palaeoproterozoic to palaeozoic magmatic and metamorphic events in the shackleton range, east...

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Precambrian Research 172 (2009) 25–45

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Palaeoproterozoic to Palaeozoic magmatic and metamorphic events in theShackleton Range, East Antarctica: Constraints from zircon andmonazite dating, and implications for the amalgamation of Gondwana

T.M. Will a,∗, A. Zeha, A. Gerdesb, H.E. Frimmela, I.L. Millarc, E. Schmädicked

a Geodynamics and Geomaterials Research Group, Department of Geography, University Würzburg, Am Hubland, D-97074 Würzburg, Germanyb Department of Geosciences, University Frankfurt, Altenhöferallee 1, D-60438 Frankfurt, Germanyc British Antarctic Survey, c/o NERC Isotope Geoscience Laboratory, Nicker Hill, Keyworth, Nottingham NG12 5GG, UKd Department of Geology and Mineralogy, University Erlangen-Nürnberg, Schlossgarten 5a, D-91054 Erlangen, Germany

a r t i c l e i n f o

Article history:Received 29 April 2008Received in revised form 3 March 2009Accepted 13 March 2009

Keywords:Shackleton RangeEast AntarcticaZircon and monazite U–Pb and Th–U–PbdatingGondwana assembly

a b s t r a c t

A comprehensive set of new geochronological data from different parts of the Shackleton Range in EastAntarctica, comprising U–Pb single zircon and Th–U–Pb single and multi-grain monazite data, combinedwith published results, reveal a complex tectono-thermal history of the Shackleton Range. Three dis-tinct, spatially separated terranes or units with different magmatic and metamorphic history are nowrecognised: (i) the Southern Terrane (Unit I) contains detrital components as old as 2850 Ma, experiencedmagmatism between 1850 Ma and 1810 Ma and underwent a medium- to high-grade metamorphic eventat 1710–1680 Ma and, locally, again at 510 Ma; (ii) the Eastern Terrane (Unit II) occurs in the easternmostpart of the Shackleton Range and contains c. 1060 Ma old Grenvillian granitoids, which experienced meta-morphism at c. 600 Ma; and (iii) the Northern Terrane (Unit III) is characterised by 530 Ma old granitesand diorites, which are hosted within paragneisses as well as mafic and ultramafic rocks. All rocks ofUnit III experienced upper amphibolite- to granulite-facies and, locally, eclogite-facies metamorphism at510–500 Ma.

The geologic features of Palaeoproterozoic tectonism in the Southern Terrane are very similar to thoseof the Australo-Antarctic Mawson Continent. This may indicate that the Mawson Continent extends acrossthe East Antarctic Shield into the Shackleton Range. The 1060 Ma and 600 Ma events in the Eastern Ter-rane have not been documented for any part of the Shackleton Range before and are correlated withGrenvillian and Pan-African tectonism in Dronning Maud Land. By implication, this suggests that the Pan-African Mozambique/Maud Belt continues into the Shackleton Range. The associated suture is located inthe easternmost Shackleton Range and is related to the amalgamation of the Indo-Antarctic plate withWest Gondwana. This was followed by further collision of the combined Indo-Antarctic/West Gondwananblock with East Gondwana at approximately 510 Ma in the Northern Terrane. A suture related to this lattercollision can be traced in the Northern Shackleton Range and may continue northwards to the Sør Ron-dane Mountains and the Lützow Holm Bay area. Our data support the model that East Antarctica finallyassembled during the Pan-African orogeny, rather than during earlier Mesoproterozoic events.

© 2009 Elsevier B.V. All rights reserved.

1. Introduction

In the last decade, the Palaeoproterozoic to Cambrian evolu-tion of the East Antarctic Shield has become a major focus ofattention with respect to supercontinent formation during Meso-proterozoic and late Neoproterozoic/Cambrian times (e.g. Rogers,1996; Fitzsimons, 2000, 2003; Harley, 2003; Yoshida et al., 2003;Boger and Miller, 2004; Collins and Pisarevsky, 2005). The recogni-

∗ Corresponding author. Fax: +49 931 8884620.E-mail address: [email protected] (T.M. Will).

tion of Pan-African tectonism in orogenic belts that were consideredpreviously as purely Mesoproterozoic (“Grenvillian”) across EastAntarctica led several authors (e.g. Shiraishi et al., 1994; Meert et al.,1995; Hensen and Zhou, 1997; Boger et al., 2001; Fitzsimons, 2003;Kelsey et al., 2008) to suggest that East-Gondwana did not existas an entity during the Proterozoic, but became itself assembledduring the Pan-African orogeny.

