Provenance characteristics of the Brumunddal sandstone in the Oslo ...

NORWEGIAN JOURNAL OF GEOLOGY Provenance characteristics of the Brumunddal sandstone in the Oslo Rift 

Provenance characteristics of the Brumunddal sandstone 

in the Oslo Rift derived from U-Pb, Lu-Hf and trace element 

analyses of detrital zircons by laser ablation ICMPS 

Tom Andersen, Ayesha Saeed, Roy H. Gabrielsen & Snorre Olaussen 

Andersen, T., Saeed, A., Gabrielsen, R.H. & Olaussen, S.: Provenance characteristics of the Brumunddal sandstone in the Oslo Rift derived from 

U-Pb, Lu-Hf and trace element analyses of detrital zircons by laser ablation ICMPS. Norwegian Journal of Geology, vol. 91, pp 1-19. Trondheim 

2011. ISSN 029-196X. 

The 800 m thick Brumunddal sandstone is partly an eolian, partly a fluvial sandstone deposited in a fault-bounded basin in the northern part of the 

Oslo Rift in Permian time.The sandstone is the youngest rift-related deposit in the northern part of the Oslo Rift. Well rounded detrital zircons are 

common accessory mineral grains in the sandstone. U-Pb dating of detrital zircon from a sample of the Brumunddal sandstone by LAM-ICPMS 

gives a range of ages from (rare) late Archaean ages to Permian (283±4 Ma). The age and initial εHf pattern of zircons in the sediment match the 

main rock- forming events in Fennoscandia from Archaean to Phanerozoic time. This kind of diverse provenance was most likely obtained by 

repeated recycling of clastic sediments of Fennoscandian origin, with Silurian, continental sandstones as the most probable direct precursors. Trace 

element distributions show a conspicuous absence of patterns with the high level of U and Th enrichment typical of zircon from granitic rocks. 

This is consistent with a complex transport and redepositional history: High U-Th, metamict zircons were selectively removed by abrasion during 

repeated transport-deposition-erosion cycles. In addition to recycled material, Caledonian syn-orogenic intrusions and Permian intermediate to 

felsic plutonic rocks in the Oslo Rift itself were minor, but still significant sources of detrital zircon. 

Tom Andersen & Roy H. Gabrielsen, Department of Geosciences, University of Oslo, PO Box 1047 Blindern, NO-0316 Oslo, Norway (tom.andersen@geo. 

uio.no). Ayesha Saeed, GEMOC, Macquarie University, NSW-2109, North Ryde, Australia. Snorre Olaussen, The University Centre in Svalbard, P.O. Box 

156, N-9171 Longyearbyen, Norway 

Introduction 

The Brumunddal sandstone (informal name) is the 

uppermost preserved member of the late Paleozoic, riftrelated, 

volcanosedimentary succession in the northern 

part of the Oslo Rift. It was deposited on a substrate of 

rhomb porphyry lava, and is preserved in a ca. 6.5 km by 

2.5 km, downfaulted block, the Narud halfgraben (Fig. 

1). The Brumunddal sandstone consists of red and yellow, 

cross-bedded, eolian dune deposits and fluvial channel 

and floodplain deposits (Rosendal 1929, Larsen et al. 

2008). Its maximum depositional age is constrained by 

a Rb-Sr mineral isochron age of the underlying rhomb 

porphyry lava at 279±9 Ma (Sundvoll et al. 1990). The 

absence of overlying, datable rocks leaves the younger 

limit for the age of deposition unconstrained. Larsen et 

al. (2008) regarded the sandstone as a remnant of graben 

fill that formed during the main rifting stage of the Oslo 

Rift, which is characterized by widespread plateau lavas 

intercalated by continenal deposits. 

Through the U-Pb and Lu-Hf isotope systems and trace 

element distribution patterns, detrital zircon grains 

in a sandstone retain a robust memory of the age, 

petro genetic history and compositional characteristics 

of the magmatic or metamorphic rock(s) in which they 

formed, i.e. of the protosource (e.g. Fedo et al. 2003, 

Kinney & Maas 2003, Belousova et al. 2002, 2006). As 

the youngest pre-Quaternary sedimentary deposit in the 

northern part of the Oslo Rift, this sandstone may have 

sampled a range of Precambrian and Paleozoic protosources 

in the Fennoscandian Shield, the Caledonian 

nappes and within the Oslo Rift itself. However, detrital 

zircons in a sedimentary rock may have been recycled 

through one or more intermediate reservoirs by erosion 

and redeposition of older sedimentary rocks (e.g. Røhr 

et al. 2010). 

In this paper, we present the results of an analytical study 

applying laser ablation inductively coupled mass spectrometry 

(LAM-ICPMS) and electron microprobe analysis 

to detrital zircon in the Brumunddal sandstone. 

Data include single-zircon U-Pb ages, Lu-Hf isotope 

ratios and trace element distribution patterns, which 

provide new insights into the sources and recycling patterns 

of clastic material in Fennoscandia in late Paleozoic 

time, and which help to refine the position of the sandstone 

within the evolutionary history of the Oslo Rift. 

1

2 T. Andersen et al. NORWEGIAN JOURNAL OF GEOLOGY 

Figure 1 

a) Simplified geological map of the southern part of the Rendal graben segment in the Oslo Graben, Norway (after Høy and 

Bjørlykke 1980). Legend: i: Brumunddal sandstone (Mauset Fm), ii-iv: Permian lavas and volcaniclastic sediments (Bjørgeberg 

Fm.), v: Bruflat Fm. (Upper Silurian), vi: Ekre shale (Silurian). vii: Pentamerus limestone (lower Silurian), viii: 

Mjøsa limestone (middle Ordovician), ix: Fluberg Fm (middle Ordovician), x: Cambrian shale and sandstone, xi: Ringsaker 

quartzite (Neoproterozoic). The broken line is an inferred, unexposed fault. 

b) Main geological features of the Oslo Rift, after Larsen et al. (2008). Legend: SG Skagerrak Graben, VG: Vestfold Graben, 

AG: Akershus Graben, RG: Rendal Graben. 1: Mesoproterozoic basement, 2: Hedemark Basin (Neoproterozoic). 3: Cambrian 

to Silurian sedimentary rocks, 4: Permo-Carboniferous volcanic and sedimentary rocks. 5: Permian intrusive rocks. 

6: Central volcanoes (cauldrons), 7: Transfer zones, 8: Brittle faults. The red rectangle marks the position of Fig. 1a. 

c) The outline of the Oslo Rift system, the preserved parts of the Permian deposits (red and deep purple; after Larsen et al. 

2008) and its relation to the tectonostratigraphic units mentioned in the text: Legend: C: Caledonian nappes, H: Hedemark 

Basin (Neoproterozoic), D: Dala-Trysil basin (Mesoproterozoic). TIB: Transscandinavian Igneous Belt (Late Paleoproterozoic). 

