by rhomb + slip (e.g. Lister & Dornsiepen, 1982).The large quartz ribbons have an irregular size anddistribution of subgrains, with inclined subgrainboundaries and a strong central maximum sometimesaccompanied with minor submaxima oriented closeto the lineation. This typical granulite pattern (Behr,1961) results from predominance of the prism slip, only rarely accompanied by subordinate prism slip (Schmid & Casey, 1986).In plagioclase, the LPO shows an incomplete girdleof (010) poles normal to the lineation direction andmaxima of [201], [101] and [001] subparallel to thelineation. It indicates the activity of primary [001] (010)Fig. 14. Electron back-scattered diffraction measurements of a perthite domain undergoing recrystallization to granular matrix andinitial D2 deformation. The crystal lattice orientation of both K-feldspar background and plagioclase exsolutions in parental perthiteare compared to both feldspars in the granulitic matrix located inside the perthite, at the edge of the perthite and in the deformedmatrix adjacent to the perthite grain. The elongation of quartz grains reflects D2 strain intensity and depicts orientation of S2 fabric.Lower hemisphere non-polar equal-area projections.Ó 2010 Blackwell Publishing Ltd359

122 J. FRANĚK ET AL.(a)(b)Fig. 15. Polar diagrams of lattice preferred orientation of quartz, K-feldspar and plagioclase from XZ sections of felsic granulitesexhibiting penetrative (a) S2 and (b) S3 fabrics. Only the slip planes significant for each mineral are presented; lower hemispherenon-polar equal-area projections.as well as secondary [101] (010) and [201] (010) slipsystems according to Kruse et al. (2001). The LPO ofK-feldspar is similar to that of plagioclase suggestingactivity of the same slip systems. Additional slip system[100] (010) in K-feldspar makes the only differencebetween both of the feldspars. It is noteworthy thatpole figures of both feldspars show double maxima ofsingle crystal orientations close to each other.Ó 2010 Blackwell Publishing Ltd360

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 123Type III microstructureThe quartz LPO shows a high angle Type II crossgirdleof c-axes distribution with a maximum either inthe centre or near edges of the girdle. The LPO issimilar to those measured from the S2 matrix microstructureand implies a combination of prism and rhomb + slip (Fig. 15b). Plagioclaseshows the (010) and (001) planes subparallel to foliationand a girdle distribution of [100] and [201] alongthe foliation with single maxima parallel to the lineation.This pattern points to activity of three slip systems,namely [100] (010), [201] (010) and [100] (001). Inthe case of K-feldspar, all samples show a pronouncedmaximum of (010) planes parallel to foliation, whereasother common slip planes do not show such relationship.A girdle distribution is observed for the [001] and[100] directions, with a maximum of the [001] directionoriented parallel to the stretching lineation. Theseattributes indicate that the activity of [100](010) slipsystem is the characteristic for both feldspars.DISCUSSIONWe discuss the origin and petrological significance ofthe S1 fabric as a precursor for development of characteristicgranular microstructure of Variscan felsicgranulites. Subsequently, an attempt is made to discussthe origin of granular fabric, its specific quantitativemicrostructural characteristics and its importance forunderstanding deep crustal flow processes. Finally, thepetrology, microstructure and LPO of various types ofgranulites are discussed in terms of the rheologicalevolution of orogenic lower crust and its extrusion(Franěk et al., 2011) during the Variscan orogeny inEurope. Figure 16 puts the microstructural evolutionin context with the tectonic history.Interpretation of S1 fabric: HP orthogneissThe observed structural succession, field distributionof the S1 relicts and microstructural relations implythat the perthitic alkali feldspar, plagioclase and quartzaggregates preserved locally within the S1 compositionallayering represent a remnant of the precursor ofthe Blansky´ les felsic granulites. The best preservedrelicts of the S1 layering contain a substantial amountof large alkali feldspar (presently perthite) and largequartz grains in the microstructure. The inclusions ofkyanite, high-grossular garnet and rutile suggest thatthe large alkali feldspar crystallized at HP conditions(Figs 8 & 16). Despite diffusional re-equilibration ofmost minerals, the peak P–T conditions related to thegrowth of large