The Shackleton Range in East Antarctica contains various base-ment units (e.g. Tessensohn et al., 1999) and could providefurther information towards the question of supercontinent for-mation during the Meso- and/or Neoproterozoic/Cambrian. Sparsegeochronological data as summarised below point to a protracted

0301-9268/$ – see front matter © 2009 Elsevier B.V. All rights reserved.doi:10.1016/j.precamres.2009.03.008


26 T.M. Will et al. / Precambrian Research 172 (2009) 25–45

tectonothermal history of the Shackleton Range from the Palaeo-proterozoic to the Cambrian. Recently, Schmädicke and Will (2006)and Romer et al. (2009) inferred a suture of Pan-African age in thenorthern Shackleton Range by demonstrating subduction-related,high-pressure eclogite-facies metamorphism of alpine-type ultra-mafic and associated mafic rocks in that area. As such ultramaficrocks are indicative of orogenic sutures (e.g. Coleman, 1971), thisfinding is exceptionally important in the context of the Pan-Africanamalgamation of Gondwana. Consequently, constraining the ageand tectonic evolution of the various basement units in the Shack-leton Range becomes particularly critical and this constitutes theoverall aim of this study.

Towards this aim, we present results from a geochronologi-cal study of zircon and monazite grains from high-grade ortho-and paragneisses from the Shackleton Range using the U–Pb andTh–U–Pb isotope systems. In order to establish the timing ofthe various tectonothermal events, isotope dilution U–Pb ther-mal ionisation mass spectrometry (ID-TIMS), U–Pb laser ablationsector field inductively coupled mass spectrometry (LA-SF-ICP-MS) zircon age dating and in situ Th–U–Pb electron microprobe(EMP) monazite age dating were carried out. Zircon and monaziteare particularly useful minerals for providing geochronologicalinformation about high-grade rocks because the radiogenic Pb isbelieved to be retained in the minerals even under granulite-faciesmetamorphic conditions (e.g. Copeland et al., 1988; Cherniak et al.,2004).

The overall goal is to better assess the tectono-metamorphicevolution of the Shackleton Range, i.e. to distinguish betweenPalaeo-, Meso- and Neoproterozoic/Cambrian crustal components,and, in the context of the Gondwana assembly, to compare theseresults with other terranes of the East Antarctic Craton and beyond.

2. Geological setting of the Shackleton Range

The Shackleton Range is situated between 80◦ and 81◦S andforms an elongated east-west trending mountain belt of some240 km length and 50–90 km width (Fig. 1). It is located at thenorthwestern edge of the stable East Antarctic Craton and is bor-dered by the N-S trending Transantarctic Mountains (TM) in thewest. The TM formed during the Pan-African Ross Orogeny at about500 Ma and trace the ancient Pacific margin of the East AntarcticCraton (Kleinschmidt and Tessensohn, 1987). Despite the fact thatparts of the Shackleton Range were also affected by Pan-Africantectonometamorphism (e.g. Zeh et al., 1999; 2004) it appears to bedifferent because its structural trend lies at almost right angles tothat of the TM (Clarkson, 1982). Furthermore, orogenic granitoids,which are typical of Andean-type orogens and ubiquitous in the TMare rare in the Shackleton Range. Many workers regard the Shackle-ton Range as a thin-skin thrust- and nappe-type collisional orogen(e.g. Tessensohn, 1997; Talarico et al., 1999; Zeh et al., 1999) thatwas caused by oblique, sinistral collision between the East Antarc-tic and Kalahari Cratons and led to the closure of the MozambiqueOcean (Tessensohn et al., 1999).

The Shackleton Range consists of a southern and northern beltof mostly metamorphic basement complexes, which are separatedby very low- to low-grade sedimentary rocks of the allochthonousMt. Wegener Formation (e.g. Clarkson, 1982; Buggisch et al., 1994).