O: Orsa sandstone (Silurian), Ö: Öved sandstone (Silurian). On-shore lower Paleozoic sediments in the Oslo Rift are 

shown in pale blue. 

To evaluate the potential of post-Caledonian continental 

sediments outside of the rift area as intermediate repositories 

for detritus in the Brumunddal sandstone, we have 

also analysed U-Pb and Lu-Hf isotopes in detrital zircons 

from a sample of the Silurian continental sandstone from 

Orsa in central Sweden.


Geological background 

The late Paleozoic Oslo Rift (Fig. 1) encompasses four 

separate graben units. The southernmost unit (the 

Skagerrak Graben) is entirely situated offshore and features 

a symmetrical geometry. In contrast, the three onshore 

units (Vestfold, Akershus and Rendalen grabens) 

are half-grabens of contrasting polarities. The graben 

units are separated by NNW-SSE- and NNE-SSE-striking 

transfer zones (Fig. 1b). Volcanic rocks and batholiths 

dominate the the Vestfold and Akershus grabens, 

whereas the Rendalen Graben is devoid of batholiths. 

Still, some lava flows exist here, albeit minor in volume 

as compared to the southern graben units (Larsen et al. 

2008). 

The rift evolution can be divided into several distinct 

stages (Ramberg and Larsen 1978, Neumann et al. 2004, 

Larsen et al., 2008). The proto-rift stage is characterized 

by alluvial and minor marine deposits of Pennsylvanian 

age, the Asker Group (Olaussen et al., 1994, Larsen et 

al. 2008). Volcanism during the main rifting stage produced 

basaltic lava (B1) which form thick sequences 

in the Vestfold Graben, thinning northwards and to 

zero around Oslo. These lavas are of late Carboniferous 

to earliest Permian age (300 ± 1Ma, Corfu & Dahlgren 

2008). Graben subsidence continued and large volumes 

of rhomb porphyry lava (latite) and minor basalt erupted 

through large fissure volcanoes. Also, thick volcanoclastic 

and sedimentary units were deposited during this 

stage. The composite Larvik pluton in the Vestfold graben 

intruded in this stage, and has been dated to 298-292 

Ma by ID-TIMS U-Pb on zircons (Dahlgren et al., 1996, 

1998, Dahlgren 2010). In the subsequent central volcano 

stage, volcanic activity was concentrated in distinct 

volcanic centres with local trends of magmatic evolution, 

terminating in caldera collapse. The central volcanoe s 

are preserved as between 15 and 20 cauldron structures, 

representing sections through sub volcanic magma 

chambers , ring-dike systems and downfaulted blocks of 

volcanic rocks. U-Pb ages are not yet available for the 

rocks that formed during this stage, but Rb-Sr ages suggest 

a relatively long duration (ca. 280-240 Ma, Sundvoll 

et al. 1990). The final stages of rift evolution were characterized 

by emplacement of intermediate to felsic batholiths, 

ranging in composition from larvikite through syenite to 

granite. U-Pb ages on zircons from different intrusions 

suggest emplacement ages between 272 Ma and 286 Ma 

(Pedersen et al. 1995, Haug 2007, Andersen et al. 2009c), 

but Rb-Sr isochron ages range down to 241 Ma (Sundvoll 

et al. 1990). 

The Rendalen Graben is delineated eastwards by the 

westerly facing extensional Rendalen fault and is superseded 

to the north by several smaller sub-graben of varying 

polarity. One of these, informally termed the Narud 

halfgraben, occupies the Brumunddalen area, its Permian 

sequence being preserved in the rotated hanging wall 

fault block of the westerly dipping Gråkvern fault (Fig. 

1a). The fill of the Narud halfgraben is subdivided into 

two formations, namely a lower 250 m thick volcanic 

unit, the (Bjørgeberget formation) of Early Permian age 

(Sundvoll et al. 1990), and the upper, conformable c.750 

m thick undated Brumunddal sandstone unit, which has 

been suggested by Eliassen et al. (in prep.) to be renamed 

the “Mauset formation” (Fig. 2). 

The Bjørgeberg formation rests unconformably on folded 

Cambro-Silurian sediments and consists of four rhomb 

porphyry lava flows (RPx-u; Rosendal 1928), intercalated 

with up to 30m thick, volcanoclastic conglomerates 

and sandstones. The conglomerates consist entirely 

of rhomb porphyry lava clasts while the sandstones 

contain 63-65% quartz, 30-32 % feldspar (mostly plagioclase) 

and 4-6% lava fragments and hence classify as an 

arkosic arenite. The volcanoclastic rocks are interpreted 

as alluvial fans derived from evolving basin-bounding 

fault escarpments or local highs. Deposition of these fans 

took place either in interludes between extrusive activity, 

or represent contemporaneous, lateral variations of the 

volcanic-sedimentary sequence (Eliassen et al. in prep.). 

The Mauset formation (informally the Brumunddal sandstone) 

is dominated by red and yellow sandstones with 

subordinate mudstones and carbonates. The rocks of the 

Mauset formation can be subdivided into three units. 

The lowermost unit (c. 200 m thick) consists of red sandstone 

with a mineralogical composition and texture similar 

to that of the arkosic arenites interlayered with the 

lavas in the Bjørgeberg formation. They are interpreted 

as wadi deposits characterized by a stable high watertable. 

This unit is superseded by a 500 m thick sequence 

of texturally mature, fine to medium grained subarkosic 

arenite with up to 77% quartz, 21% feldspar and less 

than 2% lava fragements. The lowermost 250 m of this 

unit is interpreted as eolian dune, interdune and sandflat 

deposits interbedded by fluvial stream and alluvial 

plain sandstones and siltstones with evaporite pseudomorphs. 

The upper 250 m of the middle unit consists 

entirely of eolian sands. The Mauset formation is topped 

by a 50 m thick sequence of texturally and mineralogically 

im mature sandstones somewhat similar to the 

lower unit, but which may have more than 40% feldspar 

and less lava fragments. This arkosic arenite is interbedded 

by siltstones and carbonates that include fresh-water 

stromatolites. It was deposited in ephermal lakes, ponds 

and sabkhas typical of a mixed playa and alluvial environment 

and is indicative of a second event marked by 

high water-level (Eliassen et al., in prep.). 