alkali feldspar are estimated at1.6)1.8 GPa and 850)880 °C. The abundant idiomorphicor oval-shaped quartz inclusions in the largeperthite porphyroclasts point to a lack of plasticdeformation during alkali feldspar growth. Suchinclusions may develop either in granites, or duringmelt-assisted recrystallization in migmatites (Mehnert,1968, pp. 111, 192). The lenticular plagioclase-richaggregates are interpreted as completely recrystallizedremnants of large plagioclase grains complementary tothe alkali feldspar porphyroclasts (Figs 9c,d & 16, 3Dblock diagram 1). Consequently, the original rock was acoarse-grained granite ⁄ orthogneiss crystallized underHP conditions in the kyanite stability field. Existence oftwo complementary feldspars at peak conditions contrastswith laboratory experiments of Tropper et al.(2005) focused on the South-Bohemian felsic granulites,where only a single alkali feldspar existed above 850 °C.The S1 layering of initially coarse-grained quartz,K-feldspar and plagioclase-rich layers may be interpretedas a result of solid-state deformational segregationof minerals with different plasticity at hightemperatures, a process common in formation of, e.g.layered orthogneisses. The length of the bands thenindicates significant strain during the D1 phase.Assuming growth of large alkali feldspar during theD1 episode, the overgrowth of quartz, garnet andbiotite by the large feldspar indicates a dominance ofgrowth processes resulting in overall coarsening of thegneiss (Higgins, 1998; Lexa et al., 2005). Such crystalgrowth indicates that the temperature ⁄ strain rate ratiowas rather high and thermodynamic modelling (Fig. 8)suggests contemporaneous partial melting. High-Ppeak conditions (1.6–1.8 GPa) indicate that the flowoccurred at the bottom of thickened crust as alsosuggested by other studies (e.g. Sˇtípska´ & Powell, 2005;Racek et al., 2006; Tajcˇmanova´ et al., 2006). Therefore,the S1 fabric probably originated during lowercrustal flow at the bottom of a thickened crustal rootand at progressively increasing temperatures (Fig. 16).The inclusions within the perthite porphyroclasts,the S1 compositional layering and the S2 myloniticfoliation in the BLG are similar to the oldest structuresdescribed from the Saxonian Granulitgebirge (Behr,1961). In addition, all of the Variscan felsic granulitesreveal similar geochemistry and geochronology(340 Ma peak). Consequently, both the mentionedmassifs and other Variscan felsic granulite massifs aswell may have evolved from such HP coarse-grainedorthogneisses.Origin of granular matrix in granulitesThe microchemical and textural relations indicate thatthe initial large alkali feldspar underwent heterogeneousstep-by-step recrystallization during the D2event, which resulted in formation of a fine-grainedmatrix. This recrystallization probably postdatedexsolution of the coarse perthitic plagioclase because(i) the newly formed Type I K-feldspar in the granuliticmatrix is devoid of such large braided exsolutions, and(ii) the perthite exsolutions reveal 1–2% higheranorthite content than the cores of Type I matrixplagioclase indicating a bit higher temperature of formation(Fig. 16, 3D block diagram 2). Nevertheless, itÓ 2010 Blackwell Publishing Ltd361

124 J. FRANĚK ET AL.Fig. 16. Interpretative 3D scheme of the microstructural evolution of felsic granulites from the coarse-grained layered HP granuliteprecursor to granulite facies mylonite developed during exhumation in a crustal scale ÔdiapiricÕ dome according to Franeˇk et al. (2011).The block diagram 1 shows the tentative model of original layered orthogneiss microstructure with numerous kyanite, quartz andgarnet crystals in large feldspar porphyroblasts. The model of ÔdiapiricÕ exhumation (Franeˇk et al., 2011; Lexa et al., 2011) shows theposition of this microstructure at the bottom of the crust and its P–T conditions (1). The extent of preservation of the S1 fabric ismarked by light shading at the bottom of diapiric extrusion. The block diagram (2) shows the development of perthites at the beginningof exhumation with decompression and limited cooling. The block diagram (3) exhibits the development of typical granulite fabric,such as platy quartz and granular microstructures of Types I and II. Relicts of kyanite preserved in perthite porphyroclasts and matrixwith decompression kelyphitic haloes are shown. This microstructure is developed in the main vertical channel of the ÔdiapirÕ and ismarked by intermediate shading in the structural model. The ÔdiapirÕ head shows the stability of Type III granular microstructure (darkshading) and rotated relict domains preserving Type II microstructure.Ó 2010 Blackwell Publishing Ltd362

