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Age, lithology and structural evolution of the c. 3.53 Ga Theespruit Formation in the 

Tjakastad area, southwestern Barberton Greenstone Belt, South Africa, with 

implications for Archaean tectonics 

M.J. Van Kranendonk a, ⁎, A. Kröner b , E. Hegner c , J. Connelly d 

a Geological Survey of Western Australia, 100 Plain St., East Perth, WA 6004, Australia 

b Institut für Geowissenschaften, Universität Mainz, 55099 Mainz, Germany 

c Department für Geo-und Umweltwissenschaften, Universität München, Theresienstrasse 41, D-80333 München, Germany 

d Department of Geological Sciences, University of Texas at Austin, Austin, Texas, 78712-1101, USA 

article info 

Article history: 

Accepted 4 November 2008 

Keywords: 

Archean 

Barberton greenstone belt 

Theespruit Formation 

Tectonics 

1. Introduction 

abstract 

A major debate in the Earth sciences concerns the style of early 

Archaean tectonics, whether dominated by vertical movements due to 

vigorous mantle convection, plume-derived magmatism, and crustal 

diapirism, or horizontal plate interactions akin to the modern tectonic 

paradigm (e.g. de Wit, 1998 versus Hamilton, 1998). Although both 

tectonic processes have been separately identified as affecting 

different Archaean cratons, or different parts of such cratons (e.g., 

Myers and Kröner, 1994; Van Kranendonk et al., 2007), it is the relative 

contribution of these two processes that is still widely debated. This is 

particularly true regarding the formation of Palaeoarchaean granitoidgreenstone 

terrains (GGTs) that have a characteristic dome-and-basin 

map pattern, the likes of which do not occur in modern geological 

environments (Anhaeusser et al., 1969). 

This debate is particularly vigorous regarding the formation of 

Archaean crust in the Barberton Mountain Land part of the Kaapvaal 

Craton, southern Africa (Fig. 1). On the one hand, Viljoen and Viljoen 

⁎ Corresponding author. 

E-mail address: martin.vankranendonk@doir.wa.gov.au (M.J. Van Kranendonk). 

Chemical Geology 261 (2009) 115–139 

Contents lists available at ScienceDirect 

Chemical Geology 

journal homepage: www.elsevier.com/locate/chemgeo 

0009-2541/$ – see front matter. Crown Copyright © 2008 Published by Elsevier B.V. All rights reserved. 

doi:10.1016/j.chemgeo.2008.11.006 

A field and petrographic re-assessment of rocks from the Theespruit Formation near Tjakastad, southwestern 

Barberton Greenstone Belt, is combined with new zircon U–Pb ages and whole-rock Sm–Nd isotopic results 

to reveal several important inconsistencies with a previous thrust-accretion model for this area. It was found 

that faults previously interpreted as ‘thrusts’ are extensional, ‘thrust slices of basement orthogneiss’ are little 

deformed quartz-feldspar porphyries or sheared felsic volcaniclastic sedimentary rocks, and ‘sedimentary 

diamictites’ are felsic agglomerates. These observations, combined with consistent facing directions of 

bedding, the recognition of distinct stratigraphic variations across strike, and U–Pb zircon dates from 13 felsic 

metavolcanic samples, all suggest that the Theespruit Formation is a moderately strained, coherent 

stratigraphic succession deposited at c. 3530 Ma and affected by extensional shear deformation at c. 3230 Ma. 

A 3453±6 Ma population of zircons analysed previously from an amphibolite-facies Theespruit Formation 

felsic schist is interpreted as being of metamorphic, rather than detrital, origin, and arising from heat 

associated with felsic plutonic rocks that stitch the Theespruit Formation to the overlying Komati Formation 

at ca. 3450 Ma, 220 Ma earlier than the proposed thrusting event. 

Crown Copyright © 2008 Published by Elsevier B.V. All rights reserved. 

(1969) first interpreted the evolution of this area in terms of a littledeformed 

volcano-sedimentary (greenstone) succession, the Barberton 

Greenstone Belt, intruded by granitoid magmas and then tilted 

and folded by the vertical rise of tonalitic diapirs (Anhaeusser, 1984). 

However, the diapiric model was challenged by de Wit and colleagues 

who identified nappes, reverse faults, high strain zones, and local 

repetitions of stratigraphy in the greenstones (de Wit, 1982, 1983, 

1991; de Wit et al., 1983, 1987a). These authors deduced that the 

Barberton Mountain Land part of the Kaapvaal Craton formed through 

horizontal thrust-accretionary plate tectonic processes in which 

granitoid diapirism played no part (de Wit et al., 1992). 

Whereas nappes and reverse faults may form in a variety of 

environments, including diapirism (e.g. Ramberg, 1967; Dixon and 

Summers, 1983), the single critical piece of geological evidence 

supporting the thrust-accretion hypothesis is the interpreted largescale 

occurrence of older-over-younger rocks in the Theespruit area, 

located in the southwestern part of the Barberton Greenstone Belt 

(Fig. 1) (de Wit et al., 1983, 1992; Armstrong et al., 1990). In this area, 

ca. 3.49–3.48 Ga (Lopez-Martinez et al., 1992; Dann, 2000) komatiites, 

basalts and a layer of felsic tuff of the Komati Formation occur 

structurally above interlayered pillow basalt, sedimentary ‘diamictite’, 

and ‘orthogneiss’ of the Theespruit Formation, which has been

116 M.J. Van Kranendonk et al. / Chemical Geology 261 (2009) 115–139 

Fig. 1. Simplified geological map of southwestern part of the Barberton Mountain Land, Kaapvaal Craton, showing the location of Figs. 3 and 4 (rectangle), and of Fig. 7a, b, and c 

(dots). Inset shows location of the Barberton Greenstone Belt in the Kaapvaal Craton, within southern Africa. Age data for granitic rocks cited in text. 

interpreted to be ≤3458 Ma – and therefore younger than the Komati 

Formation – based on a population of zircons from a felsic ‘diamictite’ 

within the Theespruit Formation (Armstrong et al., 1990). The 

kilometre-wide Komati schist zone (KSZ) that separates the two 

formations in the southwestern part of the Barberton Greenstone 

Belt was interpreted as a plate-boundary thrust (de Wit et al., 1992) 

(Fig. 1). In this model, the Komati Formation was interpreted to 

represent structurally thickened oceanic crust (“Jamestown Ophiolite 

Complex”; de Wit et al., 1987a), obducted onto a “tectono-stratigraphic 

assemblage” of Theespruit Formation rocks across the KSZ at 

ca. 3230 Ma (de Wit et al., 1983, 1987a, 1992; Kamo and Davis, 1994). 

Tectonic thickening was interpreted to have given rise to partial 

melting of the lower parts of the overthickened crust and the 

generation of granitoid rocks which were emplaced into active thrusts 

(de Wit et al., 1987b). Evidence of local high-pressure metamorphism 

from the granitoid domains flanking the belt has also been used to 

support models of continental collision (e.g., Dziggel et al., 2001, 2002; 

Diener et al., 2006; Stevens and Moyen, 2007). 

However, several papers have cast doubt on the thrust-accretion 

hypothesis for this part of the Barberton Greenstone Belt. First, 

detailed volcanological studies have shown that a section with 

proposed sheeted dykes within the Komati Formation, which de Wit 

et al. (1987a) suggested was part of an ophiolite, is rather a series of 

thick flows with primary spinifex textures and sills or massive flows 

(Cloete and King, 1991; Cloete, 1999; Dann, 2000). In addition, more 

detailed volcanological studies indicate that the Komati and Hooggenoeg 

Formations, the latter of which contains felsic volcanic rocks, 

represent largely intact, autochthonous volcanic stratigraphy, more 

than 6 km thick (e.g., Cloete, 1999; Dann, 2000; Dann and Grove, 

2007), and is thus very different from modern ophiolite sections. A 

gabbro interpreted as part of the ophiolite stratigraphy was shown to 

be some 100 Ma younger than the Komati volcanics (Kamo and Davis, 

1994). These data preclude an ophiolite interpretation for the lower 

Onverwacht Group. 

Second, Byerly et al. (1996) showed that the greenstones above the 

KSZ are a contiguous, progressively upward-younging succession 

(cf. Viljoen and Viljoen,1969) which formed over more than 200 Ma of 

Earth history. This implies that much of the succession is not a 

tectono-stratigraphic assemblage as implied in the thrust-accretion 

hypothesis (de Wit et al., 1987a) — at best, an accretion component is 

reduced to a single thrust (the KSZ). 

Third, Kisters and Anhaeusser (1995) showed that granitoid rocks 

flanking the greenstones show several features that are consistent with 

emplacement of granitoids as diapirs, thus indicating a more complicated 

story than proposed by the thrust-accretion hypothesis alone. 

The above, when combined with more detailed observations 

presented below, are incompatible with the proposed thrust-accretion 

hypothesis, as it has previously been described. The cumulative weight of 

these arguments, together with our new data, requires a re-assessment 

of the critical Tjakastad area, the results of which are presented here. 

The new results indicate that the published thrust-accretion model 

is unlikely to have been the primary tectonic mechanism responsible 

for development of the Palaeoarchaean rocks in the Tjakastad area. 

Instead, the structural features of this area are more consistent with 

formation through partial convective overturn of a dense succession of 

continually upward-younging, dominantly volcanic, supracrustal 

rocks and a structurally lower, more buoyant layer of intraplated 

granitoid sheets, following previously published diapiric models. The 

conclusions have important implications for early Archaean tectonic 

models, as the Barberton area has been widely cited as a key example 

in support of modern plate-tectonic processes operating in the ancient 

past (cf. Windley, 1995).

2. Regional geology and previous geochronology 

The Barberton Mountain Land part of the Kaapvaal Craton is a 

Palaeoarchaean (3.55–3.20 Ga) granite–greenstone terrain that forms 

part of the eastern Kaapvaal Craton (Brandl et al., 2006). The volcanosedimentary 

strata, known as the Barberton Supergroup, are preserved 

in the tightly folded, ENE trending Barberton Greenstone Belt (BGB), 

which is surrounded by granitoid rocks of a variety of ages (Lowe and 

Byerly,1999). The most recent overview of the BGB is by Lowe and Byerly 

(2007), who also provide an extensive literature survey. 

2.1. Barberton Supergroup 

The Barberton Supergroup comprises a basal Onverwacht Group of 

komatiitic, basaltic and felsic volcanic rocks, and conformably to 

unconformably overlying argillaceous and arenaceous metasedimen- 

M.J. Van Kranendonk et al. / Chemical Geology 261 (2009) 115–139 

tary and minor felsic volcanic rocks of the Fig Tree and Moodies 

Groups, respectively (Fig. 2) (Viljoen and Viljoen, 1969; Brandl et al., 

2006). The Onverwacht Group has been broadly subdivided into a 

lower ultramafic unit and an upper mafic to felsic unit, separated by a 

prominent unit of chert known as the Middle Marker (Viljoen and 

Viljoen, 1969). The lower and upper units are each composed of three 

formations. The structurally lowermost is the undated Sandspruit 

Formation composed of strongly deformed and metamorphosed 

mafic–ultramafic schists adjacent to, and infolded with, TTG rocks in 

the southwestern part of the BGB. Apparently conformably overlying 

the Sandspruit Formation is the heterolithic Theespruit Formation 

which regionally contains komatiites and basalts and characteristic 

units of metamorphosed felsic lava and tuff and which, in the 

Tjakastad area, has been interpreted to also include slivers of 

orthogneiss and diamictite (de Wit et al., 1983). The Theespruit 

rocks occur in two areas separated by younger granitoid intrusives, 

Fig. 2. Generalized stratigraphic section of Barberton Supergroup, showing available age constraints (see text for references). SF = Sandspruit Formation. Modified from Lowe and 

Byerly (2007) and Byerly et al. (1996); age data cited in text. 

117


namely around and east of the village of Tjakastad and partly in the 

Songimvelo Game Reserve (in this paper referred to as the Tjakastad 

area), and farther east in the area around Steynsdorp, close to the 

border between South Africa and Swaziland; here referred to as the 

Steynsdorp area (Fig. 1). 

Single zircon 207 Pb/ 206 Pb evaporation ages of 3548±3 to 3544± 

3 Ma for four separate samples of felsic tuff from the Theespruit 

Formation in the Steynsdorp area and flanking the ca. 3510 Ma 

Steynsdorp Pluton provide a well-constrained date of deposition for 

this unit (Kröner et al.,1996). Zircon xenocrysts in the Steynsdorp Pluton 

vary in age from 3553 to 3531 Ma, providing further evidence on the age 

of the underlying Theespruit and Sandspruit Formations (op cit.). In the 

Tjakastad area, along strike to the west, a sample from what was 

interpreted as a thin sliver of basement orthogneiss (but see below) in 

the Theespruit Formation yielded similar results of 3538±6 Ma and 

3538+4/−2 Ma, respectively, in two separate zircon dating efforts 

(Armstrong et al., 1990; Kamo and Davis, 1994). A felsic volcaniclastic 

rock from immediately above the so-called basement gneiss sliver, but 

still within the Theespruit Formation, yielded zircons with identical ages 

of 3538±10 Ma (Armstrong et al., 1990). However, this sample also 

contained a second population of zircons with an age of 3453±6 Ma, 

and it is this result that has been widely cited as a maximum age for the 

deposition of the Theespruit Formation (see below: Armstrong et al., 

1990). An arkosic rock from the Sandspruit Formation to the west of our 

study area contained concordant zircons at 3521±6 Ma, with older 

grains up to ca. 3540 Ma (Dziggel et al., 2001, 2002). 

The Sandspruit and Theespruit Formations are separated from the 

overlying, lower-grade, Komati Formation by a broad zone of ductile 

shear, 600–800 m wide, and known as the Komati schist zone (KSZ). The 

overlying Komati Formation is composed of komatiites and interlayered 

basalts dated at between ca. 3490 Ma (Lopez-Martinez et al., 1992)and 

3481 Ma (Dann, 2000). The Middle Marker at the top of the lower 

ultramafic unit is a persistent horizon of silicified tuffaceous rocks dated 

at ca. 3472 Ma (Armstrong et al., 1990), the same age as quartz-feldspar 

porphyry dykes and sills intruding the Komati Formation (Kamo and 

Davis, 1994). 

Conformably overlying the Middle Marker is the Hooggenoeg 

Formation of tholeiitic basalt, komatiitic basalt, komatiite, and felsic 

volcanic flows and sills, dated at 3470 Ma (Byerly et al., 2002). Intruding 

the top of the Hooggenoeg Formation is a 3457–3438 Ma subvolcanic sill 

and related felsic lavas and volcaniclastic rocks (de Wit et al., 1987b; 

Kröner and Todt, 1988; Armstrong et al.,1990; Kröner et al., 1991; Byerly 

et al., 1996; de Vries et al., 2006). These rocks are locally eroded and 

conformably overlain by metamorphosed flow and volcaniclastic rocks 

of the Kromberg and Mendon Formations, dated at ≤3416 and 

N3298 Ma (Byerly et al., 1996; Lowe and Byerly, 2007). 