Taken collectively, the Mauset formation is a typical 

eolian/fluvial wadi deposit (Rosendahl, 1929, Olaussen 

et al 1994, Eliassen et al. in prep.) much like what has 

been described for the Upper Rotliegendes in the Northern 

and Southern Permian Basins in northern Europe 

(e.g Glennie 1990). The limited outcrops do not allow 

interpretation of eolian dune types (e.g as barchans, 

transverse, seif or dras), nor can reliable determinations 

3


of paleo-wind direction be made. A similar uncertainty 

exists for the sediment transport directions of 

the ephemeral alluvial stream deposits, which are likely 

to have been determined by local dune topography and 

structural grain as developed by graben margin faulting 

(e.g. Gabrielsen et al. 1995). Still, it is considered 

unlikely that the 750 m thick eolian/wadi deposit of the 

Mauset Formation was constrained to the 6.5 x 2.5 km 

half graben in the Brumundal area alone. It is more likely 

that the Narud halfgraben is one preserved fragment of 

a wider desert area that covered most of the northern 

part of the Oslo Rift and its vicinity. Hence the deposits 

of the Brumundal sandstone might be interpreted as 

a northern arm or extension of the Rotliegendes of the 

Northern Permian Basin (Olaussen et al. 1994, McCann 

1998, Larsen et al. 2008, Eliassen et al. in prep). Consistent 

with this interpretation is the occurrence of up to 1 

m thick beds of red coloured, cross-stratified sandstone 

with well-sorted and well-rounded grains interpreted as 

Figure 2: Stratigraphy of the Brumunddal 

sandstone 

eolian or reworked eolian deposits (Olaussen et al 1994). 

These beds occur within the volcanoclastics interbedded 

with the lava pile in the Akershus and Vestfold Graben 

further south in the Oslo Rift. In a reconstructed paleogeographic 

setting, the provenance for the Mauset Formation 

is suggested to be both of local origin and from 

the basin fill of the rift further south. 

The Orsa sandstone is an erosional remnant of Silurian 

continental sandstone preserved within the Siljan impact 

structure in central Sweden (“O” in Fig. 1c). It is the 

youngest preserved member of the lower Paleozoic sedimentary 

sequence in the area, unconformably overlying 

limestones and shales of Llandowery to Wenlock age 

(Thorslund 1960). The deposit consists of a cross-bedded, 

conglomeratic to coarse sandy lower part, overlain 

by a main sequence of loosely consolidated, fine-grained 

quartz arenite, which is locally carbonate cemented and 

interlayered with clay-rich units (Hjelmqvist 1966). A


sample was included in this study to evaluate the potential 

of post-Caledonian sediments as intermediate hosts 

for detrital material, rather than to establish a detailed 

provenance history. The Orsa sandstone was chosen as 

an example rather than other Silurian continental sandstones 

(e.g. the Öved sandstone in Scania (Jeppson & 

Laufeld 1986), “Ö” in Fig. 1c; the Ringerrike group in the 

Oslo Rift) because it is situated between the site of deposition 

of the Brumunddal sandstone and potential Paleoproterozoic 

and older protosources in the Svecofennian 

and Archaean domains of Fennoscandia. The sample 

studied is a fine-grained, red, friable sandstone without 

clay or carbonate (Hjelmqvist 1966). 

Zircon petrography 

The present study was based on a sample of red, friable 

sandstone from the upper part of the Mauset formation 

(N 60° 54.59’, E 10° 58.89’). The sample is a carbonate-cemented 

quartz sandstone pigmented by finely 

disseminated iron oxide and containing minor to accessory 

amounts of detrital zircon. Most detrital zircon 

Figure 3. Backscattered electron (BSE) 

images of detrital zircon from the 

Brumunddal sandstone. Each image 

is identified by analysis number in 

tables of the Supplementary Material, 

U-Pb age and ε Hf at the U-Pb 

age. Each image is identified by analysis 

number, U-Pb age and eHf at the 

U-Pb age. Rare Archaean grains (a) 

are well-rounded and show little internal 

structure, except for a weak, BSE 

bright overgrowth (lower part). Paleo- 

to Mesoproterozoic grains (b, c, d) 

have well developed, more or less complex 

oscillatory zoning patterns typical 

for magmatic zircon, and are commonly 

well rounded. Permian grains 

(e, f) vary in shape from angular to 

well-rounded grain shapes, they show 

faint oscillatory zoning struc tures, and 

contain abundant, BSE-dark inclusions 

(feldspar, apatite). 

grains are well rounded, suggesting significant abration 

during transport (Fig. 3). A few grains giving Permian 

U-Pb ages are less well-abraded (Fig. 3e), but lack 

of rounding is not a universal feature of the youngest 

group of detrital zircons in the sandstone (compare Fig. 

3e with Fig. 3f). Most grains show the low-amplitude, 

short-wavelength oscillatory zoning in backscattered 

electron (BSE) intensity that is characteristic of magmatic 

zircons (Fig. 3b, c, d). Complex inner structures 

seen in some grains suggest that they are fragments of 

twins or intergrown groups of crystals (Fig. 3b). BSEbright 

overgrowths occur in some grains (e.g. Fig. 3a,c); 

these are generally too thin to allow separate analysis 

by LAM-ICPMS. Mineral inclusions in zircon (e.g. Fig. 

3c) were avoided during U-Pb and Lu-Hf analysis, but 

a few were penetrated during trace element ablations. 

Inclusions which were large enough to be integrated 

separately showed elevated phosphorus and middle or 

light REE, and have therefore been tentatively identified 

as apatite and monazite, respectively. 

5


Analytical methods 

Sample preparation was done at the Department of Geosciences, 

University of Oslo, and analysis of zircon by 

LAM-ICPMS at the GEMOC National Key Centre, Macquarie 

University and at the Department of Geosciences, 

University of Oslo. Zircon grains were separated from 

rock samples by standard crushing and mineral separation 

techniques, followed by hand-picking under a binocular 

microscope. Selected grains from the two least 

magnetic fractions (magnetic and non-magnetic at 1.5 

A magnet current in the Franz separator at 15º standard 

tilt and slope) were cast in epoxy disks for LAM-ICPMS 

analysis. The mounts were ground and polished to expose 

the zircons, and thoroughly cleaned before analysis. Zircons 

were examined in transmitted and reflected light, 

and by electron backscatter imaging in a Cameca SX50 

electron microprobe. Each imaged zircon was assigned a 

number, which also serves to identify analyses in Tables 

2, 3 and 4 of Supplementary Material. Standards were 

mounted in separate epoxy disks. 

LAM-ICPMS U–Pb dating. 

U–Pb dating of the Brumunddal sandstone sample 

was made with a HP 4500 series 300 inductively coupled 

plasma quadrupole mass spectrometer (QICPMS), 

attached to a Merchantec/New Wave LUV213 Nd-YAG 

laser microprobe at GEMOC. Analytical procedures 

were as described by Jackson et al. (2004). The analytical 

protocol allows accurate correction for U/Pb fractionation 

by sample-standard bracketing. Standard zircon 

GJ-1 (608 Ma, Jackson et al. 2004) was used for calibration, 

while standard zircons 91500 (Wiedenbeck et 

al. 1995) and Mud Tank (Black and Gulson 1978) were 

run frequently as unknowns to provide an independent 

control on reproduci bility and instrument stability. 