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 125cannot be fully excluded that the large perthitic exsolutionsin alkali feldspar porphyroclasts originatedduring later stages of cooling and decompression alongthe exhumation path.The recrystallization must have initiated by deformation-inducedconversion of the (perthitic) parentalalkali feldspar directly into small K-feldspar grains byan uncommon process, which is discussed in the nextparagraphs. Plagioclase was then gradually redistributedalong the new K-feldspar grain boundaries (e.g.Fig. 9b,f), presumably via grain-boundary diffusion.The sharp compositional boundaries between An 23cores and An 12 rims of the Type I plagioclase(Fig. 10a) indicate that they did not develop by continuousgrowth and that they were not significantlyaffected by later diffusion. We suggest that at first, theAn 23 plagioclase grains coalesced in triple junctions ofthe new K-feldspar grains by grain-boundary diffusionmechanism. They originated by reduction of surfaceand coalescence of preexisting coarse and compositionallysimilar An 24 plagioclase perthitic exsolutions,assuming perthitic nature of the parental alkali feldspar.The new grain boundaries were subsequentlyovergrown by An 12 plagioclase, the components forwhich were probably released from surroundingK-feldspar by volume diffusion. This process is indicatedby gradual zoning at rims of adjacent K-feldspargrains (Fig. 10a). The CSD patterns (Fig. 13) showthat the formation of the Type I K-feldspar–plagioclasemosaic originated by a nucleation-dominatedprocess accompanied by limited crystal growth, comparedwith the later Types II and III microstructure(Lexa et al., 2005; Hasalova´ et al., 2008b). The nucleationdensity of plagioclase was significantly higherthan that of K-feldspar.The most important manifestation of the perthiterecrystallization is redistribution of Or, Ab and Ancomponents on a micro-scale in a chemically closedsystem, resulting in development of the fine-grainedgranular matrix (Fig. 16, 3D block diagram 3). Thehigh amount of unlike grain boundaries in the granularmatrix preserved inside parental perthite indicates atendency to regular distribution and thus active intermixingof the K-feldspar and plagioclase. The regulardistribution represents the energetically lowest state ofsuch a polyphase material (Seng, 1936; DeVore, 1959)and progressive straightening of grain boundarygeometries of K-feldspar and plagioclase further minimizesthe surface energy by grain shape simplificationtypical for grain boundary reduction (Passchier et al.,1992). According to this model the recrystallizationthen appears to be substantially driven by a decrease ofchemical and surface energy in the metastable (perthitic)alkali feldspar. The concept of surface energyminimization was suggested by Flin (1969), who proposedthat the regular distribution is a consequence ofa smaller interfacial energy of unlike boundaries incomparison with like–like boundaries. However,Ramberg (1952) considered the interfacial energies tobe too small to drive diffusional mass transfer ingranulites.The weak but systematic scatter of lattice orientationsof new plagioclase and K-feldspar grains withrespect to the parental perthite (Fig. 14) points tocontribution of stress to the bulk recrystallizationprocess. Such a process of non-coherent decompositionof metastable alkali feldspar may be triggeredeither by water addition or by the application ofexternal stress (Brown & Parsons, 1989 and referencestherein). As a lack of hydrous phases and presumablylack of water is typical for the granulite facies evolutionof these rocks, the activity of external stressseems to be better candidate to trigger the recrystallizationprocess.The almost identical LPO of plagioclase andK-feldspar, regular grain distribution and high nucleationdensity CSD pattern of the polymineralicaggregate developed by recrystallization of solidsolution crystal has not been previously described innature. The sharp recrystallization front advancinginto parental perthites represents a moving grainboundary driven by both chemical and deformationalprocesses. According to Stu¨ nitz (1998), such recrystallizationcan be considered as a combination ofstrain- and chemically induced grain boundarymigration processes. At the temperature and chemicalconditions given during the D2, the