The Fig Tree Group conformably overlies the Onverwacht Group and 

is largely composed of clastic and pelitic metasedimentary rocks, but 

includes felsic volcanic and volcaniclastic rocks dated at between 3258 

and 3226 Ma, the same age as the Kaap Valley Tonalite and related 

intrusions to the north of the BGB (Kröner et al., 1991; Kamo and Davis, 

1994; Byerly et al., 1996; Lowe and Byerly, 2007). The Fig Tree Group is 

conformably to unconformably overlain by conglomeratic and arkosic 

units of the Moodies Group which, near the top of the section, contain 

several intra-formational unconformities indicative of syntectonic 

deposition in a contractional regime (Lamb, 1987). Deposition of the 

Moodies Group commenced about 3230 Ma and ended some time 

before 3110 Ma (Huebeck and Lowe, 1994). Synchronous folding and 

tilting of the entire Barberton Supergroup to near vertical occurred 

before 3216+2/−1 Ma, the age of the discordant and undeformed 

Dalmein Pluton (Fig. 1; Kamo and Davis, 1994). 

2.2. Granitoid rocks 

Tonalitic, trondhjemitic and granodioritic (TTG) rocks intrude and 

envelope the BGB and may be divided into five age groups. The oldest 

of these is the trondhjemitic Steynsdorp Pluton, dated at between 

3510 and 3502 Ma (Kröner et al., 1996). The second, more voluminous 

group of plutons is represented by linked, foliated to gneissic TTG 

rocks of the Stolzburg, Theespruit and Doornhoek plutons along the 

southwestern flank of the BGB. These plutons are not isolated 

intrusions but represent three domal culminations, or lobes, of a 

compositionally and geochronologically homogeneous granitoid layer 

separated by infolded, synformal septae of strongly deformed greenstones 

of the Sandspruit and Theespruit Formations (Fig. 1). These 

rocks display a characteristic dome-and-keel geometry in which 

radially-plunging lineations in the greenstone keels trend away from 

granitoid domes and increase in intensity and plunge towards the 

cores of the tightly folded greenstone keels, where the rocks are 

deformed into L-tectonite schists (Kisters and Anhaeusser, 1995). 

Anhaeusser (1984) and Kisters and Anhaeusser (1995) showed that 

the domal plutonic lobes were not formed by cross-folding but display 

characteristic features of diapiric emplacement. These plutonic rocks 

have been dated between 3469 and 3437 Ma, the same age as quartzfeldspar 

porphyry dykes intruding the Komati Formation and felsic 

volcanism in the Hooggenoeg Formation (Fig. 2: Kröner and Todt, 1988; 

Armstrong et al., 1990; Kamo and Davis, 1994; Kröner et al., 1996). 

Titanite from a 3476±10 Ma quartz-feldspar porphyry dyke cutting the 

Komati Formation in the Tjakastad area was dated at 3458±2 Ma and 

interpreted to represent the minimum age of intrusion of the dyke 

(Kamo and Davis, 1994). A second zircon generation from the ca. 

3445 Ma Stolzburg Pluton, considered by Kamo and Davis (1994) to be of 

metamorphic origin, is ca. 3237 Ma. Metamorphic titanite from the ca. 

3445 Ma Doornhoek and Stolzburg plutons was dated at ca. 3215 Ma and 

ca. 3201 Ma, respectively (Kamo and Davis, 1994). 

The third group of plutons is represented by the foliated Kaap 

Valley Tonalite and related granitoids along the northern flank of the 

BGB, dated at 3229±5 Ma (Tegtmeyer and Kröner, 1987), the same 

age as felsic volcanism in the Fig Tree Group (Kröner et al., 1991). The 

undeformed Dalmein Pluton was emplaced at 3216+2/−1 Ma 

(Kamo and Davis, 1994). Finally, a number of potassic granitoids, all 

∼3105 Ma in age, surround the BGB, among them the 3105±3 Ma 

Mpuluzi Batholith in the south (Kamo and Davis, 1994). These coarsegrained, 

porphyritic granites are largely undeformed. 

2.3. Metamorphism 

Viljoen and Viljoen (1969) identified an amphibolite-facies 

metamorphic aureole in greenstones around granitoids that surround 

the BGB. de Wit (1983) noted that greenstones of the Theespruit 

Formation adjacent to the Theespruit Pluton in the Tjakastad area 

were transformed into garnet amphibolites, and aluminous felsic 

schists contained lineated kyanite porphyroblasts. 

Cloete (1993, 1999) found that pressure estimates from fluid 

inclusion geobarometry in the Komati and Hooggenoeg Formations 

increased with stratigraphic depth to a maximum of 3.9 kbar and that 

the obtained values correspond to the lithological thickness of the 

overlying Barberton Supergroup, thus indicating that these rocks had 

only been subject to burial metamorphism. Temperature estimates 

using the plagioclase-amphibole thermometer vary from 320–420 °C 

in the Hooggenoeg Formation, to 490–530 °C in the Komati Formation. 

Rocks adjacent to the Theespruit Pluton were affected by a dynamic 

metamorphic overprint in addition to the static metamorphism, as 

evidenced by retrograde temperatures of 280 °C. Prograde P–T paths 

for Komati and Hooggenoeg Formation rocks have a slightly anticlockwise 

shape, from which Cloete (1993, 1999) concluded that 

metamorphism of the BGB was not analogous to that in Phanerozoic 

thrust-accretion belts. 

Van Vuren and Cloete (1995) studied the metamorphism of rocks 

in and around the Theespruit Pluton. These authors found that 

amphibolite-facies metamorphic conditions within the pluton were 

3–5 kbar and 470 °C. Marginal greenstones recrystallised under

conditions of 4.8–7 kbar, 500 °C. The highest metamorphic grade was 

found in the narrow septum of strongly deformed greenstones 

between the Theespruit and Stolzburg Plutons (Fig. 1), where 

garnet-clinopyroxene mafic granulites yielded P–T estimates of 

9 kbar and 700 °C. 


Fig. 3. Geology of the Tjakastad area of the Barberton Greenstone Belt, according to de Wit et al. (1983), showing inferred thrust faults (f–f), inferred thrust packages (double arrows 1 

through 6), and “diatexites” within the Theespruit Formation (modified slightly after Fig. 4 of de Wit et al., 1983). Localities A, D, E of inferred basement orthogneiss are discussed in 

the text. Zircon age data at locality D are from Armstrong et al. (1990) and show the results of two dated zircon populations from a sample of the ‘basement orthogneiss’ (but see text 

for discussion). 

Dziggel et al. (2002) recorded high-P, low-T metamorphic conditions 

of 8–11 kbar and 650–700 °C in remnants of the Theespruit 

and Sandspruit Formations within the western part of the Stolzburg 

Pluton. Two phases of titanite were dated from rocks of this area, at 

3418 ± 8 Ma and 3236±5 Ma, indicating two periods of metamorphism 

Fig. 4. Geological map of the Tjakastad area, based on data presented in this paper, showing the localities of dated felsic volcaniclastic samples and samples analysed for Sm–Nd, as discussed 

in text. Blue line with rectangles denotes transition from amphibolite- to greenschist-facies. Numbers inside boxes refer to other figures in this paper. DP = Doornhoek Pluton. 

119

120 M.J. Van Kranendonk et al. / Chemical Geology 261 (2009) 115–139

(and deformation) (Dziggel et al., 2001, 2002). Similar observations 

were made by Diener et al. (2006) for the area just west of Tjakastad, 

with prograde conditions of 5.5–6.3 kbar and 490 °C. Higher grade 

conditions in rocks closer to the granitoids gave peak P–T estimates of 7– 

7.7 kbar and 550 °C, with retrograde assemblages of 3.8 kbar and 543 °C, 

indicating nearly isothermal decompression. The estimates of peak 

metamorphic conditions indicate low geothermal gradients of 20 °C/km 

for the western part of the Barberton Mountain Land (Dziggel et al., 

2002; Diener et al., 2006). A syn-tectonic tonalite dated at 3231±5 Ma 

from the Schapenburg schist belt in the southwestern part of the 

Barberton Mountain Land gives a maximum age of amphibolite-facies 

metamorphism in this area, as does a 3219 Ma date on metamorphic 

zircon from an amphibolite in the western Tjakastad schist belt, only 10– 

20 Ma after sediment deposition (Stevens et al., 2002; Kisters et al., 

2003; Diener et al., 2006). 

3. Details of the proposed thrust-accretion model for the 

Tjakastad area 

As evident from the above overview, there is only one exception to a 

generally conformable interpretation for the Barberton Supergroup. This 

occurs in the Tjakastad area, where purported evidence of large-scale 

tectonic inversion of older-over-younger rocks has been documented. de 

Wit et al. (1983) suggested that the Theespruit Formation in the 

Tjakastad area was a tectono-stratigraphic assemblage of tectonically 

intercalated mafic volcanic rocks, metasedimentary diamictite and 

associated quartz-sericite schist, quartz-feldspar porphyries, and slivers 

of banded granitoid basement gneiss. In this model, the Theespruit 

Formation was “…most readily explained as a set of imbricate slices… 

that represents an inclined (oblique?) section through an imbrication 

zone beneath a major thrust … represented by the Komati schist zone…” 

(de Wit et al., 1983, p. 27: word in italics added). 

The most critical field evidence in support of a tectono-stratigraphic 

interpretation for the Theespruit Formation in the Tjakastad 

area was the recognition of tectonic slivers and cobbles of inferred 

basement orthogneiss (de Wit et al., 1983). Orthogneiss slivers were 

recognised in three distinct settings: (1) as a 500 m long wedge within 

quartz-sericite schist, diamictite, and quartz-plagioclase porphyries 

within the KSZ (locality E on Fig. 3); (2) as tectonic slivers within the 

Theespruit Formation, in one location being unconformably overlain 

by sedimentary diamictite (locality D on Fig. 3); (3) as cobbles in 

metasedimentary diamictite (locality A on Fig. 3). In setting 1, de Wit 

et al. (1983) suggested that the orthogneiss had experienced an 

earlier deformation history as indicated by isoclinal folds of 

metamorphic layering, the axial planes of which were oriented at a 

high angle to the regional foliation (Fig. 6b and c of de Wit et al., 

1983). Similarly, cobbles of orthogneiss within diamictite in setting 3 

were interpreted to exhibit a pre-existing gneissosity oriented at a 

high angle to both bedding and the regional upright cleavage in the 

matrix diamictite. 

That thrusting had occurred in the Tjakastad area was interpreted 

from the penetratively deformed nature of the rocks, the presence of 

faults that offset the stratigraphy and were interpreted to indicate a 

reverse sense of displacement, the presence of tectonic contacts as 

indicated by changes in the orientation and intensity of mineral 

elongation lineations across unit contacts, the occurrence of a southfacing 

bedding top direction within the KSZ, and along-strike differences 

in the thickness and components of the “tectono-stratigraphy”. 


Supporting a tectono-stratigraphic interpretation for the Theespruit 

Formation were zircon ages of 3538+4/−2 Ma and 3538 ± 

6 Ma (Armstrong et al., 1990; Kamo and Davis, 1994) for the 

“basement gneiss sliver” at locality D. At the time these results 

were obtained, they were the oldest recorded from the BGB and 

apparently confirmed a basement interpretation for these rocks relative 

to the overlying Komati Formation and younger conformable stratigraphy. 

Furthermore, euhedral to subhedral zircons with well-developed 

facets from the unconformably overlying ‘diamictite’ (which 

Armstrong et al., 1990 interpreted as a volcaniclastic rock), about 2 m 

above the gneiss wedge, yielded two distinct age populations: an older 

group at 3531±10 Ma, indistinguishable from the “gneiss sliver”,anda 

younger group at 3453±6 Ma. Although “…core and rim structures are 

common…” (Armstrong et al., 1990, p. 94) in zircons from these 

greenschist- to amphibolite-facies schists, the younger ages were 

interpreted to be derived from detrital grains and thus to indicate a 

maximum age of deposition for the Theespruit Formation. These data 

apparently confirmed both a younger age of the ‘diamictite’ relative to 

the gneiss sliver, and that the KSZ represents a major thrust across which 

older rocks of the Komati Formation (3490 Ma) were tectonically 

emplaced over the younger rocks of the Theespruit Formation 

(b3453 Ma). 

3.1. Controversial or unexplained elements of the previously proposed 

model 

Several features of both the geological interpretation by de Wit 

et al. (1983; Fig. 3) and the geochronology (Armstrong et al., 1990)do 

not appear easily reconcilable with a thrust-accretion origin for the 

rocks in this region: 

1) Why do the identified faults within the Theespruit Formation (f–f 

on Fig. 3) extend the lithological layering, rather than duplicate it 

as required for a thrust-accretion model? 

2) What is the sense of shear across the KSZ? If, as proposed, the KSZ 

represents a crustal-scale thrust, then it should contain evidence 

of reverse (Komati-over-Theespruit Formation) displacement. 

3) What is the origin of the ‘diamictites’ and how do they differ from 

volcaniclastic rocks or igneous intrusive breccias described 

elsewhere in the map area, or to felsic volcanic schist and metaagglomerate 

along strike to the east and west as described by 

Kisters and Anhaeusser (1995) and Kröner et al. (1996)? 

4) Why does a large unit of quartz-plagioclase porphyry with 

intrusion breccia textures pass laterally along strike into 

‘diamictite’ (just east of locality E in Fig. 3)? 

5) Why is there a discrepancy in the interpreted origin of felsic 

schists of the Theespruit Formation between the Tjakastad area 

(sedimentary diamictite and orthogneiss: de Wit et al., 1983) and 

those areas along strike to the west? 

6) What is the relationship between the proposed orthogneiss 

wedge (locality E in Fig. 3) and adjacent ‘diamictite’ and quartzplagioclase 

porphyries that surround, or structurally overlie it; 

does the wedge represent the limb of an inverted nappe, or is it 

located within a tectonic mélange? 

7) What role, if any, did the ca. 3450 Ma granitoid rocks play in the 

deformation of the Tjakastad area? Is it the same as for ca. 