The GLITTER software (Griffin et al. 2008) was used for 

data reduction and calculation of individual U–Th–Pb 

ages. The contents of common lead were estimated from 

the 3D discordance pattern in 206Pb/ 238U – 207Pb/ 235U – 

208 232 Pb/ Th space, and corrected accordingly (Andersen 

2002). 

Additional U-Pb analyses of detrital zircon from the the 

Orsa sandstone were made with a Nu Plasma HR multicollector 

ICPMS and a NewWave LUV213 laser microprobe 

at the Department of Geosciences, University of 

Oslo, using protocols for data acquisition and standardization 

described by Rosa et al. (2009). 

Lutetium-hafnium isotopes. 

Lu-Hf analyses of zircon from the Brumunddal sandstone 

were made with a NuPlasma multicollector ICPMS 

attached to a LUV213 Nd-YAG laser microprobe at 

GEMOC, using protocols described by Griffin et al. 

(2000) and Andersen et al. (2002b). Ablation conditions 

were: Beam diameter: 55 μm (aperture imaging mode), 

pulse frequency: 5 Hz, beam energy density: ca. 2 J/cm 2 , 

static ablation. Each ablation was preceded by a 30 seconds 

on-mass background measurement. Masses 172 to 

180 were measured in Faraday collectors, using the standard 

collector block of the NU Plasma mass spectrometer. 

The total Hf signal obtained was in the range 1.5- 

3.0 V. Under these conditions, 120-150 seconds of ablation 

are required to obtain an internal precision of ≤ ± 

0.000020 (1SE). Isotopic ratios were calculated using the 

Nu Plasma time-resolved analysis software. The raw data 

were corrected for mass discrimination using an exponential 

law, the mass discrimination factor for Hf was 

determined assuming 179 Hf/ 177 Hf= 0.7325, the observed 

2 SE uncertainty of the Hf mass discrimination factor 

was better than 0.5%, and was propagated through to 

the error of the final result. The isobaric interferences 

on 176 Hf by 176 Lu and 176 Yb were corrected from measurements 

of interference-free 172 Yb and 175 Lu. The observed 

mass discrimination for Hf was used as a proxy for those 

of Lu and Yb. 176 Lu/ 175 Lu= 0.02669 was used for correction 

of Lu interference on mass 176 (DeBievre & Taylor 

1993). The 176 Yb/ 172 Yb ratio used in the correction procedure 

(0.5870) was determined from multiple runs of 

Yb-doped JM475 Hf standard solutions. Lu-Hf analyses 

of zircon from the Orsa sandstone were made with 

the Nu Plasma HR instrument in Oslo, using procedures 

described by Andersen et al. (2009a) and Heinonen et al. 

(2010). 

A value for the decay constant of 176 Lu of 1.867 x 10 -11 

a -1 has been used in all calculations (Söderlund et al. 

2004; Scherer et al. 2001, 2007). For the calculation of 

ε Hf values we used present-day chondritic 176 Hf/ 177 Hf = 

0.282785 and 176 Lu/ 177 Hf = 0.0336 (Bouvier et al. 2008). 

We have adopted the depleted mantle model of Griffin 

et al. (2000), modified to the λ 176 Lu and chondritic composition 

used, which produces a present-day value of 

176 Hf/ 177 Hf (0.28325, εHf = +16.4) similar to that of average 

MORB over 4.56 Ga, from chondritic initial hafnium at 

176 Lu/ 177 HfDM = 0.0388. 

Trace elements. 

Trace elements in zircon were analysed using an Agilent 

7500 quadrupole ICPMS and a New Wave Research 

LUV213 laser microprobe at GEMOC. Ablation 

conditions were: Beam diameter: 40 μm (aperture imaging 

mode), pulse frequency: 5 Hz, beam energy density: 

ca. 5 J/cm 2 , giving a laser energy at the sample of 0.07 mJ. 

A fast scanning protocol was employed, analysing masses 

from 29 (Si) to 238 (U), with 10 ms dwell times on all 

masses analysed except 206 (15 ms), 207 (30 ms) and 238 

(15 ms). Ablations lasted for 120 seconds, and each was 

preceded by a 60 seconds background measurement with 

the laser off. NIST 610 silicate glass was used as external 

standard, with 29 Si as the internal standard, assuming 

a stoichiometric composition for zircon with 32.78 

weight % SiO 2 . Data reduction was made off-line using 

the software package GLITTER 4.4 (Griffin et al. 2008).


Figure 4. U-Pb data for detrital zircon 

from the Brumunddal sandstone. 

a: Tera Wasserburg diagrams showing 

data for zircon crystals in the magnetic 

and non-magnetic fractions from 

the Franz magnetic separator at 1.5 A 

separately. Note that Permian zircons 

occur only in the magnetic fraction. 

b: Conventional concordia diagram 

showing the Permian fraction. The 

three grains represented by white 

fill define a concordia age at 285±4 

Ma, the fraction as such a weighted 

average 206 Pb/ 238 U age at 283±4 Ma, 

which is the preferred estimate of 

the crystallization age of the source 

rock(s). 

Accuracy and external precision of the individual elements 

estimated from repeated analyses of reference zircon 

GJ-1 are given in Table 1 of Electronic Supplement. 

Results 

0.20 

2800 

0.16 

0.12 

0.08 

Uranium-lead ages of detrital zircons in the Brumunddal 

sandstone. 

Of a total of 105 grains dated by U-Pb (Table 2 of Electronic 

Supplement), 94 were concordant or less than 10 

% normally discordant (central discordance). The zircons 

yield a large range of U-Pb ages, from late Archaean 

207 Pb/ 206 Pb 

a 

2000 

1200 

0 4 8 12 16 20 24 

Non-magnetic fraction 

400 

0.04 

0 4 8 12 16 20 24 

b 

206 Pb/ 238 U 

0.052 

0.047 

2800 

2000 

1200 

238 U/ 206 Pb 

280 

300 

400 

Concordia Age =285 ±4 Ma 

(2 , decay-const. errs ignored) 

MSWD of concordance and equivalence = 1.6 

Magnetic fraction 

320 

0.042 260 

Weighted average 

0.037 

240 

0.26 0.28 0.30 0.32 0.34 0.36 0.38 

206 Pb/ 238 U age: 

283±4 (95% conf.), 9of10 grains 

Wtd by data-pt errs only, 1of10 rej. 