The southern belt is exposed in the Read Mountains and com-prises medium- to high-grade metamorphic rocks of the ReadGroup, mainly of quartzitic, basic, calcareous and pelitic compo-sitions. These rocks are partly migmatised, in places interlayeredwith gneissic granites and locally intruded by granites and minorbasic rocks (Talarico and Kroner, 1999). Available geochronologi-cal data are restricted to a few Rb–Sr whole-rock isochron ages of1763 ± 32 Ma and 1599 ± 38 Ma from metagranites (Pankhurst et al.,1983) and Rb–Sr and K-Ar mineral cooling ages of 1650–1550 Ma

(Hofmann et al., 1980; Pankhurst et al., 1983). The older Rb–Srisochron age was interpreted to reflect the time of igneous activity,the younger age as that of metamorphism. A metagranite with alow 87Sr/86Sr initial of 0.704 was interpreted to indicate a juvenile,depleted mantle source for the metagranite precursor (Pankhurstet al., 1983).

The northern belt extends from the Northern Haskard High-lands in the west to the Pioneers Escarpment in the east (Fig. 1)and, based on lithostratigraphy, structural style and metamorphicgrade, was subdivided into three units: the Pioneers and Strat-ton Groups and an ophiolite complex (e.g. Clarkson et al., 1995;Talarico et al., 1999). The latter was interpreted as a relic of theMozambique Ocean by Tessensohn et al. (1999). The PioneersGroup comprises Meso- to Neoproterozoic, high-grade metamor-phic supracrustal cover rocks of pelitic, quartzitic, calcareous, basicand, subordinate, ultramafic compositions, whereas the StrattonGroup represents reworked Archaean to Mesoproterozoic base-ment that consists of medium- to high-grade, intermediate tofelsic gneisses, migmatites, minor granitoids and ultramafic rocks.According to Clarkson et al. (1995), the Stratton Group rocks arelithologically similar to those of the Read Group. The ophiolitecomplex consists of amphibolites, metagabbros and serpentinisedrocks, which are predominately exposed in the Herbert Mountains(Talarico et al., 1999; see Fig. 1). However, this unit may continuetowards the Haskard Highlands in the west. This is indicated bythe occurrence of ultramafic and associated mafic rocks in thatarea, which, as recently demonstrated by Schmädicke and Will(2006), underwent a subduction-related eclogite-facies metamor-phic event (20–23 kbar and 710–810 ◦C, corresponding to a peakmetamorphic gradient of∼11 ◦C/km). Unmetamorphosed sedimen-tary rocks of the Blaiklock Glacier group unconformably overlie thevarious basement rock units of the northern belt.

Geochronological data exist only for a few samples from widelyscattered locations (open squares in Figs. 1 and 2). U–Pb zir-con data of 2328 ± 47 Ma, 1715 ± 6 Ma and a Sm–Nd garnet-wholerock age of 535 ± 22 Ma for a granitic gneiss from the La GrangeNunataks in the northwestern part of the Shackleton Range (Fig. 1)were interpreted to date the emplacement of the gneiss precur-sor, subsequent migmatisation and a Pan-African metamorphicevent, respectively (Brommer et al., 1999). In the Northern HaskardHighlands, the Pioneers and Stratton Groups are separated by theWSW-trending Northern Haskard Fault (Fig. 2), a ductile to brit-tle, sinistral shear zone (Zeh et al., 1999). Geochronological studiesby these authors revealed that a granulite-facies sample from theMt. Weston Gneiss, south of the Northern Haskard Fault, wasemplaced at 1810 ± 2 Ma (U–Pb zircon) and experienced peak meta-morphism prior to 1670 ± 60 Ma (Sm–Nd garnet-whole rock). Incontrast, Pioneers Group rocks from the same area, but to the northof the Northern Haskard Fault, were intruded by dioritic magmas at532 ± 18 Ma (U–Pb zircon) and experienced a high-grade metamor-phic overprint during the Pan-African orogeny at c. 510 Ma (Sm–Ndgarnet: 506 ± 6 Ma and Rb–Sr biotite cooling age: 500 ± 10 Ma; Zehet al., 1999). Similar ages of 500 ± 10 Ma (K-Ar mica and amphi-bole) were determined by Talarico et al. (1999) and Brommerand Henjes-Kunst (1999) for rocks from the Bernhardi Hights andMount Sheffield areas some 100 km to the east/northeast of theHaskard Highlands (Fig. 1). U–Pb, Sm–Nd and Rb–Sr data of Zeh etal. (2004) reveal that rocks exposed on the Pioneers Escarpment(Fig. 1) experienced either a sole Pan-African metamorphic eventat c. 515–500 Ma (Lord Nunatak) or two metamorphic overprints atc. 1700 Ma and 500 Ma (Meade Nunatak).