3230 Ma granitoid rocks higher up in the stratigraphy which were 

emplaced during regional thrusting (cf. de Wit et al., 1987b), or do 

Fig. 5. Outcrop and thin section features of Theespruit Formation rocks: a) Well-preserved pillow basalts near the base of the formation. View is to the west-northwest, and drainage 

cavities (arrowed) and pillow tails (i.e. at hammer head) indicate way-up to north-northeast; b) Well-rounded cobble of fine-grained siliceous metasediment in foliated and lineated 

felsic agglomerate with stretched lapilli (not visible in photo): clast shows coarse layering defined by white and light grey colour banding; c) Cross-sectional view of crude layering 

parallel to bedding defined by thin, irregular sills of darker grey, weakly plagioclase feldspar-porphyritic dacite in fine-grained felsic volcaniclastic siltstone; d) Cross-sectional view of 

massive, coarse felsic volcaniclastic rock (agglomerate) with irregular, ragged-edged fiamme of feldspar-phyric, devitrified rhyolite glass (arrow) and well rounded clasts of chert; e) 

View, looking down on a horizontal surface, of alkali trachytic breccia: note angular, unstrained nature of clasts in this view; f) View, looking horizontally, on alkali trachytic breccia, 

showing strongly downdip lineated nature of fragments; g) Cross-polarised thin section view of alkali trachyte fragment in breccia, showing Carlsbad-twinned subhedral K-feldspar 

phenocryst in fine matrix of flow-aligned albite. Pen knife in b)–f) is 10 cm. All locations are shown in Fig. 4. 

121

122 M.J. Van Kranendonk et al. / Chemical Geology 261 (2009) 115–139

they display evidence of diapiric emplacement as previously 

proposed (Anhaeusser, 1984; Kisters and Anhaeusser, 1995)? 

8) What is the significance of high-grade metamorphic minerals 

(kyanite in pelitic schists and garnet in amphibolite) within the 

Theespruit Formation? Is the variation in metamorphic grade across 

the Tjakastad area consistent with, and exclusive to, a thrustaccretion 

model of formation, or can it be explained in other ways? 

9) What is the significance that the ‘basement gneiss sliver’ in the 

Tjakastad area is identical in age to felsic volcanic schists of the 

Theespruit Formation along strike to the southeast? 

10) What is the significance of the younger population of zircons in the 

dated, amphibolite- to greenschist-facies ‘diamictite’ (ca. 3453 Ma) 

which display core–rim overgrowth relationships and are identical 

in age to titanite in a quartz-feldspar porphyry sill cutting the 

Komati Formation (ca. 3458 Ma) and to widespread TTG plutonism 

(3470–3437 Ma) and felsic volcanism (3452–3438 Ma)? 

4. Results of field mapping and petrography 

Field mapping was undertaken in order to investigate the apparent 

inconsistencies in the thrust-accretion model for the Tjakastad area. The 

results are presented in Fig. 4, and below, under subheadings which 

individually address each of the apparent inconsistencies. The map 

differs in several key ways from that published by de Wit et al. (1983), 

particularly by way of different interpretations for the origin of several 

rock units in the Theespruit Formation, a stratigraphic interpretation for 

the Theespruit Formation, the nature of contact relationships between 

units, the position of the northeastern part of the Komati Fault, and the 

kinematics of faulting/shearing across the Tjakastad area. 

4.1. Lithology 

4.1.1. General stratigraphy 

The map area spans the transition from amphibolite-facies schists 

of the Theespruit Formation, which lie adjacent to the trondhjemitic 

Theespruit Pluton, to lower greenschist-facies komatiites of the 

Komati Formation across the high-strain KSZ (Fig. 4). The best and 

most continuous exposures are in a small streambed in the hills east of 

Tjakastad. The Theespruit Formation consists of lower pillowed and 

ocellar metabasalts (Fig. 5a), three compositionally and texturally 

distinct horizons of felsic schists and metachert, and interlayered 

mafic–ultramafic schists of indeterminate extrusive or intrusive 

origin. Way-up indicators of bedding are defined by graded bedding 

in felsic volcaniclastic layers and pillow structures in basalts. These 

consistently face to the northeast, away from the Theespruit Pluton 

(Fig. 4), except on the limbs of small-scale tight folds of bedding. 

Lithologic units of the Theespruit Formation continue into the KSZ, but 

then change across the Komati Fault into the base of the Komati 

Formation, which includes the type area of bladed olivine spinifextextured 

komatiites, massive ultramafic rocks, talc-serpentine schist, 

and intrusive quartz-feldspar porphyry. 

4.1.2. Felsic volcanic rocks 

Three compositionally and texturally distinct horizons of felsic 

rock occur in the Theespruit Formation (Fig. 4). The structurally and 


stratigraphically lowest felsic horizon is ≤65 m thick. It consists of 

irregularly layered felsic volcanic schist, grey-and-white layered, finegrained, 

cherty metasediment (siltstone/mudstone), kyanite-bearing 

quartz-sericite schist derived from aluminous greywacke, subordinate 

felsic tuff, and some massive felsic volcanic rocks. This horizon 

contains the dated unconformity locality that is discussed in detail in 

Section 4.1.4.2. The kyanite–andalusite schists contain faint relict 

bedding and lineated kyanite porphyroblasts that locally retain fresh 

cores of andalusite (see Section 4.3). The cherty metasediment is 

characterised by irregular, wispy layering and common chaotic folds 

that were possibly developed during, or soon after sedimentation, as 

they contain no axial planar foliation and have highly non-cylindrical 

fold axes, in contrast to folds of clearly tectonic origin elsewhere. 

Locally, cherty metasediment occurs as mis-oriented blocks in 

massive, fine-grained quartz-sericite rock (Fig. 5b). Nearby, the 

layered metasediment has an irregular contact with massive, weakly 

feldspar-phyric quartz-sericite rock that cuts bedding within the 

metasediment (Fig. 5c). These textures are interpreted as indicative of 

coeval sedimentation of felsic volcaniclastic detritus and felsic 

magmatism, the latter formed as either a felsic volcanic flow or a sill 

emplaced directly below the sediment–water interface. 

The middle horizon of felsic schist, which de Wit et al. (1983) 

interpreted to contain extensive sedimentary diamictite with local clasts 

of basement granitoid orthogneiss, is similar in thickness to the lower 

horizon, but is texturally heterogeneous along strike. The best exposures 

of the coarse clastic rocks are in the small streambed mentioned above, 

where they comprise almost the entire thickness of the horizon. This 

unit contains subrounded to subangular, pebble to boulder size clasts of 

abundant grey to black chert, less abundant white and grey banded 

chert, clastic metasediment, and fine-grained felsic volcanic rock, and 

rare mafic volcanicrockandgraniteinafine grained quartz-sericite 

matrix (Fig. 5d). We interpret these rocks to be proximal felsic volcanic 

pyroclastic agglomerates and not sedimentary diamictite, because: 

1. Well-preserved parts of the unit contain common, irregular, 

ragged-edged fiamme of feldspar-phyric, devitrified rhyolite glass 

(arrow in Fig. 5d), whose delicate shapes could not have survived 

transport by water together with the rounded boulders of chert, as 

they would have almost instantly been disaggregated. 

2. The unit locally contains elliptical bombs of vesicular to pumiceous 

felsic volcanic rock in an otherwise massive, fine-grained felsic 

matrix that lacks bedding. 

3. The coarsely fragmental rocks lack bedding, are unsorted, and 

laterally pinch and swell along strike, grading into felsic lavas and/ 

or fine-grained volcaniclastic sediment. 

4. In thin section, the matrix is seen to be generally fine-grained with 

no evidence of original detrital grains, and it contains elliptical to 

irregular-shaped patches of slightly coarser quartz–biotite–plagioclase 

with serrate, gradational contacts with the matrix, that are 

interpreted to represent fiamme. 

The stratigraphically highest felsic horizon locally displays metreto 

centimetre-scale bedding and is characterised by layers of quartzphyric 

felsic porphyry with subordinate felsic schist, fine-grained 

quartz-biotite metasedimentary rock of possible volcaniclastic origin, 

and minor felsic volcanic breccia. 

Fig. 6. a) Representative outcrop example of the quartz-feldspar porphyry intrusion at ‘basement orthogneiss’ locality of de Wit et al. (1983) (locality E in Fig. 3 of this paper), 

showing medium-grained, equigranular texture and crude igneous layering; b) Cross-polarised thin section view of a), showing unstrained, euhedral plagioclase feldspar 

phenocrysts in weakly foliated matrix of fine-grained quartz and feldspar; c) boulder of unstrained quartz-feldspar porphyry – previously interpreted as basement orthogneiss (de 

Wit et al., 1983) – with well defined igneous flow banding that outlines isoclinal folds; d) Closeup view of hinge region of isoclinal igneous flow fold in c), showing unstrained, 

bipryramidal quartz phenocrysts (arrows); e) Cross-polarised thin section view of rock shown in parts c) and d), showing unstrained, euhedral bipyramidal quartz phenocrysts in an 

undeformed matrix of fine-grained quartz and feldspar: unstrained, euhedral plagioclase phenocrysts are also visible in this thin section, but are not shown; f) Outcrop view, looking 

south-southeast along strike, of the ‘unconformity locality’ of de Wit et al. (1983) (locality D in Fig. 3 of this paper), showing foliated quartz-sericite schist (right side, under penknife) 

and breccia-textured unit (left side, oblique view onto bedding/foliation plane); g) Plane polarised light thin section view of homogeneous quartz-sericite schist from the 

‘unconformity locality’, showing tourmaline crystals and chloritoid porphyroblasts; h) View looking west-southwest onto steeply dipping bedding/foliation surface of felsic volcanic 

breccia unit shown on the left side of (f), showing angular nature of felsic clasts in this volcaniclastic breccia unit and well-developed stretching lineation at this outcrop. Pen knife in 

a), c) and f) is 10 cm long. All locations are shown in Fig. 4. 

123


Fig. 7. a) Large, unstrained amphibolite-facies pillow basalt in a 100–200 m xenolith from core of Theespruit Pluton (see Fig. 1 for location): pillows face to the southeast; b) Vertical 

cross-sectional view of strongly lineated, amphibolite-facies pillows within narrow greenstone septum between lobes of the Theespruit Pluton; c) Sheeted, heterogeneous margin of 

Theespruit Pluton, which contrasts with the unstrained, homogeneous core of the pluton; d) Rodded L-tectonite felsic volcanic schists from the KSZ, used as building materials; 

e) Plane polarised light view of a vertically-oriented thin section, looking southeast, showing SW-side-up displacement in sheared metabasalt from the KSZ (NE = northeast, SW = 

southwest); f) Plane polarised thin section view of aluminous felsic schist of Theespruit Formation, showing kyanite (Ky = dark grey) overgrowing andalusite (And = light grey, with 

well develop cleavage) along extension cracks, indicative of an increase in pressure during metamorphism and the development of linear fabrics in these rocks. Locations are shown 

in Figs. 1 and 4.

Two thin units of strongly sheared, texturally variable felsic schist 

in the western part of the KSZ are similar to the lower felsic unit in the 

Theespruit Formation. The topmost unit contains fine-grained tuff, 

grey chert, and fine-grained quartzite. In the southern-central part of 

the KSZ, a ∼100 m thick unit of felsic schist that de Wit et al. (1983) 

interpreted as sedimentary diamictite varies, from base to top, from a 

massive rock with plagioclase-phyric texture, a middle unit of felsic 

volcanic breccia with elliptical, ragged-shaped inclusions of carbonate-altered 

felsic material, and ≤50 m of grey and white layered 

chert. 

In the eastern part of the KSZ in the study area, de Wit et al. (1983) 

described an intrusion breccia within the core of a quartz-plagioclase 

porphyry intrusion that they interpreted to pass along strike 

westwards into sedimentary diamictite (see Fig. 3). However, it was 


Fig. 8. a) Structural map of the Tjakastad area, showing stretching lineations, kinematic indicators, and areas of high strain within the Komati schist zone (KSZ). Locations of transects 

in b) are also indicated; b) Transects across the Theespruit Formation and KSZ, showing variations in orientation of stretching lineations (see text for discussion). 

found that the rocks in this area comprise two distinct units (see 

Fig. 4): 1) an intrusive unit of quartz-feldspar porphyry to the north 

(described below); and 2) a fine-grained, highly siliceous felsic 

volcanic rock with closely packed, monomict, angular breccia 

fragments (Fig. 5e). Detailed outcrop mapping showed that parts of 

this unit display crude layering on a metre scale, with lenses or layers 

of massive rock alternating with layers of homogeneous, monomict, 

angular breccia, and other layers of less well sorted breccia with 

irregular-shaped blocks of massive rock up to a metre in size. Overall, 

bedding becomes more clearly defined at a finer-scale to the north, 

suggestive of graded bedding and a top to the north facing 

direction. Despite excellent preservation of igneous textures on 

horizontal outcrops, the rocks have been deformed into verticallyplunging 

L-tectonites (Fig. 5f). In thin section, the massive unit at this 

125


locality is characterised by euhedral K-feldspar phenocrysts with 

Carlsbad twinning in a matrix of fine, flow-aligned albite with 

trachytic texture (Fig. 5g). Quartz phenocrysts are lacking in this 

lithology, and thus the rock is classified as an alkali trachyte breccia. 

The textural uniqueness of each of the felsic horizons in the 

Theespruit Formation and the KSZ indicate that they do not represent 

a tectonically-duplicated series of thrust slices as suggested by de Wit 

et al. (1983). Rather, the above data indicate distinct lithological 

variations across strike, and the presence of consistent bedding top 

indicators to the north in both pillowed mafic volcanics and felsic 

volcanic/volcaniclastic horizons suggests that the rocks of the 

Theespruit Formation in the Tjakastad area, through into the KSZ, 

represent a strained, but otherwise intact volcano-sedimentary 

succession. This hypothesis is discussed more fully in Section 4.2. 

4.1.3. Granitic dyke in the Theespruit Formation 

A7–8 m wide sheared granite dyke occurs at one locality in the small 

stream bed to east of Tjakastad, about 100 m north of the Theespruit 

Formation/Theespruit Pluton contact. The foliation in this granite is 

parallel to that in the adjacent Theespruit schists, and the dyke could not 

be followed along strike due to poor exposure. The dyke is similar in 

appearance to granitic to aplitic dykes and veins emplaced into the 

Theespruit and Stolzburg plutons in the Tjakastad region. The dyke was 

sampled for zircon dating, and the results are reported below. 

4.1.4. “Granitoid gneiss” 

The detailed map of the Theespruit area by de Wit et al. (1983) 

shows the precise locations where exposures of granitoid gneiss were 

thought to occur in three distinct tectonic settings. These included a 

large wedge of orthogneiss in the northeastern part of the KSZ, tectonic 

slivers in the lower part of the Theespruit Formation, and cobbles 

within sedimentary “diamictite”. Re-examination of these easily 

recognisable sites and detailed petrographic studies of collected 

samples indicate that granitoid rocks were misidentified at these 

localities and that no such rocks occur in the Theespruit area. Instead, 

these rocks are re-interpreted as either quartz-plagioclase porphyries 

or volcaniclastic metasedimentary rocks, for reasons detailed below. 