MSWD = 1.7 

207 Pb/ 235 U 

0.20 

0.16 

0.12 

0.08 

0.04 

(2.6 Ga) to Permian (Fig. 4a), with a pronounced frequency 

maximum at ca. 1 Ga, and subordinate maxima 

at ca. 1400 Ma, 1600 Ma, 1800 Ma and in the Phanerozoic, 

with marked hiatuses in the Neoproterozoic and 

early Paleoproterozoic (Fig. 5a). Although the magnetic 

and non-magnetic zircon fractions show largely similar 

age distribution patterns, a distinct group of late Paleozoic 

zircons (270-310 Ma) has been observed only in 

the magnetic fraction (Fig. 4a). Of eleven zircons in this 

group, eight are 10 % or more discordant. The three concordant 

zircons give a concordia age of 285±4 Ma, and 9 

zircons define a weighted average 206 Pb/ 238 U age of 283±4 

Ma (Fig. 4b, 5b), indistinguishable from the concordia 

age. Another group of similar size (9 grains) give ages 

7


Number 

Number 

20 

15 

10 

5 

between 526 and 398 Ma, with a peak at 440 Ma, which 

is compatible with igneous or metamorphic source rocks 

in Caledonian nappes (e.g. Gee et al. 2008). Two analyses 

give ages of 346±4 Ma and 361±4 Ma. These grains do 

not differ from the Caledonian fraction in terms of trace 

element distributions and Lu-Hf characteristics (below), 

and it is therefore suggested that they are Caledonian 

grains which have suffered sufficient lead-loss in Permian 

or recent time to modify their ages, while remaining 

concordant within error. 

Orsa sandstone, Sweden (Silurian) 

Brumunddal sandstone 

0 

200 600 1000 1400 1800 2200 2600 3000 

5 

4 

3 

2 

1 

a 

Detrital zircon age, Ma 

b Weighted average 206Pb/ 238U 

age: 283±4 Ma 

0 

250 260 270 280 290 

300 310 

Zircon 206 Pb/ 238 U age, Ma 

Relative probability 

Relative probability 

Figure 5. Accumulated probability distribution 

diagrams showing: 

a): The whole population of dated zircons 

( 207Pb/ 206Pb ages for t> 600 

Ma, 206Pb/ 238U ages for t< 600 Ma), 

analyses with more than ± 10 % 

discordance excluded. The colourfilled 

background histogram shows 

the age distribution of detrital zircons 

in the Silurian Orsa sandstone, 

Sweden (Table 5a of Supplementary 

material). 

b): Permian zircons (n=11, including 

discordant grains). Broken outline 

indicates zircons that are not used to 

calculate ages. 

Lutetium-Hafnium isotopes of detrital zircon in the 

Brumund dal sandstone 

In Fig. 6 (data from Table 3 of Electronic Supplement), 

initial 176 Hf/ 177 Hf for 82 concordant or near-concordant 

zircons are plotted at the preferred U-Pb age (i.e. 

207 Pb/ 206 Pb ages for zircons older than 600 Ma, 206 Pb/ 238 U 

ages for younger zircons). 

The late Paleozoic zircons form a coherent group with 

distinct initial Hf character. When zircons that are more 

than 10% discordant in U-Pb are included in the calculation, 

nine zircons give a weighted average 176 Hf/ 177 Hf=


Initial Hf/ Hf 

176 177 

0.2835 

DM 

0.2830 

CHUR 

0.2825 

0.2820 

0.2815 

0.2810 

0.2805 

+10 

0 

-10 

Oslo Rift 

magmas 

TIB crust 

Age range of Caledonian 

magmatism 

Sveconorwegian granites 

Inherited zircon in 

Sveconorwegian 

granites 

Inherited zircon in TIB, 

Intrusions in the 

Archaean Domain 

0.28272±0.00004 (95% confidence) or ε Hf =4.2±1.5, with 

no rejections. The spread of initial 176 Hf/ 177 Hf in this 

group is comparable to the analytical error, suggesting 

that these zircons originated from one or more protosource 

rocks within the rift with homogeneous Lu-Hf 

characteristics. 

Zircons older than these show considerable variation in 

initial 176 Hf/ 177 Hf, ranging from well below CHUR (i.e. 

ε Hf


could be analysed for Hf isotopes, and yielded nearchondritic 

176 Hf/ 177 Hf at its 207 Pb/ 206 Pb age. 

Trace element compositions of detrital zircon in the Brumunddal 

sandstone 

The trace element distribution patterns observed in 

the analysed samples are typical for magmatic zircon 

(Hoskin and Schaltegger 2003, Belousova et al. 2002), 

with positive anomalies for Hf, U, Th and Ce (Fig. 7, data 

from Table 4 of Electronic Supplement). Hf concentrations 

range from ca. 0.6 to ca. 2.4 percent, and total REE 

from 33 to 2000 ppm. Source-dependent trace element 

combinations spread from the less U and LREE enriched 

part of the field of variation of zircon from granitoids 

towards that of zircon from mafic rocks (Fig. 8, fields 

from Belousova et al. 2002). This is reflected by the 

CART classification algorithm of Belousova et al. (2002), 

which gives a dominance of “dolerite” and different types 

of “granitoids” as suggested source rocks (Table 1). A few 

zircon analyses indicate source rock types such as “kimberlite” 

and “carbonatite”. Although both rock types 

are known to exist in Fennoscandia (e.g. O’Brien et al. 

2005, Rukhlov & Bell 2010), they are rare, and unlikely 

to have yielded significant amounts of detrital material 

to the Brumunddal sandstone. Zircon grains indicating 

uncommon source rock types show anomalous REE patterns 

with low HREE and /or HREE / LREE ratios, combined 

with Eu/Eu* > 0.5 (NM15A-02, 04, 17, 21, 29, 32, 

45, 66, Fig. 9) and probably formed by crystalli zation 

from hydrothermal solutions or during metamorphism 

rather than from magma; the relatively flat HREE pattern 

in NM15A-32 suggests coexistence with garnet (e.g. 

Hoskin & Schaltegger 2003). 

Chondrite normalized REE distribution patterns are 

strongly enriched in the heavy REE, as is common in zircon; 

all grains analysed have a marked positive Ce anomaly 

and a variable, negative Eu anomaly (Fig. 9). A majority 

of the analyses have (Yb/Sm) CN = 30 to 200, and deep, 

negative Eu anomalies (Eu/Eu* < 0.4). With the exception 

of a few grains with apparent LREE enrichment due 

to the presence of solid inclusions (see below), none of 

the zircons analysed from the Brumunddal sandstone 

show the characteristic flattening of the LREE pattern 

observed in zircons from “granitoids” by Belousova et al. 

(2002), and the patterns in general resemble those of zircon 

from “dolerite” and “basaltic” sources (Fig. 9). 

Anomalous LREE-enrichment seen in a few analyses 

(e.g. NM15A-57 and -67) are from ablations that 

have picked up minor amounts of LREE-rich inclusions 

(monazite?) in the zircon. Other anomalous analyses 

show relative enrichment of middle REE and a shallow 

Eu anomaly (e.g. NM15A-37, Fig. 9). Again, this is due 

to accidental ablation of small inclusions in the zircon, in 

this case most probably of apatite. 