activation energyof the grain boundary migration process was probablylower than that of dynamic recrystallization of themetastable (perthitic) alkali feldspar. Material scienceliterature (e.g. Saheb et al., 1995; Sennour et al., 2004)has examples of sharp recrystallization fronts advancinggradually into intact grains of alloys (Fig. 9a,b,f)explained by the chemically induced grain boundarymigration or a similar discontinuous precipitationprocess, which in many aspects resembles the alkalifeldspar recrystallization. According to Yoon (1995),the discontinuous precipitation in metastable alloysrepresents largely autocatalytic recrystallization drivenmainly by coherency strain energy. Activation of thisprocess depends upon composition, temperature, stressand strain, elastic anisotropy, crystallographic relationsand boundary curvature of involved grains(Yoon, 1995). Once triggered, e.g. by external strain, itcauses very efficient heterogeneous decomposition ofoversaturated solid solutions (like alkali feldspar) via(mainly chemically driven) grain boundary migration(Hay & Evans, 1987; Saheb et al., 1995; Sennour et al.,2004).In conclusion, the formation of the granular matrixresembles the discontinuous precipitation triggered bystraining of the probably perthitic alkali feldspargrains during D2 deformation (Fig. 16). The processstarts preferably on pre-existing grain boundaries withother feldspar porphyroclasts or with garnet. Themigration of the recrystallization front is driven mainlyby the metastability of the (perthitic) alkali feldspar. Inturn, the metastability of parental feldspar controls theÓ 2010 Blackwell Publishing Ltd363

126 J. FRANĚK ET AL.development of the irregular to amoeboid domains(e.g. Fig. 9a,b) of fine-grained regularly distributedK-feldspar and plagioclase matrix.Lower crustal ascent and strength evolution of felsicgranulitesThe onset of the D2 deformation is marked by highlynon-cylindrical folding of the S1 anisotropy prior tothe S2 cleavage development (Figs 4a,b & 16, blockdiagram 3). The formation of a typical Variscan felsicgranulite fine-grained microstructure with highlyelongated quartz ribbons is the main result of the D2deformation (Figs 5e & 9e) that is contemporaneouswith the development of the granular Type I microstructure.Petrological modelling and mineral zoning,particularly the decrease of the grossular componentin garnet, suggest that between the perthite growthand the end of the D2 there was significant decompressionof 0.8 GPa at slightly decreasing temperatureover a time span of several million years (Figs 8& 16). That means the S2 microstructure developedat slightly lower temperature compared with the S1and under continuously decreasing pressure as indicatedby a number of decompression microstructuressuch as plagioclase rims around kyanite and garnet(Figs 5e & 6h). This conclusion agrees with theobservations of spinel–plagioclase and garnet coronasdeveloped around kyanite, which form characteristicdecompression microstructures in S2 (Sˇtípska´ et al.,2010). The common denominator of these coronas isthat they are elongated parallel to S2 foliation(Fig. 6h) and therefore developed at the end of theD2 deformation.The small interstitial quartz grains, discontinuousalbite films at plagioclase or K-feldspar boundariesand cuspate feldspar grain shapes (Fig. 9d) are consistentwith the presence of syn-deformational intergranularpartial melt in the granulitic S2 matrix, aconclusion previously suggested by Franeˇk et al.(2006) and supported also by dykes of felsic granulitecrosscutting mafic boudins embedded in S2. Tajčmanova´et al. (2006, 2007) inferred a low content ofpartial melt (

PRECURSOR AND RHEOLOGY OF VARISCAN GRANULITES 127D3 switch in rheology of feldspar v. quartzThe microstructural characteristics of the S3 microfabric(Type III microstructure) differ significantlycompared with those of the S2. The high number oflike–like boundaries in the feldspar-dominated matrix(Fig. 12) indicates significant coalescence of individualphases into monomineralic aggregates, and the SPO,GBPO and axial ratio of both feldspars also increase.In contrast, quartz grains exhibit a decrease in SPO aswell as in axial ratio. The Type III microstructure isclearly characterized by a growth-dominated process,especially in the case of K-feldspar. The N0–Gt valuesof Types I, II and III microstructures project on acurve indicating simultaneous change of N0 and Gtdue to temperature and strain variations and suggestthat strain rate change probably plays