In the present study, the focus was on detailed analyses of indi-vidual zircon and monazite grains or domains from a carefullyselected suite of ortho- and paragneiss samples. The geochrono-logical data were obtained by three different methods. First, highprecision zircon and monazite age data were obtained by isotope


T.M. Will et al. / Precambrian Research 172 (2009) 25–45 27

Fig. 1. Tectonic sketch map of the Shackleton Range, modified after Tessensohn et al. (1999) and Clarkson et al. (1995). The black patches show the location of nunataks.Circles and diamonds indicate the locations of U–Pb (LA-SF-ICP-MS zircon and ID-TIMS zircon and monazite) isotope and Th-U–Pb (EMP monazite) samples investigatedin this study; the squares indicate the sample locations of previously dated rocks. BH-Bernhardi Hights, BN-Baines Nunatak, DTN-Du Toit Nunatak, FN-Fairfield Nunatak,HP-Hatch Plain, LGN-La Grange Nunatak, LN-Lord Nunatak, MN-Meade Nunatak, NHH-Northern Haskard Highlands, MS-Mount Sheffield, SC-Shaler Cliff. Abbreviations ininset: DML-Dronning Maud Land, EAC-East Antarctic Craton, NVL-North Victorialand, TM-Transantarctic Mountains.

dilution thermal ionisation mass spectrometry (ID-TIMS) on twometapelitic rocks: one from the Pioneers Group north of the North-ern Haskard Fault (sample N4-3a), the other from the ophiolitecomplex at the Shaler Cliff in the Herbert Mountains (sample SC).Second, zircon grains from eleven samples were analysed by laserablation sector field inductively coupled mass spectrometry (LA-SF-ICP-MS) using heavy mineral concentrates obtained from ortho-and paragneisses from several parts of the Shackleton Range. Twosamples are from the Read Mountains (samples 167, 173), four fromthe Northern Haskard Highlands (samples G10, HM1, N4-3, 333),one from the central Pioneers Escarpment (sample Me3-3) andfour from the easternmost basement exposures on the PioneersEscarpment (samples 520, 525, 593, 612), henceforth referred toas the Eastern Basement. Third, monazite grains in 14 thin sectionswere analysed in-situ by electron microprobe (EMP). The ID-TIMSand LA-SF-ICP-MS samples investigated are described in Table 1,the EMP samples in Table 2. The sample localities are indicated inFigs. 1 and 2, and the analytical procedures are described in detailin Appendices A–C.

3. Results

3.1. ID-TIMS zircon and monazite dating

For ID-TIMS dating (for method see Appendix A) zircon and mon-azite grains were separated from heavy mineral concentrates of two

metapelitic rocks: one from a high-grade K-feldspar-kyanite-garnetgneiss of the Pioneers Group from the Northern Haskard Highlands(sample N4-3a), the other from a cordierite-bearing paragneissfrom the Shaler Cliff in the Herbert Mountains (sample SC). Theresults are summarised in Supplementary Table S1. As shown inFig. 3, all monazite and zircon analyses from sample N4-3a lie ona discordia with an upper intercept at 509 ± 4 Ma, those of sampleSC define an upper intercept age of 510 ± 4 Ma. We interpret bothages to define the time of peak metamorphism that affected theserocks.