The zircons of late Paleozoic age form a consistent 

group with relatively low HREE enrichment (Yb/Sm) CN 

= 20-30). Of the 11 grains in this group, four have distinctly 

shallower Eu anomalies than the rest (Eu/Eu*>0.4, 

compared to Eu/Eu*≤ 0.3, Fig. 9). The patterns of the 

latter resemble those of Oslo Rift larvikite analysed by 

Belousova et al. (2002) in terms of pattern curvature 

and Eu anomaly, but at an overall lower REE enrichment 

level. The differences in concentration level may, 

however, be an effect of the different methods of internal 

standardization used in the two studies (assumed 

stoichiometric SiO 2 value in this study vs. Hf analysed by 

electron microprobe). 

U-Pb and Lu-Hf characteristics of Silurian sandstone 

The age pattern of detrital zircon in the Silurian Orsa 

sandstone mimics that of the Brumunddal sandstone 

with age maxima in the late Paleoproterozoic, at ca. 1500 

Ma and in the late Mesoproterozoic (Fig. 5a, data from 

Table 5a of Electronic Supplement). There is also a small 

contribution of Phanerozoic zircon, presumably from a 

Caledonian source. The variation of initial Hf isotope 

character also overlaps that of the Brumunddal sandstone 

(Fig. 6, data from Table 5b of Electronic Supplement 

l). The clearest difference between the two distribution 

patterns is due to a single, highly U-Pb discordant 

zircon with ε Hf =-21 at 1376 Ma (Orsa-38, Table 4); this 

is most probably a late Archaean or Paleoproterozoic zircon 

that has lost a significant amount of lead in a thermal 

event in Mesoproterozoic time. 

Discussion 

U-Pb, Lu-Hf and potential protosources of detrital zircon 

in the Brumundal sandstone. 

The detrital zircon ages from the Brumunddal sandstone 

range from late Archaean to Phanerozoic, with initial Hf 

istope signatures from near depleted mantle to crustal 

ratios. The distribution of age and initial 176 Hf/ 177 Hf or ε Hf 

in Fig. 6 is, however, far from random, and can be related 

to known crustal evolution events in Fenno scandia. 

Precambrian protosources. The near-CHUR initial 

176 Hf/ 177 Hf of the single late Archaean detrital zircon analysed 

for Lu-Hf falls within the age and 176 Hf/ 177 Hf range 

of late Archaean granitoids from eastern Fennoscandia 

(Patchett et al. 1981, Lauri et al. 2011). Archaean rocks 

are exposed only in E Finland, NE Sweden N Norway 

and NW Russia, at considerable distance from their sites 

of deposition. 

The Precambrian rocks of the Fennoscandian shield 

immediately bordering the Oslo Rift (Fig. 1c) belong 

to the late Paleoproterozoic Transscandinavian Igneous 

Belt (TIB, Högdahl et al. 2004) and the early Mesoproterozoic 

(“Gothian”) magmatic arc fragments (Andersen 

et al. 2004). The TIB was formed by multiple episodes


Figure 7. Spider diagrams 

showing the overall 

variation of zircon 

compositions in the 

Brumunddal sandstone, 

normalized to chondritic 

values of McDonough 

and Sun (1995). The 

shaded background field 

is the interquartile range 

of zircon compositions in 

granitoids (Belousova et 

al. 2002). 

Hf ppm 

Y ppm 

25000 

20000 

15000 

10000 

5000 

1 

0 

10 6 

10 5 

10 4 

10 3 

10 2 

10 

ppm in zircon / ppm in chondrite 

Mafic rocks 

1.0 6 

1.0 4 

1.0 2 

1.0 

1.0 -2 

Pb La Ce Pr Nd Sm Eu Gd Th Dy Y Ho Er U Yb Lu Hf Nb Ta Ti P 

Granitoids 

Y ppm Y ppm 

1 

1 10 10 1 10 

Y ppm U ppm 

2 

102 10-2 10-1 103 103 104 104 105 105 0.1 1 10 100 

Nb/Ta 

10 6 

10 5 

10 4 

10 3 

10 2 

10 

10 6 

10 5 

10 4 

10 3 

10 2 

10 

1 

1 10 100 

Yb/Sm 

Figure 8. Source-indicat ing element and element -ratio plots showing that a significant proportion of the detrital zircons analysed 

falls between the main fields of zircon derived from granitoid and mafic (i.e. basalt, dolerite) sources (reference fields 

from Belousova et al. 2002). 

11 

10 3


ppm zircon / ppm chondrite 

10 4 

10 3 

10 2 

10 

1 

10 -1 

10 4 

10 3 

10 2 

10 

1 

10 -1 

10 4 

10 3 

10 2 

10 

1 

10 -1 

10 -2 

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 

< 320 Ma 

Larvikite, 

interquartile range, 

Belousova et al. (2002) 

900-1250 Ma 

M15A-01 

NM15A-67 

NM15A-57 

1400-1500 Ma 

NM15A-26 

NM15A-62 

NM15A-29 

320-900 Ma 

900-1250 Ma 

Basalt Dolerite Granitoids 

> 1600 Ma 

NM15A-04 

NM15A-37 

NM15A-35: inclusion 

NM15A-21 

NM15A-32 

La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu 

Figure 9. Chondrite normalized REE patterns for zircon from the Brumunddal sandstone, divided into six groups (1600 Ma, 

see text for further explanation). Reference fields for all except the


of magmatic intrusion in the period 1.88 to 1.68 Ga; the 

TIB intrusions close to the Oslo Rift belong to the youngest 

part of the belt, known as TIB3, with intrusive ages 

around 1.68-1.67 Ga (Larson and Berglund 1992, Heim 

et al. 1996, Andersen et al. 2009a). Despite the large size 

and long time of emplacement of the TIB province, the 

initial Hf signature of magmatic zircon from TIB granitoids 

shows little variation, which can be represented 

by initial ε Hf = +2±3 at 1.88 Ga and 176 Lu/ 177 Hf≈ 0.015 

(Andersen et al. 2009a). Given the short distance to 

potential source rocks within the TIB, the low abundance 

of TIB-like detrital zircons in the Brumunddal sandstone 

(Fig. 5a, 6) is somewhat surprising. On the other hand, 

a small group of Paleoproterozoic zircons with negative 

ε Hf overlaps the field of inherited zircon of possible 

Archaean origin in a TIB intrusion at the eastern margin 

of the belt in Sweden (Andersson et al. 2006, Andersen 

et al. 2009a). Furthermore, Paleoproterozoic zircons with 

such initial Hf characteristics are found in Paleoproterozoic 

granites within the Archaean domain of Finland 

(Patchett et al. 1981, Andersen & Lauri 2010). Paleoproterozoic 

crust with Lu-Hf properties similar to the rocks 

of the Transscandinavian Igneous Belt or their source in 

the deep crust has contributed source material to most of 

the Mesoproterozoic crustal magmatic events that can be 

recognized in southwestern Fennoscandia, even far west 

of the present TIB outcrop area (Andersen et al. 2001, 

2002a, 2009b, Andersen 2005, Pedersen et al. 2009). 