4.1.4.1. Northeastern wedge of granitoid gneiss. The wedge of 

‘granitoid gneiss’ in the northeastern part of the map area (locality E in 

Fig. 3) consists of a homogeneous, medium-grained rock of trondhjemitic 

composition with characteristic euhedral, bipyramidal quartz and 

plagioclase phenocrysts (≤3 mm) scattered throughout a finer-grained 

quartzo-feldspathic matrix (Fig. 6a). In thin section, igneous textures are 

well preserved, consisting of subhedral to euhedral quartz and 

plagioclase phenocrysts in a fine-grained quartzo-feldspathic matrix 

that is overprinted by radiating sheaths of riebeckite needles (Fig. 6b). 

In contrast to the map of de Wit et al. (1983), this unit occurs 

outside, to the north of the KSZ, the northern margin of which is clearly 

marked by a discrete, fault-filled, 2 m wide quartz vein (Fig. 4). The 

rocks contain none of the strong LNS tectonite fabric characteristic of 

the KSZ (see Section 4.2 below); they are almost completely 

unstrained except for a very local, weak foliation developed within 

10 m of the KSZ. These rocks contain no leucosome veins, migmatitic 

textures, or gneissic layering with palaeosome–leucosome relations 

to indicate that they were derived through metamorphic segregation 

processes related to an earlier thermo-tectonic event, as previously 

interpreted (de Wit et al., 1983). The compositional layering which 

does occur in this unit is only locally developed in the southern part of 

the unit, and in only one outcrop does it outline isoclinal folds (Fig. 

6c; same as Fig. 6b and c of de Wit et al., 1983). This instantly 

recognisable outcrop with the isoclinally folded layering contains 

undeformed, euhedral bipyramidal quartz and plagioclase phenocrysts 

throughout, even across the hinge regions and limbs of the 

folds (Fig. 6d, e). The presence of these igneous phenocrysts indicate 

that the folded compositional layering represents igneous flow 

banding rather than a metamorphic segregation, and that the rock 

is thus a quartz-plagioclase porphyry intrusion, similar to the type 

dated by Kamo and Davis (1994) at ca. 3470 Ma; it is not a basement 

orthogneiss. 

Finally, the “evidence” cited by de Wit et al. (1983) that the 

isoclinally folded ‘granitoid gneiss’ at this locality had experienced an 

earlier phase of deformation on the basis that the folds were oriented 

at a high angle to the regional foliation, cannot be substantiated since 

the rock containing the folds is a misoriented boulder that has partly 

slid down the side of a hill. 

4.1.4.2. ‘Granitoid gneiss’ at the ‘unconformity’ locality. The ‘unconformity’ 

locality described by de Wit et al. (1983: locality D in Fig. 3)is 

underlain by highly strained schists over a strike length of 10 m and 

30 cm across. The interpreted unconformity in this outcrop consists of 

a lower unit (‘orthogneiss sliver’ of de Wit et al., 1983), 28 cm wide, of 

millimetre- to centimetre-scale interbedded quartz-sericite, quartzsericite-tourmaline-chloritoid, 

quartz-chlorite, sericite-quartz-chlorite-actinolite 

schist (Fig. 6f, g). The overlying schist unit (‘diamictite’ of 

de Wit et al., 1983) is 2 cm wide and displays a distinctive breccia of 

strongly lineated felsic fragments (L=62°→107°) on foliation surfaces 

(Fig. 6h). The fragments are composed of sericite (95%)-chloritetitanite 

up to 120×40×2 mm in size (X:Y:Z axes) within a matrix of 

slightly darker quartz-sericite-biotite-chlorite schist. These rocks are 

overlain by quartz-sericite schist, finely layered grey and white 

Table 1 

Location and field relations of samples for zircon dating and Sm–Nd isotopic analysis. 

Sample Location Rock type Field relationship 

BA56 S26°00′15.3″ 

E30°50′24.3″ 

Felsic tuff Bottom of 50 m wide layer 

BA57 S26°00′04.1″ 

E30°50′15.4″ 

Felsic tuff Top of 50 m wide layer 

BA62 S26°00′03.8″ 

E30°50′00.2″ 

Felsic tuff Close to Theespr. tonalite contact 

BA65 S26°00′04.3″ Granite Sheared dyke cutting Theespruit 

E30°49′59.8″ 

Fm. 

BA66 S26°00′00.7″ 

E30°50′06.9″ 

Amphibolite Interlayered with felsic tuff 

BA69 S25°59′54.6″ 

E30°50′05.3″ 

Amphibolite Interlayered with felsic tuff 

BA70 S25°59′54.7″ 

E30°50′04.9″ 

Felsic tuff 30 m wide layer 

BA71 S25°59′54.3″ 

E30°50′05.6″ 


BA73 S25°59′41.0″ 

E30°50′12.2″ 

Felsic tuff 20 m wide layer, very fresh 

BA74 S25°59′40.7″ 

E30°50′12.6″ 

Basaltic komat. 10 m wide pillowed layer 

BA75 S26°00′33.8″ 

E30°51′11.7″ 

Gabbro Sill in felsic tuff 

BA79 S26°01′34.5″ 

E30°53′36.8″ 


BA84 S26°00′32.3″ 

E30°53′31.5″ 

Felsic tuff Interlayered with chert 

BA85 S26°00′32.2″ 

E30°53′31.5″ 

Felsic tuff Layer 10 mN of BA 85 

BA 104 S25°59′30.5″ 

E30°50′08.2″ 

fsp-porphyry 25 thick layer in felsic tuff 

BA 105 S25°59′35.8″ 

E30°49′59.4″ 

fsp-porphyry ca. 25 m wide layer, fresh 

BA 107 S26°00′52.2″ 

E30°51′07.5″ 

Felsic tuff Closure of isoclinal fold 

BA 108 S26°03′00.9″ 

E30°53′01.8″ 

Felsic agglom. ca. 50 m thick layer, fresh 

BA 109 S26°02′21.8″ 

E30°53′34.1″ 

Felsic tuff ca. 80 m wide layer, bio-rich 

BA 110 S26°02′04.7″ 

E30°53′34.1″ 

Felsic tuff Same layer as BA 109, bio-free 

160932 S25°59′41.9″ Felsic volcanic Thick, coarse felsic volcaniclastic 

E30°49′52.7″ 

conglomerate 

All samples listed in the table are plotted on Fig. 4.

Table 2 

Isotopic data from evaporation of single zircons from rocks of the Theespruit Formation and one intrusive granite, Tjakastad–Songimvelo area, Barberton Greenstone Belt. 

Sample number Zircon colour and 

morphology 


Grain # Mass scans 1 

Evaporation 

temp. in °C 

Mean 207 /Pb/ 206 Pb ratio 2 

and 2-σm error 

207 Pb/ 206 Pb age 


BA 56 Clear to yellowish, 1 109 1599 0.310606±64 3523.9±0.3 

Long-prismatic, 2 150 1601 0.310751±67 3524.6±0.3 

Near-idiomorphic 3 83 1598 0.310486±139 3523.3±0.7 

4 102 1599 0.310533±93 3523.6±0.5 

Mean of grains 1–4 1–4 444 0.310593 ±44 3524.0 ±0.2 

BA 57 Yellow-brown, 1 128 1603 0.310214±148 3522.0±0.7 


Idiomorphic 3 82 1599 0.310337±148 3522.6±0.7 

Mean of grains 1–3 1–3 320 0.310336±80 3522.6±0.4 

BA 62 Clear to light 1 98 1598 0.310602±110 3523.9±0.5 

Yellow-brown, 2 111 1598 0.310678±118 3524.3±0.6 

Long-prismatic, 3 103 1600 0.310546 ±122 3523.6±0.6 

and stubby 4 52 1601 0.310623±164 3524.0±0.8 

Mean of grains 1–4 1–4 364 0.310612±62 3524.0 ±0.3 

BA 65 Clear to yellow- 1 64 1602 0.288251±126 3408.2 ±0.7 

Brown, long- 2 64 1600 0.288265±81 3408.3 ±0.4 

Prismatic, idiom. 3 65 1599 0.288197±65 3407.9±0.4 

Mean of grains 1–3 1–3 193 0.288238±54 3408.1±0.3 

BA 70 Clear, thin, long- 1 66 1595 0.310381 ±86 3522.8±0.4 

Prismatic, 2 88 1599 0.310363 ±64 3522.7±0.3 

Slightly rounded 3 66 1598 0.310962 ±73 3525.7±0.4 

Mean of grains 1–3 1–3 220 0.310548±56 3523.6±0.3 

BA 71 Dark red-brown, 1 75 1608 0.311461±51 3528.2±0.3 


Idiomorphic to 3 51 1602 0.311498±91 3528.4±0.5 

Slightly rounded 4 68 1600 0.311572±65 3528.7±0.3 

Mean of grains 1–4 1–4 252 0.311470±35 3528.2±0.2 

BA 73 Clear to yellowish, 1 52 1599 0.310174±92 3521.8±0.5 



Mean of grains 1–3 1–3 220 0.310241±53 3522.1±0.3 

BA 84 Clear, thin, long- 1 95 1595 0.312116±87 3531.4±0.4 

Prismatic, 2 44 1599 0.312437±87 3533.0±0.4 


Mean of grains 1–3 1–3 221 0.312273±71 3532.2±0.4 

BA 85 Clear, 1 74 1608 0.312382±125 3532.7±0.6 


Well rounded 3 96 1618 0.312393±110 3532.8±0.5 

Terminations 4 95 1612 0.312448 ±136 3533.1±0.5 

5 97 1620 0.312384±59 3532.7±0.3 

6 104 1614 0.312467±70 3533.1±0.3 

Mean of grains 1–6 1–6 581 0.312468±43 3533.2 ±0.2 

BA 105 Dark grey, round, 1 128 1597 0.311002 ±37 3525.9±0.2 

Detrital 2 82 1594 0.310763 ±46 3524.7±0.2 

Grains 3 92 1595 0.310867±58 3525.2±0.3 

Mean of grains 1–3 1–3 302 0.310896 ±43 3525.4±0.2 

BA 107 Grey-brown, 1 127 1600 0.312022±48 3531.0±0.2 


Well rounded 3 54 1601 0.311789±85 3529.8±0.4 

Terminations 4 114 1600 0.312024±52 3531.0±0.3 

Mean of grains 1–4 1–4 382 0.311919±37 3530.4±0.2 

BA 108 Clear, 1 110 1608 0.311212±45 3526.9±0.2 


Well rounded 3 142 1618 0.311146±35 3526.6±0.2 

Terminations 4 204 1612 0.311154±27 3526.7±0.1 

5 107 1620 0.311346±50 3527.6±0.2 

Mean of grains 1–5 1–5 656 0.311207±18 ⁎3526.9±0.2 

BA 109 Clear, 1 104 1608 0.311834±74 3530.0±0.4 


Well rounded 3 114 1618 0.311893±62 3530.3±0.3 

Terminations 4 64 1612 0.311882±99 3530.3±0.5 

Mean of grains 1–4 1–4356 0.311852±40 3530.1±0.2 

(continued on next page) 

127


Table 2 (continued) 

Sample number Zircon colour and 

morphology 

tuffaceous sandstone/siltstone, and wispy to irregularly layered felsic 

schist. Layering in the lowermost felsic unit (the so-called ‘granitoid 

gneiss sliver’ of de Wit et al., 1983) is parallel to the contact with the 

overlying brecciated unit, such that there is no evidence of an angular 

relationship to indicate the presence of an unconformity. Nor is there 

any textural evidence in the lower schist unit that metamorphic 

segregation (e.g., leucosome–palaeosome) processes, or introduced 

granitoid melt, were involved in the formation of the compositional 

layering. Rather, the high quartz content, the presence of tourmaline, 

and textural features of the layering are consistent with an interpretation 

as a sheared felsic, aluminous, metasediment. 

The origin of the overlying, brecciated unit is more obscure. 

However, the monomict nature of the ‘clasts’ and the petrographic 

evidence that the ‘clasts’ are similar in composition to the matrix 

suggests that this rock is a felsic agglomerate, as with other rocks in 

the Theespruit Formation. Such an origin is supported by the fact that 

this unit passes along strike to the northwest, as well as across strike, 

into the felsic volcanic rocks described above. 

4.1.4.3. ‘Basement orthogneiss’ cobbles in ‘diamictite’. The third 

setting in which ‘basement orthogneisses’ were identified by de Wit 

et al. (1983), was as boulders within ‘sedimentary diamictite’. 

Although the boulder depicted in Fig. 6a of de Wit et al. (1983) 

could not be relocated, boulders of identical appearance and texture, 

with a similar misoriented compositional layering with respect to the 

regional foliation, were observed at several places (e.g. Fig. 5b). All 

such boulders consist of fine-grained, siliceous, weakly and irregularly 

colour-layered rock, identical to the grey-white cherty metasediment 

(silicified tuff?) in the lowermost felsic unit. The irregular layering in 

the boulders is defined only by a change in colour, not in grain size or 

composition, as both grey and white layers are composed of N95% 

fine-grained silica. Locally, the darker layers are formed by remobilisation 

of grey-black chert into fractures within a white-cream quartz 

siltstone/chert. Again, no evidence of leucosome veins, metamorphic 

differentiation, or even a granitic composition was observed in the 

boulders to confirm an origin as orthogneisses. As such, and because 

of the volcanic origin of the host agglomerate, the layered boulders are 

re-interpreted here as fragments of underlying sediment entrained in 

felsic magma during explosive volcanism. 

4.2. Strain 

Grain # Mass scans 1 

4.2.1. Granite-greenstone contact 

In the map area, the Theespruit Pluton is a leucocratic, mediumgrained 

trondhjemite that intrudes the greenstones as a series of 

little-deformed, bedding-subparallel sheets (sills). Here it represents 

the structurally highest and least deformed part of the pluton and has 

intruded the highest stratigraphic level in the greenstones. Anhaeusser 

(1984) and Kisters and Anhaeusser (1995) described contact 

agmatites and ring dykes in other low-strain, high-level parts of the 

pluton; the ring dykes are parallel to bedding and thus probably 

represent sills turned on edge together with greenstones, during 

doming. 