a key role in theresulting CSD shape. Such an evolutionary trend isoften interpreted as a result of grain coarsening, whichunusually occured at lower temperature related toupper amphibolite facies metamorphism during D3(700–800 °C) compared with the D2 granulite faciesconditions (800–850 °C) (Figs 3 & 8). Therefore, thestarting D2 recrystallized grain size was not in equilibriumwith the temperature-corrected strain rate(Zener & Holomon, 1944, parameter Z), such that themicrostructure was located in the grain-coarseningfield (fig. 1 of Ulrich et al., 2006 modified after Sakai &Jonas, 1984). In conclusion, only a significant strainrate decrease can explain the observed S3 graincoarsening at decreased temperature.The S3 microstructure tendency for monomineralicaggregate distribution, high SPO and the generalprevalence of grain growth over nucleation comparedwith S2 are in agreement with the strong LPO, suggestingthat the quartz, K-feldspar and plagioclasedeformed predominantly by dislocation creep. Theirregular grain boundaries in the S3 microstructure arealso typical of dislocation creep rather than GBS ordiffusional mechanisms. The dominance of dislocationcreep indicates hardening of the felsic granulites, wherethe less plastic feldspar forms a low viscosity contrastload-bearing framework (LBF, Handy, 1990) that resultedin high bulk strength of the retrograde felsicrock in mid-crustal levels.The precursor of the Blansky´ les felsic granulites wasan HP alkali feldspar–plagioclase–quartz–garnet–biotite–kyanite bearing coarse-grained layered orthogneissthat developed via deformational solid-statesegregation of quartz and feldspars. It records peakP–T conditions of 1.6)1.8 GPa and 850)880 °C(Fig. 16).Granular matrix (Type I microstructure) formed byrecrystallization of 10 mm large and probablyperthitic strong alkali feldspar porphyroclasts intoa weak fine-grained (0.055 mm) plagioclase–K-feldsparmatrix. This transition occurred via a processresembling discontinuous precipitation mechanism inmetal alloys. It was driven mainly by chemically andstrain induced grain boundary migration, but triggeredby stress build up during early stages of D2. Newlyformed matrix grains maintained roughly the latticeorientation of the parental alkali feldspar crystal. Theyhave a sufficiently small grain size to allow flow byGBS.The fine-grained S2 granulitic matrix (Type IImicrostructure) with a very low content of silicatemelt deformed predominantly via diffusion creepcontrolledGBS. The large quartz ribbons were rheologicallystronger than the feldspar-dominatedmatrix due to the activity of different deformationalmechanisms. During this stage, in only several millionyears, the linearly viscous granulite matrix containinglow amount of interstitial melt was extruded upwardsthrough a vertical channel (Franeˇk et al., 2011) frompressures of 1.7 GPa at 850 °C to 1.0 GPa at800 °C (Fig. 16).After cooling and crystallization of the syn-D2partial melt along grain boundaries, the granulitesdeformed at lower P–T conditions in the middle crustduring D3 episode by less-efficient dislocation creep.Local hydration produced partial melt that segregatedfrom solid matrix into isolated bands instead ofcoating grain boundaries.ACKNOWLEDGEMENTSThe work was supported by Grants of the Czech ScienceFoundation (GACR 205 ⁄ 05 ⁄ 2187 and205 ⁄ 09 ⁄ 0539) and the internal project of the CzechGeological Survey (No. 326700). Stays of J. Franěkat Strasbourg University were funded by theFrench Government Foundation (BGF). The FrenchNational Grant Agency (No.06-1148784) and GrantMSM0021620855 of the Ministry of Education of theCzech Republic are acknowledged for financial supportof K. Schulmann and O. Lexa. We are grateful toJ.-E. Martelat for constructive suggestions. We alsothank R. Vernon, H. Stu¨ nitz and J. Tullis for theirconstructive revisions. We also thank M. Brown andR. White for careful editorial work.CONCLUSIONSREFERENCESAbart, R., Petrishcheva, E., Rhede, D. & Wirth, R., 2009.Exsolution by spinodal decomposition: II: perthite formationduring slow cooling of anatexites from Ngornghoro, Tanzania.American Journal of Science, 309, 450–475; doi: 10.2475/06.2009.02.Baratoux, L., Lexa, O., Cosgrove, J.W. & Schulmann, K., 2005.The quantitative link between fold geometry, mineral fabricand mechanical anisotropy: as exemplified by the deformationof amphibolites across a regional metamorphic gradient.Journal of Structural Geology, 27, 707–730.Behr, H.J., 1961. 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