Detrital zircons of early Mesoproterozoic age with initial 

176 Hf/ 177 Hf approaching the DM curve resemble zircons 

from both Mesoproterozoic continental margin arc complexes 

(Andersen et al. 2002b, 2004) and those from felsic 

rocks from the Telemark block, which are still poorly 

constrained by Lu-Hf data. A small, coherent group 

of 1.45-1.50 Ga zircons with ε Hf = -3 to -8 can either be 

derived from a coeval, crustal source, or they can be TIBderived 

zircons that have lost lead in a Mesoproterozoic 

tectonothermal event. 

During the Sveconorwegian period, southwestern Fennoscandia 

was affected by several episodes of mafic and 

felsic magmatism in different tectonic settings, ranging 

from ca. 1300 Ma to 920 Ma (Bingen and van Breemen 

1998, Andersen et al. 2001, 2002a,b, 2007, 2009b, 

Vander Auwera et al. 2003, Pedersen et al. 2009). Magmatic 

events involving crustal underplating with mafic 

material with a depleted mantle Hf isotope signature 

(ε Hf ≈ +12) at 1.3 and 1.22 Ga have been documented in 

the Telemark block; after each of these events, the initial 

Hf isotope composition of zircon from magmatic rocks 

evolved towards less radiogenic compositions due to 

increasing effects of crustal source components (Andersen 

et al. 2007, Pedersen et al. 2009). Initial ε Hf of the 

large group of Sveconorwegian zircons overlaps well 

with the combined range of magmatic and inherited zircon 

from Sveconorwegian granitoids (Andersen et al. 

2002a, 2007, Pedersen et al. 2009). 

13 

Caledonian protosources. Synorogenic intrusions in the 

nappes of the Scandinavian Caledonides that could be 

sources of early Paleozoic detrital zircon have been dated 

to 430-500 Ma (Gee et al. 2008 and references therein). 

The Caledonian mountain chain was mainly levelled 

by mid-Carboniferous time (Gabrielsen et al. 2005, 

2010), but late Permian – Triassic uplift is likely to have 

enhanced relief to the west and northwest of the Oslo 

Graben system (Rohrman et al. 1995, Gabrielsen et al. 

2010). By the early Triassic, sediments of Caledonian origin 

were transported southwards to become deposited 

in the Central European Basin (Paul et al. 2009). Unfortunately, 

no Hf isotope data are available from Caledonian 

intrusions. The considerable range of ε (+ 1.4 

Hf 

to -11) observed suggests, however, that these zircons 

either crystallized from magmas with variable amounts 

of a TIB-like crustal component, or that they are zircons 

from Sveconorwegian protoliths that have lost lead in 

Caledonian metamorphism. 

Within-rift sources. Permian grains make up ca. 10 % of 

the analysed detrital zircon population (Fig. 5). Rocks 

from the Larvik pluton have been dated to 292-298 Ma 

by U-Pb on zircon and baddeleyite (Dahlgren et al. 1998, 

Dahlgren 2010). This event was, however, followed by 

the emplacement of syenitic to granitic plutons around 

280 Ma (Pedersen et al. 1995, Haug 2007). Magmatic 

zircon from larvikite, nordmarkite and biotite granite 

in the ca. 280 Ma Sande and Drammen plutons have 

ε Hf in the range - 4 to +6, whereas zircons from pegmatites 

in the Larvik pluton are slightly more radiogenic 

at ε Hf = +8 ± 2 (Andersen et al. 2009b); no evidence has 

so far been found of hafnium derived from a late Paleozoic 

depleted mantle reservoir (ε Hf ≈ + 15) in any of these 

rocks. The 176 Hf/ 177 Hf at 280 Ma of the late Paleozoic 

zircons falls within the age and Hf isotope range of zircon 

in the younger group of intermediate to felsic intrusions 

in the rift. Two distinct types of REE patterns have 

been observed, with “deep” and “shallow” Eu anomalies, 

respectively (Fig. 9). The first group has REE pattern 

geometry similar to that of zircon from larvikite reported 

by Belousova et al. (2002), and therefore most likely originates 

from intermediate intrusive rocks. The other type 

must originate from other syenitic or (alkali) granitic 

plutonic rocks, which cannot be identified. 

Recycling of sedimentary rocks 

Most of the main documented crust-forming events in 

Fennoscandia are represented in the detrital zircon population 

of the Brumunddal sandstone, including accessory 

late Archaean and Paleoproterozoic zircons that 

must have originated in protosources east of the Transscandinavian 

Igneous Belt. It is highly unlikely that zircon 

from such a wide range of sources could have been 

brought into the rift in a single event of one-way transport 

from protosource to depositional basin. Rather, zircon 

must have been recycled through deposition in intermediate, 

sedimentary reservoirs followed by erosion and


new transport. This has eventually caused a relatively 

well mixed sample of the entire Fennoscandian surface to 

be brought into the Oslo Rift, including far-transported 

material from eastern Fennoscandia. 

The close similarity in age and initial 176 Hf/ 177 Hf distribution 

of detrital zircon in the Brumunddal and Orsa sandstones 

indicates that Silurian sandstone is a very likely 

immediate precursor. Silurian, continental deposits may 

have had a wide distribution in western and central Fennoscandia 

in post-Caledonian time; outside of the Oslo 

Rift they are now preserved only as small erosional remnants. 

Like the Orsa sandstone, the late Silurian Ringerrike 

Group within the Oslo Rift itself (Davies et al. 2005 

and references therein) contains material derived from 

both Fennoscandian and Caledonian sources (Turner & 

Whitaker 1976, Dahlgren & Corfu 2001). 

The Silurian continental sediments make up the youngest 

of several cover sequences in Fennoscandia, ranging 

in age from Paleoproterozoic (e.g. Claesson et al. 1993, 

Lundqvist & Persson 1999, Lahtinen et al. 2002, Bergman 

et al. 2008), through Mesoproterozoic (e.g. Bingen et al. 

2001, 2003, 2005b, Kuhonen & Rämö 2005, Andersen et 

al. 2004, Lamminen & Köykkä 2010) and Neoproterozoic 

(e.g. Nystuen et al. 2008, Bingen et al. 2005a, Lamminen 

et al. 2008, 2009) to Phanerozoic (e.g. Thorslund 

1960, Bjørlykke 1974a, Turner & Whitaker 1978, Davies 

et al. 2005). Each of the successive cover sequences must 

have contained clastic material recycled from older 

deposits, as well as detritus derived directely from crystalline 

protosource rocks. 