Central parts of the pluton to the south of our map area are little 

deformed, medium- to coarse-grained leucocratic trondhjemite. Large 

Evaporation 

temp. in °C 

Mean 207 /Pb/ 206 Pb ratio 2 


207 Pb/ 206 Pb age 


BA 110 Dark grey, round, 1 88 1597 0.310479 ±52 3523.1±0.3 

Detrital 2 103 1594 0.310681±84 3524.3±0.4 

Grains 3 65 1595 0.310560±110 3523.7±0.5 

Mean of grains 1–3 1–3 0.310580±48 3523.8±0.2 

1 207 206 2 

Number of Pb/ Pb ratios evaluated for age assessment. Observed mean ratio corrected for non-radiogenic Pb where necessary. Errors based on uncertainties in counting 

statistics. ⁎Error of combined mean ages (bold print) is based on reproducibility of internal standard at 0.000027 (2σ). 

xenoliths of amphibolite-facies pillow basalt in the core of the pluton 

are almost completely undeformed, although they are gently tilted 

and face radially away from the core of the pluton, indicating its 

domical shape (Fig. 7a). Greenstones along the margins of the pluton 

were transformed into strongly foliated and lineated, medium- to 

coarsely recrystallised amphibolites and host a transposed network of 

granitoid veins emplaced in part during folding. In the narrow septum 

of greenstones between the Theespruit and Stolzburg plutons 

(Anhaeusser, 1984; Kisters and Anhaeusser, 1995), all rocks have 

been intensely deformed into subvertical L-tectonites (Fig. 7b) under 

metamorphic conditions up to granulite-facies (9 kbar, 700 °C; Van 

Vuren and Cloete, 1995). The steeply dipping margins of the 

Theespruit Pluton are strongly foliated and heterogeneous, with 

multiple foliated sheets of variable composition and textural complexity 

(Fig. 7c). Such variation was explained by Anhaeusser (1984) as 

due to progressive and differential assimilation of marginal greenstones 

by successive intrusive trondhjemite sheets and concomitant 

deformation of earlier sheets during successive intrusive pulses under 

ductile conditions imposed by the heat of the magmas. 

4.2.2. Theespruit Formation and KSZ 

Rocks in the Tjakastad region vary greatly in strain. In the 

southwest, adjacent to the Theespruit Pluton, little deformed rocks 

with well-preserved sedimentary and/or volcanic textures are 

preserved (e.g. pillows, ocelli, intra-pillow hyaloclastite breccia, and 

pillow shelves). Strain increases progressively towards the southern 

contact of the KSZ, which is marked by the Theespruit Fault, filled by a 

≤12 m wide quartz vein. The southern 300 m of the KSZ is underlain 

by high strain ultramylonites derived from mafic, felsic and cherty 

precursors (Fig. 8a). Strain then decreases again northeast through the 

remainder of the KSZ until the Komati Fault, a locally sharply defined 

structure filled by a quartz vein that marks the northern margin of the 

KSZ. North of the Komati Fault, only those rocks within ∼20 m of the 

fault display a weak foliation, but otherwise, rocks of the Komati 

Formation have not been penetratively deformed. Fabric elements 

across the area are generally co-axial, characterised by beddingparallel 

foliations that are locally axial planar to tight small-scale folds 

and lineations which progressively steepen in plunge away from the 

Theespruit Pluton (see below). This variation in strain accompanies a 

progressive decrease in metamorphic grade from amphibolite-facies 

in the southwest, adjacent to the granitoid rocks, to greenschist-facies 

in the northeast. 

Amphibolite-facies mafic volcanic rocks of the Theespruit 

Formation adjacent to the Theespruit Pluton contain a moderate 

S–L fabric (3:2:1) in which elongation lineations plunge moderately 

to the northwest (Fig. 8b). A narrow high strain zone is developed 

along the southern contact of the first felsic horizon, as noted by 

de Wit et al. (1983). North of this contact, lineations become more 

steeply plunging and more northerly trending in rocks of variable 

preservation, but generally the succession becomes more strongly 

deformed towards the KSZ. The trend and plunge of extension 

lineations across this zone vary between different rock types, as 

shown in Fig. 8b, only becoming fully transposed within the 

southern part of the KSZ where chloritic ultramylonites contain 

subvertical, NNW-SSE striking schistosities and downplunge


mineral elongation lineations. In the northeastern part of the study 

area, rocks of the KSZ are pure L-tectonites, with strongly developed 

vertical rodding lineations (Fig. 7d).Inthinsectionsofthemost 

highly strained schists, kinematic asymmetry could not be identified 

due to intense flattening strain. However, in slightly coarser-grained 

porphryoblastic and originally porphyritic rocks along the northern 

margin of the KSZ, kinematic indicators of SW-side-up displacement 

were observed (Fig. 7e). Northeast of the ultramylonitic schists, in 

the western part of the KSZ, strain decreases and local sedimentary 

features can again be recognised within felsic tuffaceous horizons. 

Pillow facing directions to the northeast were identified in basalts. 

In the eastern part of the KSZ, the strain is more heterogeneous, 

with high strain shear splays near its northern boundary (Fig. 4). 

The Komati Fault on the northern margin of the KSZ is exposed in 

the western part of the map area, to the south of which SW-side-up 

kinematic indicators were observed in serpentinized, schistose metakomatiite 

(Fig. 8). In the eastern part of the map area, the Komati Fault 

is represented by a quartz vein that separates L-tectonite mylonitic 

felsic agglomerates from undeformed quartz-feldspar porphyry at the 

base of the Komati Formation (Fig. 4). 

4.3. Metamorphism of the Theespruit Formation in the Tjakastad area 

Metamorphic grade decreases from amphibolite-facies adjacent to 

the Theespruit Pluton (garnet-actinolite pillowed mafic volcanics), to 

greenschist- facies in the KSZ, where mafic mylonites contain chloritecarbonate 

assemblages, felsic schists consist of sericite and quartz, and 

quartz-plagioclase porphyries contain chloritised riebeckite. As 

described above, P–T estimates decrease up stratigraphy from the 

western margin of the Theespruit Pluton (≤9 kbar, 700 °C; Van Vuren 

and Cloete, 1995) to the top of the Hooggenoeg Formation (Fig. 2: 

1.9 kbar, 320–420 °C; Cloete, 1993, 1999). 

The stratigraphically lowest felsic horizon in the Theespruit 

Formation contains abundant kyanite porphyroblasts that are elongate 

in the lineation direction and syn- to late-kinematic with respect 

to the deformation. Andalusite porphyroblasts contain round inclusions 

of an altered mineral that now consists of fine-grained quartzsericite 

and which may have originally been cordierite. Andalusite 

porphyroblasts are commonly overgrown by kyanite, along extension 

cracks, indicating replacement of the former by the latter under 

conditions of increasing pressure (Fig. 7f). The initial assemblage 

andalusite–cordierite indicates P–T conditions of b3.9 kbar, 550 °C 

(Spear, 1993), whereas the presence of overgrowing kyanite indicates 

an increase in pressure under near-isothermal conditions along part of 

an anticlockwise P–T path. As discussed below, these data have 

important implications for the tectonic evolution of the BGB. 

5. Zircon geochronology and Nd isotopic systematics 

Single zircons were dated from 13 felsic Theespruit lithologies and 

from one intrusive granite in the Tjakastad area, using the evaporation 

technique (Kober, 1986, 1987). The sensitive high-resolution ion 

microprobe (SHRIMP II; DeLaeter and Kennedy, 1998) and conventional 

analysis by thermal ionisation mass spectrometry (TIMS) 

Fig. 9. Histograms showing distribution of Pb isotope ratios derived from evaporation of 

single zircons from felsic metavolcanic samples of the Theespruit Formation (a–c, e–h) 

and intrusive granite (d), Tjakastad and Songimvelo areas, southern Barberton 

Greenstone Belt, South Africa (see Fig. 4 for sample locations): a) Spectrum for 4 

grains from felsic sample BA56, integrated from 444 ratios; b) Spectrum for 3 grains 

from felsic sample BA57, integrated from 320 ratios; c) Spectrum for 4 grains from felsic 

sample BA62, integrated from 363 ratios; d) Spectrum for 3 grains from sheared granite 

dyke sample BA65, integrated from 193 ratios; e) Spectrum for 3 grains from felsic tuff 

sample BA70, integrated from 220 ratios; f) Spectrum for 3 grains from felsic 

agglomerate sample BA71, integrated from 252 ratios; g) Spectrum for 3 grains from 

felsic tuff sample BA73, integrated from 220 ratios; h) Spectrum for 3 grains from felsic 

tuff sample BA84, integrated from 221 ratios. 

129


Fig. 10. Histograms showing distribution of Pb isotope ratios derived from evaporation 

of single zircons from felsic metavolcanic samples of the Theespruit Formation 

Tjakastad and Songimvelo areas, southern Barberton Greenstone Belt, South Africa (see 

Fig. 4 for sample locations): a) Spectrum for 6 grains from felsic tuff sample BA85, 

integrated from 581 ratios; b) Spectrum for 3 grains from feldspar porphyry sample BA 

105, integrated from 302 ratios; c) Spectrum for 4 grains from felsic tuff sample BA 107, 

integrated from 382 ratios; d) Spectrum for 4 grains from felsic agglomerate sample BA 

108, integrated from 549 ratios; e) Spectrum for 5 grains from felsic tuff sample BA 109, 

integrated from 461 ratios; f) Spectrum for 3 grains from felsic metavolcanic schist 

sample BA 110, integrated from 256 ratios. 

techniques were used for zircons from one sample. Sample preparation 

for the evaporation technique followed methods described in 

Kröner et al. (2003). In addition, we determined the Sm–Nd wholerock 

isotopic composition of selected samples in order to calculate Nd 

model ages and ɛ Nd(t) values. The different analytical procedures are 

summarized in Appendix A. Sample locations and short field 

characteristics are given in Table 1 and Fig. 4. 

Samples BA56, 57, 62, 70, 71, 73, 79, 84, 85, 104, 105, 107, 108, 109 

and 110 represent variably foliated felsic volcanic schist layers of the 

Theespruit Formation in the Tjakastad and Songimvelo areas (Fig. 4). 

We also collected mafic rocks from the base of the succession, 

including BA 66, 69 (massive foliated amphibolite), 74 (pillowed 

basaltic komatiite) and 75 (gabbroic sill). Samples BA 79, 84, 85, 109 

and 110 represent Theespruit felsic tuff or trachytic breccia interlayered 

with chert, along strike to the east of the main Tjakastad 

samples. The zircons in all samples vary in colour from clear to yellowbrown 

and in shape from stubby to long-prismatic. Most are 

idiomorphic, occasionally with slight rounding at their terminations. 

No core-overgrowth relationships were seen under the microscope. 

Between 3 and 6 zircons per sample were evaporated, and the mean 

207 Pb/ 206 Pb ages obtained on 13 samples vary within a narrow range 

between 3522 and 3533 Ma, with very low errors (generally b1 Ma: 

see Table 2 and Figs. 9–10). 

Felsic pyroclastic sample 160932 was collected from the middle of 

the three main felsic volcaniclastic units identified in the Tjakastad area 

(Fig. 4). The horizon is a coarse felsic volcaniclastic rock that contains 

irregular-shaped fragments of fine-grained and feldspar-phyric dacite, 

and rounded cobbles (to 20 cm diameter) of grey-black and white 

layered chert in a fine-grained felsic volcanic matrix (Fig. 5b). The 

reasoning behind dating this sample was that it is from the most 

heterogeneous felsic layer in the region and thus should have the most 

varied detrital zircon population, if derived from a sedimentary 

diamictite with basement orthogneiss clasts, as previously suggested. 

The sample contained abundant zircon grains with a single 

population of dark- to medium-brown, elongate prisms that have 

common bipyramidal terminations and internal growth zoning. The 

recovered zircon grains are commonly cracked and often broken, but 

show no cores or rims. Seven zircon grains were analysed on SHRIMP II 

(Table 3 and Fig. 11a). Grain 1 is grossly discordant and probably lost Pb 

at unspecified periods in the past; it is therefore not considered for age 

assessment. Grains 2–7 are variably discordant but are well aligned 

along a chord (MSWD=0.0025) with an upper Concordia intercept at 

3539±16 Ma (2-sigma) and a lower intercept at 0±450 Ma. The 

weighted mean 207 Pb/ 206 Pb age of the six grains is 3539±2 Ma. We 

consider this to reflect the time of eruption of the felsic pyroclastic rock. 

TIMS analysis of five single zircons from the same sample (Table 4) 

yielded three grossly discordant points (Z2, Z4, Z5) which are not 

further considered here, one near-concordant point (Z3) and one 

moderately discordant point (Z1) (Fig.11b). A chord through grains Z3 

and Z1 suggests recent Pb-loss (Fig. 11b), and we therefore consider 

both analyses meaningful. The mean 207 Pb/ 206 Pb age for both Z3 and 

Z1 is 3534±4 Ma, which is identical, within error, to the mean 

SHRIMP age. Combining all data points (SHRIMP and TIMS) yields a 

regression with an upper concordia intercept age of 3534±1 Ma. 

In view of the slightly younger age of zircons from felsic 

volcaniclastic rocks analysed by evaporation as compared to the 

SHRIMP and TIMS results, we cannot completely rule out that nonzero 

lead loss has occurred in the younger zircons, although the grains 

from individual samples all show near-identical 207 Pb/ 206 Pb ratios, 

and our oldest evaporation age of 3533 Ma is identical to a mean 

SHRIMP U/Pb zircon age of 3531±10 Ma reported by Armstrong et al. 

(1990). We therefore adopt a mean age of c. 3530 Ma for the 14 

samples of Theespruit felsic volcanic rocks analysed from the 

Tjakastad–Songimvelo area for the purpose of calculating Nd initial 

isotopic ratios and T DM model ages. It is important to note that no ca. 

3450 Ma zircons were identified during any of these analyses. 

Evaporation of three idiomorphic zircon grains from granite dyke 

sample BA 65 yielded identical Pb isotopic ratios which combine to a 

mean 207 Pb/ 206 Pb age of 3408±0.3 Ma (Table 2, Fig. 9d). We consider 

this age to reflect the time of dyke emplacement, and the dyke may 

reflect either a late, K-feldspar-rich phase of the Theespruit Pluton, or

Table 3 

SHRIMP II analytical data for spot analyses of magmatic zircons from felsic volcaniclastic sample 160932. 

Sample no. U ppm Th ppm Th/U 206 Pb/ 204 Pb 

208 206 

Pb/ Pb 

207 206 

Pb/ Pb 

206 238 

Pb/ U 

207 235 

Pb/ U 206/238 age ±1σ 207/235 age ±1σ 207/206 age ±1σ 

BA160932-1 1459 1871 1.28 2969 0.4817±10 0.2110±4 0.2358±28 6.860 ±85 1365±15 2094 ±11 2914±3 

BA160932-2 210 140 0.67 22,042 0.1758±8 0.3137±7 0.6851±85 29.636 ±381 3364±32 3475 ±13 3539±3 

BA160932-3 573 412 0.72 72,961 0.1830±5 0.3135±4 0.6230 ±76 26.930 ±334 3122±30 3381±12 3538±2 

BA160932-4 268 159 0.59 24,413 0.1468±7 0.3135±6 0.6725 ±83 29.068±370 3316±32 3456 ±12 3538±3 

BA160932-5 87 42 0.48 86,207 0.1273±11 0.3135±11 0.7122 ±93 30.786 ±428 3467±35 3512±14 3538±5 

BA160932-6 527 322 0.61 26,658 0.1517+5 0.3137±5 0.5754±70 24.886±310 2930±29 3304 ±12 3539±2 

BA160932-7 187 151 0.81 39,543 0.2123 ±10 0.3137±8 0.7109 ±90 30.750 ±404 3462±34 3511±13 3539±4 

part of a phase of granite and aplite emplacement found in all granitic 

rocks of the southern Barberton Mountain Land. Importantly, the main 

phase of deformation in the dyke and the Theespruit Formation must 

be post-3408 Ma. 