3.2. LA-SF-ICP-MS zircon dating

Zircon grains from eleven samples (Table 1) were analysed bylaser ablation sector field inductively coupled mass spectrome-try (for analytical details see Appendix B) using heavy mineralconcentrates of five orthogneiss samples (HM1, N4-3, 165, 333,593) and six paragneisses samples (G10, Me3-3, 173, 520, 525,612) from different parts of the Shackleton Range (Figs. 1 and 2).The results are summarised in Supplementary Table S2. Many zir-con grains from the orthogneiss samples reveal oscillatory zoningpatterns (Fig. 4) and most grains have Th/U ratios well above 0.1(Table 1), which, according to Rubatto (2002) points to a magmaticorigin of the zircon grains. In contrast, zircon grains from parag-neisses are either structureless-diffuse (samples G10, 173, 525,Me3-3; Fig. 5a–d) or show core-rim relationships, with magmatic



Fig. 2. Sketch map of the Northern Haskard Highlands. White circles and diamonds-locations of U–Pb (zircon) isotope and Th-U–Pb (monazite) samples investigated in thisstudy; squares-locations of samples investigated by Zeh et al. (1999). NHF-Northern Haskard Fault.

cores being surrounded by metamorphic overgrowth rims (sam-ples 520, 612; Fig. 5e and f). Invariably, the structureless-diffusezircon grains and the metamorphic overgrowth rims have Th/Uratios of less than 0.1 (Table 1) as is common in many zircon grainsthat crystallised or recrystallised during metamorphism (Rubatto,2002).

3.2.1. OrthogneissesWithin error, granitic and dioritic orthogneisses from the

Haskard Highland (north of the Northern Haskard Fault,Figs. 1 and 2) yielded identical concordant ages of 534 ± 6 Ma(metadiorite N4-3), 527 ± 4 Ma (metadiorite HM1) and 530 ± 5 Ma

(granite gneiss 333), which is shown in Fig. 6a–c. An identical U–PbTIMS age of 532 ± 18 Ma was obtained by Zeh et al. (1999) for ametadiorite from the same area. Clearly, these ages indicate a pulseof Cambrian dioritic and granitic magmatism. In contrast to theHaskard Highlands, magmatic zircon grains from a metagranitefrom the Read Mountains to the south (sample 165) gave a mucholder concordant age of 1850 ± 13 Ma (Fig. 6d), and a granitic gneiss(sample 593) from the Eastern Basement yielded a concordant ageof 1059 ± 9 Ma (Fig. 6e). Some of these late Mesoproterozoic zircongrains have dark, structureless overgrowth rims (Fig. 4e) with verylow Th/U ratios (Table 1). These rims yielded a Neoproterozoicconcordant age of 602 ± 11 Ma (Fig. 6e), which indicates a post-

Fig. 3. 207Pb/235U vs. 206Pb/238U diagrams for U–Pb ID-TIMS dating of zircon (zrc) and monazite (mzt) in samples N4-3a (a) and SC (b). The data point error ellipses are 2�.

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Table 1Sample location, description and results of U–Pb LA-SF-ICP-MS single zircon and ID-TIMS (zircon and monazite) samples.

Sample Location Mineralassemblage

Metamorphic facies andpeak conditions

Zircon characteristics (withreference to Figs. 4 and 5)

# of analyseson # of grains

U [ppm]Th [ppm]

Th/U ratio Concordia age (# ofanalysis, MSWD)

Upper interceptage (# of analysis,MSWD)

Interpretation

LA-SF-ICP-MS samplesN4-3 Metadioritic

orthogneissNHHn grt, bi, cpx, cam, AF to GF Prismatic to rounded, 100–300 �m

(dia.)28 30–470 0.02–2.26 534 ± 6 Ma

(6, 0.23)— Age of magmatism

80◦24′16′′S plag, q Up to 700 �m in length 12 2–930 Mostly < 0.1 and1.3–2.3

29◦40′23′′W Very variable zoning patterns (Fig. 4a)

HM1 Metadioriticorthogneiss

NHHn grt, bi, cpx, cam, AF to GF Clear, prismatic, 150–400 �m, 20 70–300 0.07–0.77 527 ± 4 Ma(13, 0.057)

— Age of magmatism

80◦24′24′′S plag, q 710 ± 50 ◦C/9.0 ± 1.5 kbar osc. z., some grains with overgrowth 11 18–164 (mostly0.3–0.45)

29◦32′25′′W (Zeh et al., 1999) rims, some with convoluted zoning(Fig. 4b)

333 grt-biorthogneiss

NHHn grt, ksp, plag, bi, q AF to GF Clear, prismatic, 50–100 �m, 11 130–350 0.3–0.77 530 ± 5 Ma(7, 0.22)