Basins such as the late Vendian – Cambrian Peri-Thornquist, 

and the Pre-Baltic sub-basins (Sliaupa et al. 1997, 

Poprawa et al. 1999), the southwestern part of the Baltic 

Basin (Lazauskiene et al. 2002), as well as the Scandinavian 

(Bjørlykke 1974b, Worsley et al. 1983) and German-Polish 

Caledonian foreland basins (Lazuskiene et 

al. 2002) may have acted as intermediate repositories for 

far-transported detritus, including Paleoproterozoic and 

Archaean detrital zircons of east Fennoscandian origin. 

The Caledonian forebulge may have promoted uplift and 

erosion of these basins, thus releasing zircons of an easterly 

origin. 

The most zircon-fertile rocks within each of the protosource-terranes 

are granitic. Nevertheless, the trace element 

patterns are more in line with mafic to intermediate 

source rocks than with granite. Zircons with a relatively 

less steep LREE pattern are also likely to be more heavily 

enriched in uranium and thorium (Fig. 7; Belousova 

et al. 2002), and would therefore more easily become 

metamict due to internal radiation damage than less U 

and Th enriched (and hence also more strongly LREE 

depleted) ones. Metamict zircon is brittle and much 

more likely to be destroyed by abrasion during transport 

than grains with an intact crystal structure. A scarcity of 

typical “granitoid” zircon is therefore to be expected in 

sediments with a long and complex transport history. 

The significance of Neoproterozoic detrital zircon ages in 

southwestern Fennoscandia. 

The Neoproterozoic was a quiet period in Fennoscandia, 

with little or no production of zircon-fertile igneous 

rocks. Nevertheless, Neoproterozoic detrital zircon ages 

have been reported from both the Neoproterozoic Hedmark 

basin (Bingen et al. 2005a) and from the Carboniferous 

Asker Group within the Oslo Rift (Dahlgren & 

Corfu 2001). In the Brumunddal sandstone, such zircons 

are conspicuously absent. A non-Fennoscandian source 

for Neoproterozoic zircon has been suggested (Dahlgren 

& Corfu 2001), but the range of 176 Hf/ 177 Hf in Neoproterozoic 

zircons in the Hedmark Basin with the observed 

U/Pb ages is fully compatible with Neoproterozoic loss of 

radiogenic lead in zircon from Sveconorwegian granitoids 

(Fig. 6). This could be caused by surface weathering 

prior to deposition, or by interaction with diagenetic 

fluids within the Hedmark Basin. 

Time of deposition and tectonostratigraphic status of the 

Brumunddal sandstone 

The weighted average U/Pb age of 283±4 Ma (Fig. 3b) 

confirms the previous estimate of the maximum age limit 

for sedimentation based on Rb-Sr ages of the underlying 

rhomb porphyry lavas (Sundvoll et al. 1990, Larsen et 

al. 2008). By the time of deposition of the Brumunddal 

sandstone, rocks belonging to the ca. 280 Ma “batholith 

stage” of rift evolution must have been exposed to erosion 

at the surface, which suggests that deposition took 

place during the final stages of rift evolution (stage 6 in 

the chronology of Larsen et al. 2008), rather than during 

or immediately after rhomb porphyry volcanism (i.e. in 

stage 3), which preceded emplacement of the younger, 

batholitic bodies. Plausible sources of Permian zircon are 

only exposed to the south of the site of deposition, in the 

intrusive complexes of the Akershus and Vestfold graben 

segments. These zircon grains must therefore have been 

transported from south to north along the rift axis. The 

observation that some of these zircons are quite severely 

rounded due to abrasion during transport indicates significant 

transport for some, but not for all of the young 

zircons observed. 

The present data suggest that sediment transport took 

place from south to north along the graben axis, perhaps 

suggesting a topographic gradient in this direction. It is 

of particular interest to note that one of the REE-patterns 

shows the presence of a deep (plutonic) source, suggesting 

that plutons became exposed and eroded in the 

southern part of the graben simultaneously with ongoing 

volcanism in the Brumunddal area. This situation is not 

unique and is paralleled in the present East African rift 

system (Barberi et al. 1970, 1972).


Why are Permian zircons only found in the magnetic 

fraction? 

Permian zircons were only found in the magnetic 

fraction at 1.5 A. With ca. 10% of the total number 

of analyzed zircons, this fraction is relatively large. 

However , if each of the two fractions from the magnetic 

separator are considered individually, their sizes are not 

sufficient to guarantee detection of even relatively large 

age fractions in each (Andersen 2005). However, it is still 

fully possible that the absence of syn-rift zircons from the 

non- magnetic fraction reflects a real physical property of 

these zircons. If so, relatively standard use of magnetic 

separation has efficiently removed them from the least 

magnetic fraction. The tendency of magnetic mineral 

separation to introduce undesirable bias in detrital zircon 

populations was pointed out by Sircombe and Stern 

(2002), and the present results could imply that such 

effects can be due to even a moderate use of the magnetic 

separator. 

Conclusions 

The Brumunddal sandstone was deposited during 

the final stages of evolution of the Oslo Rift, at a time 

when plutonic rocks formed during the ca. 280 “batholith” 

stage were exposed to erosion. The sandstone contains 

detrital zircon ranging in age from late Archaean to 

Permian. Combined U-Pb and Lu-Hf data suggest that 

detrital zircons from most of the region-wide petrogenetic 

events in Fennoscandia are represented in this single 

sediment, including zircons formed in events not represented 

by rocks within reasonable distance from the 

site of deposition. The likely explanation for this broad 

distribution of detrital zircon is recycling of material 

from Phanerozoic sedimentary cover on Fennoscandian 

basement, which may in turn contain material recycled 

from still older sedimentary rocks. Repeated transport 

and deposition has effectively mixed detrital zircon from 

a wide range of protosources, and depleted the detritus in 

metamict, high U zircons with typical “granitoid” trace 

element distribution patterns, leaving a population with 

a large proportion of zircons having trace element distributions 

more characteristic of “mafic” source rocks. 

Acknowledgments. - The research reproted her could be carried out 

thanks to economic support from the University of Oslo and Macquarie 

University. Thanks are due to Trine-Lise Knudsen for assistance in 

the field, Gunborg Bye-Fjeld for technical assistance with mineral separation, 

Justin Payne for help with the LAM-ICPMS, to Elena Belousova 

and Norman Pearson for many helpful discussions, and to Bernard 

Bingen and an anonymous reviewer for helpful comments. T.A. wants 

to thank GEMOC and its directors S.Y. O’Reilly and W.L. Griffin 

for repeated invitations to spend fruitful research time in Australia . 

This is publication no. 17 from the Isotope Geology Laboratory at the 

Department of Geosciences, University of Oslo and contribution 696 

from the Australian Research Council National Key Centre for the 

Geochemical Evolution and Metallogeny of Continents (http://www. 

gemoc.mq.edu.au). 

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