Whole-rock samples analysed for Sm and Nd isotopes (Table 5) 

were not, in all cases, the same as those used for zircon dating, since 

some rocks did not contain large or sufficient zircons for evaporation. 

In addition to felsic volcanic rocks of the Theespruit Formation, we 

also measured isotopic values for mafic/ultramafic samples from the 

Theespruit Formation, near outcrops of well-preserved pillowed 

basalt that are interlayered with the felsic volcanic rocks. 

The ɛNd(t) values for Theespruit felsic rocks (Table 5 and other data in 

Hamilton et al., 1979 and Kröner et al., 1996) show a moderate scatter 

from −1.4 to + 1.3 (Fig. 12a), suggesting heterogeneous crustal sources, 

including juvenile mantle-derived material, as well as recycling of older 

crust. Although we cannot preclude disturbance of the Sm–Nd system in 

these samples, their high Sm and Nd concentrations probably inhibited 

drastic modifications resulting from fluid/rock interaction (e.g., Gruau 

et al.,1996). A negative ɛNd value for the pillowed metakomatiite sample 

BA 74 may be explained by sample alteration and possibly melt 

contamination by older crustal material. 

The somewhat lower ɛ Nd(t) in the felsic samples than in the 

associated komatiites suggest that these rocks were not comagmatic, 

and that the felsic rocks were derived from a crustal source that was 

ultimately formed from the same mantle as the associated komatiites. 

This is consistent with previous Nd isotopic data for Theespruit felsic 

volcanics from the Steynsdorp area farther east that also suggested 

involvement of older continental crust (Kröner et al., 1996). 

A regression analysis of three komatiitic samples (excluding BA 74, 

which has negative ɛ Nd-values of −0.4 at 3.53 Ga that we infer is due 

to melt contamination by older crust, or alteration of the Sm–Nd 

system by hydrothermal activity that affected the pillow basalts; e.g., 

de Wit et al., 1982; Duchač and Hanor, 1987), yields an age of 3.51± 

0.07 Ga (2σ; MSWD=0.02) (Fig. 12b). Enlarging the MSWD to a value 

of 1 would reduce the error to c. 10 Ma. This age is identical to the 

emplacement age of the Steynsdorp Pluton (3510 Ma; Kröner et al., 

1996) and not much different from the mean zircon age of c. 3530 Ma 

for the Theespruit Formation obtained herein. This agreement 

suggests that the Sm–Nd system in these rocks behaved as a closed 

system during their post-eruption history. We therefore interpret 

the initial ɛNd-values of c. +0.5±1.4 (2σ) for the mantle-derived 

komatiites as representative of their mantle source. The Theespruit 

mantle value is not much different from that of other units of the 

Onverwacht Group, which yielded average ɛ Nd values of +2.5 

(Lécuyer et al., 1994), +1.9 (Chavagnac, 2004), +1.1 (Lahaye et al., 

1995), and +1.1 (Hamilton et al., 1979), with a total range in values 

from +3 to 0 (Fig. 12a). 

A regression line calculated for the Theespruit mafic and associated 

felsic metavolcanic rocks yields an errochron corresponding to an age 

of 3.63±0.13 Ga (2σ, n=9, MSWD=16; initial ɛNd-value=+1.1± 

2.1, 2σ). Addition of four samples of the Theespruit Formation from 

the Steynsdorp area (Kröner et al., 1996) produces an errorchron age 

of 3.60±0.13 Ga (2σ; MSWD=21, initial ɛ Nd-value=+0.8±2, 2σ) 

(Fig. 12b). The near-magmatic age derived from the composite 


regression line indicates that both the crustally-derived felsic rocks 

and mantle-derived komatiites were derived from overall isotopically 

similar sources. The inclusion of the felsite data in addition to the 

komatiitic samples in the regression analysis results in a rotation of 

the regression line to a higher age and increase in data scatter. The 

scatter of the felsite data may reflect open-system behaviour of the 

Sm–Nd system in the felsic rocks or their crustal protoliths. Open- 

Fig. 11. a) Concordia diagram showing SHRIMP analyses of zircons from felsic tuff 

sample 160932 (Table 3). Data boxes for each analysis are defined by standard errors in 

207 Pb/ 235 U, 206 Pb/ 238 U and 207 Pb/ 206 Pb; b) Concordia diagram showing TIMS analyses 

of zircons from sample 160932 (Table 4): recent Pb-loss line passes through zircons Z1 

and Z3, with an upper intercept age of 3534±4 Ma. Dated sample is shown in Fig. 5b 

from locality shown in Fig. 4. 

131


Table 4 

U–Pb TIMS data for sample 160932. 

Fraction Concentration Measured Corrected Atomic ratios a 

Ages [Ma] 

Weight U Pb R 

Common Pb T 206 Pb/ 

208 

Pb/ 

206 238 

Pb/ U 

207 235 

Pb/ U 

207 206 

Pb/ Pb 

206 

Pb/ 

207 

Pb/ 

207 

Pb/ 

(mg) (ppm) (pg) 

204 

Pb 

206 

Pb 

238 

U 

235 

U 

206 

Pb 

Z1 brn euh frag 0.001 236 216.7 2 8357 0.1994 0.70423 (156) 30.3641 (698) 0.31271 (26) 3437 3499 3534 

Z2 sub prsm 0.001 966 411.1 22 1284 0.2425 0.34605 (72) 8.8451 (176) 0.18538 (26) 1916 2322 2702 

Z3 sub prsm 0.010 17 16.6 4 1905 0.2640 0.72643 (208) 31.3197 (884) 0.31270 (38) 3520 3529 3534 

Z4 prsm 0.001 267 164.4 7 1600 0.2227 0.48660 (134) 16.6645 (460) 0.24838 (32) 2556 2916 3174 

Z5 prsm 0.001 284 171.2 29 331 0.2628 0.46872 (104) 14.7999 (364) 0.22901 (28) 2478 2802 3045 

Zircon fractions Z1–Z3 are of single small grains that were well abraded (Krogh, 1982). Pb R refers to radiogenic Pb; Common Pb T refers to total common Pb. 

Abbreviations: brn = brown; euh = euhedral; frag = fragment; prsm = prism; sub = subhedral. 

a 

Ratios corrected for fractionation, laboratory Pb (1 pg) and U (0.25 pg) blank. Initial common Pb calculated using Pb isotopic compositions of Stacey and Kramers (1975). Twosigma 

uncertainties on isotopic ratios, calculated with a modified unpublished error propagation program written by L. Heaman, are reported after the ratios (in brackets) and refer to 

the final digits. 

system behaviour of the Sm–Nd system has been reported in many 

studies (Moorbath et al., 1997 and references therein; Gruau et al., 

1996), and it may well be that fluid/rock interaction during 

metamorphism of the Theespruit Formation (see above) has affected 

the Sm–Nd system. However, any possible secondary disturbance of 

the Sm–Nd system had no drastic effects on the samples, as shown by 

the reasonable agreement between the whole-rock Sm–Nd ages and 

U–Pb zircon data. 

6. Discussion 

Remapping and dating of the Theespruit Formation and associated 

granitic rocks in the Tjakastad area indicates significant differences 

from previous tectonic interpretations of this area (de Wit et al., 1983). 

In particular, our data question the validity of the thrust-accretion 

model because: 

1. No tectonic slivers or clasts of basement orthogneiss were found in 

the Tjakastad area. Those previously interpreted as such by de Wit 

et al. (1983) were found to consist either of undeformed quartzfeldspar 

porphyry with igneous flow banding, deformed metasediment, 

or clasts of cherty metasediment (metatuff?) within felsic 

agglomerate. Furthermore, there is no evidence to support the 

presence of an unconformity between “basement orthogneiss” and 

“sedimentary diamictite” at locality D in Fig. 3. 

2. Rocks within the Theespruit Formation and KSZ that were 

interpreted as sedimentary diamictite and intrusion breccia by de 

Wit et al. (1983) are re-interpreted here to be felsic volcaniclastic 

rocks that form part of a series of felsic volcaniclastic horizons 

interlayered with mafic schists of indeterminate intrusive or 

extrusive origin. The unique compositional and textural features 

of each felsic horizon indicate that they are distinct from one 

another and thus do not represent slivers of a thrust-duplicated 

unit, as previously suggested (de Wit et al., 1983, 1992). 

3. The presence of consistent top-to-the-northeast bedding indicators 

across the Theespruit Formation, in combination with the results of 

(1) and (2) above, strongly favour a coherent, though strained (see 

point 4, below), stratigraphic succession. The presence of local tight 

folds and of reversals of bedding in the KSZ (de Wit et al., 1983) is 

not considered to be a significant feature, as no large-scale folds of 

mappable units have been found. 

4. Changes in the intensity and orientation of structures across the 

contact between mafic volcanic rocks and the first felsic horizon in 

the Theespruit Formation may readily be explained as an example of 

strain partitioning between the competent pillowed komatiitic unit 

and the well-layered, less competent, and more fully transposed felsic 

tuffaceous rocks (cf. figure 26.17 and related text of Ramsay and 

Huber, 1987; Hanmer and Passchier, 1991 [pp. 22–24]). Northeast of 

this contact, the strain discontinuously increases towards the KSZ, in 

concert with more steeply-plunging and more northerly-trending 

lineations. These data suggest that the southern contact of the 

stratigraphically lowest felsic schist horizon represents the first major 

increase in strain approaching the KSZ and that the abrupt change in 

lineation orientations across this contact reflects varying degrees of 

transposition within a single shear system. Significantly, strain is 

highly variable within the KSZ, and largely dependent on lithology, 

with the highest strain localised in the softest (talc-carbonate) rocks, 

and areas of low strain localised in some pillowed basalts and felsic 

volcanic rocks (Fig. 4). 

5. The distribution of strain, orientation of fabric elements, and 

kinematic data of discrete faults and shear zones within the 

Tjakastad area indicate a single set of penetrative fabric elements 

developed during greenstone belt-down, granitoid-up extensional 

deformation. No evidence of thrust kinematics, fold interference, or 

of multiple overprinting sets of structures was observed, and the 

period of earlier deformation within orthogneisses as interpreted 

by de Wit et al. (1983) has not been substantiated. Compilation of 

available data show that the extensional deformation occurred at 

ca. 3230 Ma. The KSZ is here interpreted as an extensional 

detachment zone, similar to that documented along strike to the 

west by Kisters et al. (2003) (see Fig. 1). As discussed below, we 

suggest that the most likely origin of the KSZ is as a dislocation 

plane to tight, upright, disharmonic folding between deep crustal 

levels dominated by the Stolzburg plutonic suite, and high level 

rocks of the rest of the Barberton Greenstone Belt. 

Table 5 

Sm–Nd isotopic data of metavolcanic rocks from the Theespruit Formation. 

Sample Age 

(Ga) 

Sm Nd 

147 Sm/ 

144 Nd 

143 144 

Nd/ Nd 

(meas.) 

ɛNd(t) TDM 

Felsites 

Ba 57 ∼3.53 4.06 26.79 0.09157 0.510113 −1.4 3.69 

Ba 70 ∼3.53 7.03 44.50 0.09542 0.510254 −0.4 3.62 

Ba 79 ∼3.53 6.35 35.85 0.1071 0.510551 0.1 3.60 

Ba 85 ∼3.53 5.84 34.75 0.1016 0.510457 0.8 3.54 

Ba 104 ∼3.53 7.29 46.86 0.09407 0.510223 −0.4 3.62 

Ba 108 A ∼3.53 1.29 7.093 0.1102 0.510654 0.7 3.55 

Komatiites 

Ba 66 ∼3.53 4.55 20.07 0.1369 0.511278 0.7 3.60 

Ba 69 ∼3.53 3.50 13.74 0.1539 0.511671 0.6 ⁎ 

Ba 74 ∼3.53 2.65 8.740 0.1833 0.512302 −0.4 ⁎ 

repeat ∼3.53 2.61 8.682 0.1819 0.512274 −0.4 ⁎ 

Ba 75 ∼3.53 1.40 4.287 0.1975 0.512685 0.6 ⁎ 

Standards 

BCR-1 

6.59 28.84 0.1382 0.512633 −0.1 

(n=3; 2σ) ±0.02 ±0.08 

±6 

BHVO-1 6.15 24.84 0.1496 0.512966 +6.4 

6.15 24.85 0.1495 0.512978 +6.6 

La Jolla 

(n=6; 2σ) 

0.511853 

±7 

−15.3 

Within-run error for 143 Nd/ 144 Nd b1.3× 10 −5 (2σ mean), external precision ∼1.3× 10 −5 . 

⁎ Sm/Nd-ratio too high for calculation of meaningful model age. Input errors (2σ) for age 

calculations: 147 Sm/ 144 Nd=0.3%, 143 Nd/ 144 Nd=0.003%.

Fig. 12. a) Nd isotope evolution diagram for selected Archaean rocks and the felsic and 

mafic samples from the Theespruit Formation (diamonds); the average εNd value derived 

from a whole-rock regression line of all Theespruit samples is +1.1±2.1 (2σ). Other data 

for metavolcanic rocks from the Onverwacht Group (large circles represent average ε Nd 

values, bars total range in ɛ Nd values): 1, Chavagnac (2004, calculated for 3.45 Ga); 2, 

Hamilton et al. (1979, data derived from a 3.54 Ga whole-rock isochron); 3, Lahaye et al. 

(1995, recalculated for 3.49 Ga). Solid diamond = average ε Nd value for Isua (Caro et al., 

2006). Other data points represent average values compiled from the literature and cited 

by Caro et al. (2006). Solid squares = North Atlantic Craton, Scotland; cross = North 

Atlantic Craton, Labrador; solid triangle = Qianan region (Hebai province), China (Xuan et 

al., 1986); solid circles = Yilgarn Craton, Australia; triangles = Superior Craton, Canada; 

squares = Indian Craton; open circles = Kaapvaal Craton, South Africa. The Nd-evolution 

line for differentiation at 4.4 Ga (stippled line) is from Caro et al. (2006), the other line is 

from DePaolo (1981).b)Sm–Nd isochron diagram for komatiite and felsite samples of the 

Theespruit Formation. Symbols: solid triangles = komatiites from the Tjakastad area; open 

triangles = felsic volcanics from the Tjakastad area; solid circles = felsic volcanics from the 

Steynsdorp area (data in Kröner et al., 1996). 