80◦24′24′′S osc. z., few grains with dark, 9 61–25729◦32′25′′W structureless rims (Fig. 4c)

165 grt-bearingmetagranite

RM grt, ksp, plag, bi, AF to GF Clear, idiomorphic, 50–150 �m, 18 200–500 0.3–0.6 and 1850 ± 13 Ma(7, 0.005)

1859 ± 19 Ma(17, 10.1)

Age of magmatism

80◦43′46′′S q, opx relics 600 ± 110 ◦C/4.2 ± 1.6 kbar cores: osc. z., unzoned, structureless 18 25–153 0.06–0.1325◦46′09′′W (Schubert and Will,

1994)rims: narrow, structureless, dark(Fig. 4d)

593 bi gneiss EB bi, mu, plag, q ± grt AF Clear, prismatic, 100–250 �m, 19 200–820 0.2–1.0 (cores) 1059 ± 9 Ma(4, 0.0052)

1057 ± 12 Ma(13, 1.5)

Age of magmatism

80◦30′38′′S osc. z., few grains with dark, 12 22–477 0.02 (rims) 602 ± 11 Ma(2, 0.0052)

— Age ofmetamorphism

19◦12′28′′W structureless rims (Fig. 4e)

G10 ky-st-grtparagneiss

NHHs grt, st, ky, bi, mu, AF Clear, elliptical-rounded, 30–80 �m(dia.),

64 75–430 0.03–0.2 — 1698 ± 7 Ma(64, 2.0)

Age ofmetamorphism

80◦26′42′′S plag, q 580 ◦C/6 kbar (min.) structureless, diffuse, 64 7–63 average 0.1229◦27′33′′W (Zeh et al., 1999) patchy, irregular zoning (Fig. 5a)

172 crd-sill-grtparagneiss

RM grt, sill, crd, ksp GF Xenomorphic, rounded, 50–100 �m(dia.),

13 800–1200 0.05–0.17 1695 ± 12 Ma(1)

1704 ± 57 Ma(13, 32)

Age ofmetamorphism

80◦43′32′′S plag, bi, sp, q 790 ± 50 ◦C/5.5 ± 1.6 kbar structureless, diffuse (Fig. 5b) 12 36–250 average 0.0825◦44′53′′W (own unpubl. data)

525 grt-sill/ky-kspparagneiss

EB grt, sill, ky, ksp GF Xenomorphic, rounded, 50–100 �m(dia.),

37 150–300 mainly < 0.1 1690 ± 4 Ma(26, 0.037)


80◦39′34′′S bi, q structureless, diffuse, 34 4–39 average 0.0719◦47′39′′W irregular, patchy zoning (Fig. 5c)

Me3-3 ky-st-grtparagneiss

PE grt, st, ky, bi, mu, polymetamorph AF andGF

Xenomorphic, rounded, 20–50 �m(dia.),

66 100–300 Mainly < 0.1 1687 ± 7 Ma(9, 0.059)

1687 ± 7 Ma(17, 0.71)

Age of HT/LPgranulite-faciesmetamorphism

80◦22′11′′S sill, plag, crd, q AF: 663 ± 22 ◦C/ structureless, diffuse (Fig. 5d) 55 1–49 average 0.05 ∼500 Ma (Zehet al., 2004)

Age of MT/MPamphibolite-faciesmetamorphism

21◦57′38′′W 6.8 ± 1.0 kbar (Zeh et al.,2004)(ky-st-grt assemblages)GF: high-temperature,low-pressuremetamorphism(grt-sill-crdassemblages)

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Table 1 (Continued )

Sample Location Mineralassemblage

Metamorphic facies andpeak conditions

Zircon characteristics (withreference to Figs. 4 and 5)

# of analyseson # of grains

U [ppm]Th [ppm]

Th/U ratio Concordia age (# ofanalysis, MSWD)

Upper interceptage (# of analysis,MSWD)

Interpretation

520 bi gneiss EB plag, ksp, bi, q GF? Idio- to xenomorphic and 49 50–350 0.11–1.61(cores)

1907 ± 12 Ma(5, 0.08;magmaticcore)