6. The petrographic evidence of kyanite overprinting andalusite 

presented here is interpreted as a preserved remnant of the 

prograde P–T conditions and indicative of near isothermal burial 

under moderate temperature conditions. When combined with 

previous evidence that metamorphism was syn-kinematic with 

regional extension and decreased in temperature away from the 

Theespruit Pluton (Van Vuren and Cloete, 1995) under low 

geothermal gradients (e.g., Dziggel et al., 2002; Diener et al., 

2006), this new observation suggests that metamorphism in the 

Theespruit area was not the same as that in more recent 

accretionary orogenic belts (Cloete, 1993, 1999), where inverted 

metamorphic gradients and clockwise P–T paths are commonly 

developed (e.g. Brown, 2006). Instead, the metamorphic data, 

when combined with the structural data of weakly deformed cores 

to domical granitoids and strong, vertically lineated, greenstone 

septae between domical granitoids (e.g. Fig. 7a–c), are more 

consistent with episodes of partial convective overturn of green- 


stone-over-granitoid layered crust, as initially modelled by Mareschal 

and West (1980) and discussed by many others (e.g. Collins 

et al., 1998; Collins and Van Kranendonk, 1999; Sandiford et al., 

2004; Van Kranendonk et al., 2004; Bodorkos and Sandiford, 2006). 

7. The geometry of weakly strained crests and strongly sheared, diplineated 

margins to the TTG domes to the south of the Theespruit 

Formation (e.g., Fig. 7a–c) is consistent with partial convective 

overturn of a thick, overlying greenstone succession and underlying, 

more buoyant and partially melted TTG, as originally proposed by 

Viljoen and Viljoen (1969), and subsequently supported by Anhaeusser 

(1984) and Kisters and Anhaeusser (1995), possibly within an 

overall extensional regime (Kisters et al., 2003). This structural 

geometry developed at between c. 3237 and 3216 Ma, which is the 

age range of metamorphic zircons and titanite, respectively, from the 

Stolzburg and Doornhoek Plutons, and contemporaneous with the 

age of the post-tectonic Dalmein Pluton (Kamo and Davis, 1994). 

6.1. Significance of the available geochronology 

Our new interpretation for the Theespruit–Komati Formations 

contact zone contrasts markedly with previous tectonic models based 

on geochronological data and this requires some discussion. 

Previous tectonic models based on zircon dating suggested that 

thrusting occurred at ca. 3230 Ma (Kamo and Davis, 1994). However, 

this is negated by the fact that both the higher-grade Theespruit and 

Sandspruit Formations structurally beneath the KSZ, and the lowergrade 

greenstones above the KSZ (Komati Formation and younger 

formations), experienced a common felsic magmatic event at ca. 

3450 Ma (Kröner et al., 1991; Kamo and Davis, 1994), and thus were in 

contact at, or before, this time. If thrusting did occur, then it must have 

been at, or before, ca. 3450 Ma. However, several lines of evidence 

preclude such an interpretation. First, only one phase of deformation 

is recognised in the area, and that is extensional, and/or related to 

partial convective overturn of greenstones and granitoids. Second, an 

early thrust scenario is improbable because of the conformable 

relations between the Komati Formation and overlying rocks of the 

Hooggenoeg (3445 Ma), Kromberg (≤3416–3334 Ma), and Mendon 

(3298 Ma) Formations, and Fig Tree (b3260 Ma) Group (Fig. 2: Byerly 

et al., 1996; Lowe and Byerly, 2007), which shows that the greenstone 

sequences were deposited progressively from at least 3490 to 

3260 Ma. Third, the 3408 Ma age of the granite dyke that we dated 

from the Tjakastad area (sample BA 65) is significant because it cuts 

bedding in the volcanic rocks and is deformed, indicating deformation 

occurred after its emplacement. 

The main deformation of the belt at ca. 3230 Ma was extensional, as 

evidenced by rodding lineations, extensional faults, and kinematic data 

from within the KSZ (this paper and Kisters et al., 2003). The 3418±8 Ma 

titanite age cited by Dziggel et al. (2001) is most likely related to late stage 

felsic magmatism in the area, as indicated by the 3408±2 Ma Pb–Pb 

zirconageforthefoliatedgranitedykesampleBA65intheTjakastadarea. 

The zircon ages of Armstrong et al. (1990) and Kamo and Davis 

(1994) indicate that the unit previously interpreted as a basement 

gneiss sliver, but which we re-interpret to represent a felsic 

volcaniclastic rock within the Theespruit Formation, was erupted at 

3538 Ma. This is the same age as our felsic schist sample 160932 

(Tables 3, 4; Fig. 11) and the ages of 13 samples of felsic schist of the 

Theespruit Formation located along strike in the map area and to the 

east around the Steynsdorp Anticline (Kröner et al., 1996), as well as 

for the older population of zircons in the “diamictite” (herein reinterpreted 

as felsic agglomerate) from the Tjakastad area (3531± 

10 Ma; Armstrong et al., 1990). Based on this remarkable similarity in 

age for the same lithostratigraphic unit now dated for 18 samples (this 

paper and Kröner et al., 1996), it is clear that 3530 Ma represents the 

true age of felsic volcanism in the Theespruit Formation. 

The remaining question then becomes what is the origin and 

significance of the 3453±6 Ma population of zircons within the 

133


Theespruit Formation agglomerate dated by Armstrong et al. (1990)? 

The description of zircon morphology by Armstrong et al. (1990), 

combined with subsequent information in Kamo and Davis (1994), 

allows for a reasonable alternative interpretation. 

First, it must be recognised that the dated rock was metamorphosed 

under amphibolite-facies conditions, at which grade metamorphic 

recrystallisation of zircon and growth of new zircon rims on 

older cores are known to occur (e.g., Hoskin and Schaltegger, 2003). 

The observation by Armstrong et al. (1990) of core–rim structures in 

zircons from their dated sample suggests that the younger population 

of zircons may have grown during the contact metamorphic event 

associated with the emplacement of the ca. 3450 Ma Stolzburg 

plutonic suite (e.g., Van Vuren and Cloete, 1995) and associated felsic 

intrusions that span the Komati schist zone. Supporting evidence for 

such an event was presented by Kamo and Davis (1994) who 

identified major felsic igneous activity in the Barberton region at 

3476–3440 Ma based on the following ages (listed from structural 

base to top): 3460+5/−4 Ma zircon in the older phase of the 

Stolzburg Pluton, which is located immediately west of, and connected 

to, the 3443+4/−3 Ma Theespruit Pluton (Fig. 1); 3458±2 Ma 

titanite in a 3476±10 Ma felsic dyke cutting the Komati Formation; 

3457±3 Ma zircon from a felsic sill emplaced into the base of the 

Komati Formation; and ca. 3445 ± 5 Ma from the Hooggenoeg 

Formation (de Wit et al., 1987b), which conformably overlies, and is 

genetically related to, the Komati Formation (Lowe and Byerly, 2007). 

We therefore re-interpret the geochronological information from 

the Theespruit Formation to indicate that a bimodal, mafic and felsic 

volcanic succession was erupted at ca. 3530 Ma, and then intruded by 

the Steynsdorp Pluton at 3510 Ma. Following eruption of the Komati 

Formation at 3490–3480 Ma, the emplacement of sheets of felsic 

plutonic rocks into the volcanic succession at between 3476 and 

3443 Ma caused contact metamorphism and the growth of new 

zircons and titanite in the host rocks, particularly lower in the section 

where the volumetrically significant Stolzburg plutonic suite was 

emplaced and resulted in amphibolite-facies contact-style 

metamorphism. 

6.2. Origin of the Komati schist zone and regional folding 

The regional geology of the southwestern Barberton Greenstone Belt 

(BGB)isdominatedbyasetoffoldsthatplungeinwardstowardsthecore 

of the belt, and are defined by lobate, antiformal granitoids and cuspate, 

generally synclinal greenstones (Fig. 13). The strike of fold axial traces 

rotates through 130° around the southwestern corner of the BGB, from 

almost due north-striking in the southeast, to northeast-striking in the 

southwest, to southeast-striking in the northwest (Fig. 13). The lobe-cusp 

fold geometry of this area indicates contemporaneous deformation of 

competent and incompetent lithologies, respectively. The age of 

this folding is tightly constrained by the fact that folds involve the 

3229±5 Ma Kaap Valley Tonalite, but are cut by the 3216+2/−1Ma 

Dalmein Pluton (Kamo and Davis, 1994: Fig. 1). 

Within this broad fold context, the high-strain Komati schist zone 

is located within the hinge of a regional, upright anticline, along the 

boundary between a thick, upper succession of low grade, low strain, 

volcanic flow units of the post-Theespruit Formation rocks of the 

Onverwacht Group (≤ca. 3.49 Ga) and a thinner, lower succession of 

amphibolite-facies, older greenstones (c. 3.53 Ga Theespruit and 

Sandspruit Formations) and ca. 3.45 Ga trondhjemitic intrusions. 

Indeed, the KSZ is located directly along the amphibolite-greenschist 

isograd (Fig. 4), where ultramafic rocks and greenstones are 

transformed into soft talc-carbonate and chlorite schists, respectively, 

mineral assemblages perfectly suited to accommodate large degrees 

of shear strain. Kinematic data indicate extensional, top to the eastnortheast 

shearing along this zone. Kisters et al. (2003) documented 

an eastward-dipping, top-to-the-east-northeast extensional detachment 

plane at the same structural level as the KSZ along strike to the 

west (Figs. 1, 13), which can be considered the continuation of the 

same zone. A similar high strain zone occurs at the same structural 

level along strike to the east, separating the Steynsdorp Pluton and its 

surrounding envelope of Theespruit Formation rocks from lowergrade, 

younger rocks of the upper Onverwacht Group in another hinge 

of a tight regional anticline (Fig. 1). Given that this zone separates 

younger rocks in the north from older, higher-grade rocks in the 

southwest, it is most likely that the eastern zone is also extensional; 

thus, we consider all of these zones to be a continuous structural 

dislocation zone across the southwestern Barberton Greenstone Belt, 

intimately related to fold development. 

Significantly, these dislocation zones developed at the same time 

as folding of the low-grade greenstones of the post-Theespruit 

Formation units of the Onverwacht Group, and also at the same time 

as the dome-and-keel geometry and peak metamorphic conditions 

were being developed in the higher grade rocks (see 2.3; Metamorphism: 

Dziggel et al., 2001, 2002; Diener et al., 2006). 

Studies of fold mechanics in buckled multilayers with units of 

different competencies show that high-strain dislocation zones 

commonly develop along unit interfaces and that folding may be 

accommodated by a variety of processes, including disharmonic 

folding if units are of significantly different thicknesses (e.g., Ramsay 

and Huber, 1987). Considering the contemporaneity of fold formation 

and shear zone development, the most likely origin of the KSZ and its 

equivalents along strike to the east and west, is as an extensional 

detachment surface during regional folding. That folds in the upper 

Onverwacht Group are tighter than folds of the KSZ and its equivalents 

along strike, but that folding also affected the shear zones themselves, 

suggests that folding was contemporaneous with shearing. As 

discussed below, the fact that the rocks above the detachment surface 

display one style (upright, large amplitude and large wavelength 

folds) and those below display another style (dome-and-keel 

geometry: see below) suggests disharmonic folding and/or a different 

response at different structural levels to similar overall driving forces. 

In the upper-crustal, lower-grade rocks, the inward-plunging 

nature of the folds around the southwestern BGB into the core of 

the greenstone belt, together with the greenstone-down kinematics 

across the shear zones, strongly suggest that deformation was driven 

by greenstone sinking, as previously suggested for this area (Viljoen 

and Viljoen, 1969; Anhaeusser, 1984). Greenstone sinking was 

accompanied by the exhumation of deeper level rocks beneath the 

detachment zones, as indicated by the presence of higher pressure 

metamorphic mineral assemblages in these rocks. Exhumation along 

the Komati schist zone may have excised a considerable thickness of 

stratigraphy between the Theespruit and Komati Formations, 

although this is unconstrained. 

In the higher-grade rocks, the presence of vertically-rodded Ltectonites 

in strongly deformed greenstone keels between low-strain, 

domed granitoid cores is consistent with deformation driven by 

greenstone sinking. This is in contrast to, and different from, 

deformation generated as a result of granite doming (diapirism), 

where greenstones in the roof zones of diapric granitoids should be 

extensively flattened (e.g. Ramberg, 1967): however, this is clearly not 

the case, as shown in Fig. 7a. Greenstone sinking is confirmed by the 

occurrence of high-P metamorphic mineral assemblages of 9 kbar, 

700 °C in highly transposed greenstone septae between domes 

(Fig. 13: Van Vuren and Cloete, 1995). The low geothermal gradients 

indicated by metamorphic studies in this area are also consistent with 

greenstone sinking (cf. Collins and Van Kranendonk, 1999). 

Some aspects of the geology of this area have similarities with 

metamorphic core complexes, as suggested by Kisters et al. (2003). 

However, the fact that folds and metamorphic mineral elongation and 

stretching lineations point inwards towards the core of BGB contrasts 

with the typically unidirectional trend of linear fabric elements within 

metamorphic core complexes, as does the dome-and-keel geometry of 

folding within the higher-grade footwall in the Tjakastad area. Other

differences between the geology of the Barberton Mountain Land and 

Phanerozoic metamorphic core complexes include the lack of 

evidence in the former for thinning of the lithosphere during 

extension, and for a prior episode of structural thickening of the 

lithosphere through orogenesis (see also Hickman and Van Kranendonk, 

2004; Van Kranendonk et al., 2004). 

6.3. Regional orogeny 

Of critical importance relating to interpretations of the tectonic 

evolution of the BGB, is the contrasting evidence for contemporaneous 

extensional deformation (KSZ and associated zones along strike) and 

compressional deformation at c. 3230 Ma. Extension is indicated by 

the KSZ and associated zones along strike, and by the dome-and-keel 

architecture of c. 3.45 Ga trondhjemitic rocks along the southwestern 

margin of the belt. Compressional deformation is indicated by the 

presence of upright folds around the southwestern BGB, synsedimentary 

contractional folds (including recumbent isoclinal 

folds) in the Moodies Group (e.g. Lamb, 1987), and horizontaltectonics 

style structures in the Fig Tree Group (e.g., de Wit, 1982). 

Significantly, extensional structures are restricted to the higher-grade, 

lower structural and stratigraphic parts of the BGB below the KSZ 

detachment, whereas compressional structures are restricted to 

higher structural and stratigraphic levels within the lower-grade 

core of the BGB, above the KSZ detachment. Whereas this contempor- 


Fig. 13. Simplified litho-structural map of the southwestern Barberton Greenstone Belt and surrounding area, showing the position of high strain shear zones within the regional 

lobe-cusp fold pattern. Lower greenstones = Sandspruit and Theespruit Formations; Upper greenstones = main part of Onverwacht Group (Komati, Hooggenoeg, Kromberg and 

Mendon Formations) and Fig Tree and Moodies Groups. DP = Doornhoek Pluton; KSZ = Komati schist zone; SP = Stolzburg Pluton; TP = Theespruit Pluton. Metamorphic pressure– 

temperature conditions also shown (see text for references). Towns: B = Barberton, L = Lochiel, T = Tjakastad. 

aneity of contrasting styles does occur in Phanerozoic plate-collisional 

orogens, extension in these terranes follows significant crustal 

thickening (crustal doubling) through thrusting, caused by the 

collision of tectonic plates (e.g. Searle, 2007). 

We contend that our new data, combined with the evidence of an 

upward-younging, thick greenstone stratigraphy, precludes an origin 

of the BGB through subduction–accretion, as widely supposed (e.g. de 

Wit et al., 1992; Lowe, 1994; Stevens et al., 2002; Moyen et al., 2006). 

Rather, we suggest that the available data supports an origin of the 

BGB through partial convective overturn of a dense upper crust and 

more buoyant middle crust at c. 3230 Ma, similar to models previously 

proposed by Viljoen and Viljoen (1969) and Anhaeusser (1984). 

We envisage a two-stage development of ancient Kaapvaal Craton 

crust. 

1) Long-lived, predominantly mantle-derived magmatism from 

3530–3416 Ma, which resulted in the construction of a thick greenstone 

pile, based on a substrate of older continental crust, perhaps including 

the ≤c. 3.64 Ga Ancient Gneiss Complex (AGC) in Swaziland, on the 

southeastern side of the BGB (Kröner, 2007). This was further thickened 

by the emplacement of c. 3470–3437 Ma TTG melts, for which 

geochemical data indicates melting under a range of pressures from a 

dominantly infracrustal basaltic source (Moyen et al., 2007): this implies 

that the Kaapvaal crust was already thick by this time. Emplacement of 

TTG into the mid-crust caused low-pressure metamorphism of adjacent 

greenstones (e.g. andalusite in the present study area). 

135


2) Subsequent addition of mantle-derived melts at 3334–3298 Ma 

(Kromberg and Mendon Formations) resulted in magmatic overthickening 

of the crust and development of a gravitational instability 

represented by a N10 km thick, largely komatiitic volcanic pile on top of 

granitic (TTG and AGC) middle crust. Modelling studies show that this 

type of gravitationally unstable Palaeoarchaean crustal architecture, 

when combined with the input of conductive heat from the mantle 

melts and radiogenic heat in TTGs buried to mid-crustal levels by the 

associated eruptives, results in partial melting of the TTG middle crust, 

which, in turn leads to sinking of greenstones and doming of the 

partially melted TTG during partial convective overturn (Rey et al., 2003; 

Sandiford et al., 2004; Bodorkos and Sandiford, 2006). We suggest that 

this mechanism, rather than subduction–accretion, was responsible for 

the deep burial of greenstones along the flanks of doming granitoids in 

the lower structural parts of the terrain, where high-pressure 

metamorphic assemblages have been recorded (e.g. Dziggel et al., 

2002; Diener et al., 2006). Partial convective overturn was rapid, with 

associated isothermal decompression of the lower structural greenstones 

during extension-related exhumation at ca. 3230 Ma (Kisters 

et al., 2003; Dziggel et al., 2005). Uplift of the granitoid margins to the 

BGB shed detritus into the sinking greenstone core (Fig Tree and 

Moodies Groups), which became deformed as the sediments accumulated. 

Horizontal translation of the greenstones into the zone of sinking 

from adjacent to the rising granitoids was effected across mid-crustal 

detachment zones such as the KSZ and caused horizontal-style 

deformation of the upper greenstones (including recumbent isoclinal 

folds: e.g., Ramberg, 1967). 

The model of partial convective overturn for the development of 

the BGB outlined above should be considered as a working, testable 

hypothesis and that it is not necessarily exclusive of subduction/ 

accretion tectonics in the development of the craton as a whole, nor of 

Archaean plate tectonics in general (e.g., Van Kranendonk, 2004, 

2007; Smithies et al., 2005). 

7. Conclusions 

Remapping of the southwestern part of the Barberton Greenstone 

Belt, combined with zircon dating and Sm–Nd data, has revealed several 

important inconsistencies with the previous horizontal tectonic model 

of de Wit et al. (1983, 1992).Mostsignificantly, the kinematics of shear 

deformation in the Theespruit Formation is extensional rather than 

contractional, as previously proposed. Rocks previously interpreted as 

tectonic slivers of basement orthogneiss were found instead to be 

composed of little deformed quartz-feldspar porphyry or sheared felsic 

volcaniclastic metasediment, and those interpreted as sedimentary 

diamictite were found to be felsic agglomerates. Consistent facing 

directions of bedding to the northeast, the recognition of distinct 

stratigraphic variations between several felsic volcanic and volcaniclastic 

horizons across strike, and strain data all combine to suggest that the 

Theespruit Formation in this area is a ca. 3530 Ma old, moderately 

strained, coherent lithostratigraphic succession and not a tectonostratigraphic 

assemblage as proposed by de Wit et al. (1983, 1992). 

We re-interpret a 3458 Ma population of zircons dated from an 

amphibolite-facies Theespruit Formation felsic schist by Armstrong 

et al. (1990) as being of metamorphic, rather than detrital origin, 

based on these authors' observation of core–rim structures in zircons, 

the amphibolite-facies metamorphic grade of the dated rocks, and our 

observations for the widespread emplacement of felsic plutonic rocks 

of that age across the Komati Schist Zone, which we suggest provided 

heat for contact metamorphism of the greenstones. 

In combination with previous work (Byerly et al., 1996; Lowe and 

Byerly, 2007), the data presented here show that the Barberton 

Supergroup is an upward-younging, autochthonous succession 

deposited over ∼300 Ma (3530–3230 Ma). Deformation commenced 

during deposition of the Moodies Group at ca. 3230 Ma, and involved 

partial convective overturn of upper and middle crust. 

Acknowledgements 

MVK thanks Thinus Cloete for help with fieldwork logistics and for an 

introductory fieldtrip through the Barberton Mountain Land in 1998. A.K. 

acknowledges mass spectrometric analytical facilities in the Max-Planck- 

Institut für Chemie in Mainz. Zircon analyses of sample 160932 were 

carried out on the Sensitive High Resolution Ion Microprobe Mass 

Spectrometer (SHRIMP II) operated by a consortium consisting of Curtin 

University of Technology, the Geological Survey of Western Australia and 

the University of Western Australia with the support of the Australian 

Research Council. We thank M.T.D. Wingate for assistance during SHRIMP 

analysis. A.K. also appreciates financial support from the Institute of 

Geoscience Research (TIGeR), Curtin University of Technology, Perth. Paul 

Evins and two anonymous reviewers provided comments that helped 

clarify parts of the manuscript. This paper is published with the 

permission of the Executive Director of the Geological Survey of Western 

Australia and is contribution no. 436 of the Geocycles Cluster of Mainz 

University, funded by the State of Rhineland-Palatinate, Germany. 

Appendix A. Laboratory procedures 

A.1. Single zircon evaporation 

Our laboratory procedures, as well as comparisons with conventional 

and ion-microprobe zircon dating, are detailed in Kröner et al. 

(1991) and Kröner and Hegner (1998). Isotopic measurements were 

carried out on a Finnigan-MAT 261 mass spectrometer at the Max- 

Planck-Institut für Chemie in Mainz. 

The calculated ages and uncertainties are based on the means of all 

ratios evaluated and their 2σ mean errors. Mean ages and errors for several 

zircons from the same sample are presented as weighted means of the 

entire population. During the course of this study we repeatedly analysed 

fragments of large, homogeneous zircon grains from the Palaborwa 

Carbonatite, South Africa. Conventional U–Pb TIMS analyses of six separate 

grain fragments from this sample yielded a 207 Pb/ 206 Pb age of 2052.2± 

0.8 Ma (2σ, W.Todt,unpublisheddata),whereasthemean 207 Pb/ 206 Pb 

ratio for 19 grains, evaporated individually over a period of 12 months, was 

0.126634±0.000027 (2σ error of the population), corresponding to an 

age of 2051.8±0.4 Ma, which is identical to the U–Pb TIMS age, within 

error. The above 2σ error of the population of evaporated zircons is 

considered the best estimate for the reproducibility of our evaporation 

data and corresponds approximately to the 2σ (mean) error reported for 

individual analyses in this study (Table 2). In the case of combined data 

sets, the 2σ (mean) error may become very low, and whenever this error 

was less than the reproducibility of the internal standard, we have used 

the latter value (that is, an assumed 2σ error of 0.000027). 

The analytical data are presented in Table 2, andthe 207 Pb/ 206 Pb 

spectra are shown in histograms that permit visual assessment of the 

data distribution from which the ages are derived. The evaporation 

technique provides only Pb isotopic ratios, and there is no aprioriway to 

determine whether a measured 207 Pb/ 206 Pb ratio reflects a concordant 

age. Thus, all 207 Pb/ 206 Pb ages determined by this method are 

necessarily minimum ages. However, many studies have demonstrated 

that there is a very strong likelihood that these data represent true zircon 

crystallisation ages when (1) the 207 Pb/ 206 Pb ratio does not change with 

increasing temperature of evaporation and/or (2) repeated analyses of 

grains from the same sample at high evaporation temperatures yield the 

same isotopic ratios within error. Comparative studies by evaporation, 

conventional U–Pb dating, and ion-microprobe analysis have shown this 

to be correct (Kröner et al.,1991,1999; Cocherie et al.,1992; Jaeckel et al., 

1997; Karabinos, 1997). 

A.2. SHRIMP II analysis 

U–Pb ion microprobe data for zircons from sample 160932 were 

obtained on the SHRIMP II (B) of the John de Laeter Centre of Mass

Spectrometry at Curtin University, Australia (DeLaeter and Kennedy, 

1998). Clear euhedral zircons some 80 to 150 µm in length were 

handpicked and mounted in epoxy resin together with chips of the 

Perth zircon standard CZ3. For data collection, six scans through the 

critical mass range were made, and details on the analytical procedure 

and data reduction can be found in Compston et al. (1992), Stern 

(1997), Nelson (1997), and Williams (1998). Primary beam intensity 

was about 2.0 nA, and a Köhler aperture of 100 µm diameter was used, 

giving a slightly elliptical spot size of about 30 µm. Sensitivity was 

about 26 cps/ppm/nA Pb on the standard CZ3. Analyses of samples 

and standards were alternated to allow assessment of Pb + /U + 

discrimination. Common-Pb corrections have been applied using the 

204 Pb-correction method, and because of low counts on 204 Pb it was 

assumed that common lead is surface-related (Kinny, 1986), and the 

isotopic composition of Broken Hill lead was used for correction. The 

analytical data are presented in Table 3. Errors on individual analyses 

are given at the 1-sigma level and are based on counting statistics and 

include the uncertainty in the standard U/Pb age (Nelson, 1997). 

Errors for pooled analyses are at the 2-sigma or 95% confidence 

interval. The data are graphically presented on a conventional 

concordia plot (Fig. 11a). 

A.3. TIMS analysis of zircons 

Rock samples were crushed to mineral size under clean conditions 

using a jaw crusher and disc pulveriser and initial heavy mineral 

concentrates were made using a Wilfley table at The University of Texas 

at Austin. Wilfley table concentrates were sieved to remove particles 

larger than 70 mesh and zircon was concentrated using standard heavy 

liquid and magnetic separation techniques. Zircons were characterised 

using a binocular reflected-light microscope, transmitted light petrographic 

microscope (with condenser lens inserted to minimise edge 

refraction) and a scanning cathodoluminescence (CL) imaging system 

on a JEOL 730 scanning electron microscope. 

Single grains of zircon were strongly abraded following Krogh 

(1982), subsequently re-evaluated optically and then washed successively 

in distilled 4 N nitric acid, water and acetone. They were loaded 

dry into TEFLON capsules with a mixed 205 Pb/ 235 U isotopic tracer 

solution and dissolved with HF and HNO 3. Chemical separation of U and 

Pb from zircon using minicolumns (0.055 ml resin volume; after Krogh 

1973) resulted in total Pb and U procedural blanks that are estimated to 

be 1 and 0.25 pg, respectively. Pb and U were loaded together with silica 

gel and phosphoric acid onto an outgassed filament of zone-refined 

rhenium ribbon and analysed on a multi-collector MAT 261 thermal 

ionisation mass spectrometer operating in dynamic mode with all 

masses measured sequentially by the secondary electron multiplier 

(SEM) — ion counting system. Ages were calculated using decay 

constants of Jaffey et al. (1971). Errors on isotopic ratios were calculated 

by propagating uncertainties in measurement of isotopic ratios, 

fractionation and amount of blank with a program written by JNC. 

Results are reported in Table 4 with 2σ errors. 

A.4. Sm–Nd isotopic analysis 

Whole-rock isotopic analyses were carried out at Tübingen and 

Munich Universities using a technique described in Hegner et al. 

(1995) and Hegner and Kröner (2000). The sample powders were 

totally spiked with a 150 Nd– 149 Sm tracer and dissolved initially in an 

open-beaker overnight, followed by dissolution of accessory phases in 

steel jacket-lined TEFLON bombs at 160 °C over 5 days, ensuring 

decomposition of e.g. zircon and garnet. Total procedural blanks were 

b100 pg for Nd and Sm and are not significant for the samples of the 

present study. 

143 Nd/ 144 Nd isotopic ratios were measured using MAT 262 and MAT 

261 mass spectrometers and a dynamic-quadruple mass collection 

mode. Nd and Sm concentrations were calculated from static collection 


routines. 143 Nd/ 144 Nd ratios are normalized to 146 Nd/ 144 Nd=0.7219, 

and Sm isotopic ratios to 147 Sm/ 152 Sm=0.56081. During the course of 

this study the La Jolla Nd standard yielded 143 Nd/ 144 Nd=0.511853±7 

(2σ, n=6) and Ames-Nd, prepared at Munich 0.512139±11. Analyses 

of the rock standards BCR-1 and BHVO-1 are included in Table 5.Withinrun 

precision (2σ mean) for Nd was ≤1.3× 10 −5 ; the external precision 

inferred from a large number of analyses is about 1.3×10 −5 (2σ), 

resulting in an error of ca.±0.3 ɛNd units. 

The Sm–Nd isotopic data are listed in Table 5.TheɛNd(t) values have 

been calculated for an age of 3.53 Ga using present-day Bulk Earth values 

of 143 Nd/ 144 Nd=0.512638 (normalized to 146 Nd/ 144 Nd=0.7219) and 

147 Sm/ 144 Nd=0.1967 (Jacobsen and Wasserburg, 1980; Wasserburg 

et al.,1981). The Nd model ages were calculated using the parameters of 

DePaolo (1981). We interpret the Nd model ages in terms of mean 

crustal residence ages (Arndt and Goldstein, 1987). 

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