— Age ofmagmatism/detrital

80◦39′03′′S rounded, 80–300 �m (dia. and/orlength),

25 3–190 1836 ± 12 Ma(4, 0.03;magmaticcore)


20◦33′12′′W very variable internal structures(Fig. 5e)

0.02–1.11(rims)

1707 ± 5 Ma(25, 1.70;metamorphicrim)


612 grt-stparagneiss

EB grt, st, sill(fibrolite),

AF Clear, xenomorphic, rounded, 45 65–330(U cores)

0.4–0.75(cores)

— 1834 ± 9 Ma(9, 0.084)

Age of magmatism

80◦32′14′′S bi, mu, q 50–150 �m (dia.), 29 120–190(U rims)

— 1771 ± 9 Ma(7, 0.13)

Age probablygeologicallymeaningless (seetext)

20◦34′47′′W cores: mainly idiomorphic withosc. z.,

11–249(Th cores)

0.01–0.24(rims)

1702 ± 8 Ma(7, 0.19)

1703 ± 7 Ma(13, 1.3)

Age ofmetamorphism

rims: grey, structureless (Fig. 5f) 2–94(Th rims)

Sample Location Mineral assemblage Metamorphic facies and peak conditions Upper intercept age (# of analysis, MSWD) Interpretation

ID-TIMS samplesN4-3a ky-ksp-grt paragneiss NHHn ky, ksp, sill, grt, st, GF 509 ± 4 Ma (5, 1.8) Age of metamorphism

80◦24′16′′S plag, q29◦40′23′′W

SC crd paragneiss HM crd, sill, bi, AF 510 ± 4 (4, 0.65) Age of metamorphism80◦16′46′′S plag, q 660 ◦C/7 kbar25◦28′58′′W (Zeh, 2001)

General abbreviations: NHHn-Northern Haskard Highlands (north of Northern Haskard Fault), NHHs-Northern Haskard Highlands (south of Northern Haskard Fault), PE-Pioneers Escarpment, HM-Herbert Mountains, RM-ReadMountains, EB-Eastern Basement. AF-Amphibolite Facies, GF-Granulite Facies, dia.-diameter, osc. z.-oscillatory zoning. Mineral abbreviations: grt-garnet, ksp-K-feldspar, plag-plagioclase, bi-biotite, q-quartz, opx-orthopyroxene,cpx-clinopyroxene, cam-Ca-amphibole, ky-kyanite, sill-sillimanite, crd-cordierite, st-staurolite, mu-muscovite, sp-spinel.

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Table 2Sample location, description and results of in-situ Th-U–Pb EMP monazite samples.

Sample Location Mineral assemblage Metamorphic facies Monazite characteristics (with reference toFig. 11)

# of analyses Weighted averageage ± 2� (Ma)

ThO*2 vs PbOage (Ma)

167 crd-sill-grt RM grt, sill, ksp, bi, GF Xenomorphic, 40–80 �m, 15 1672 ± 10 Ma 1668 Maparagneiss 80◦43′30′′S q 800 ± 93 ◦C/6.3 ± 0.9 kbar few grains are rounded,

25◦43′24′′W (own unpubl. data) in matrix (Fig. 11a)

173 crd-sill-grt RM grt, sill, crd, ksp, GFparagneiss 80◦43′32′′S sp, bi, plag, q 731 ± 120 ◦C/6.4 ± 1.1 kbar Angular to rounded, 30–80 �m, 28 1667 ± 5 Ma 1664 Ma

25◦44′53′′W (Schubert and Will, 1994) a few grains with optical zoning,matrix and inclusion in grt (Fig. 11b)

227 grt-bearing NHHn grt, ksp, plag, bi, AF Xenomorphic to rounded, 40–75 �m, 3 496 ± 20 Ma 494 Mametagranite 80◦24′17′′S mu, q Locally frayed grain boundaries,

29◦31′23′′W in matrix

308 ky-grt-ksp NHHn grt, ky, ksp, bi, AF to GF Xenomorphic, 40–180 �m, 7 499 ± 29 Ma 499 Magneiss 80◦23′08′′S plag, cam, q in matrix (Fig. 11 g)

29◦40′00′′W

315 grt-ksp NHHn grt, ksp, plag, bi, AF to GF Xenomorphic to rounded, 40–200 �m, 4 499 ± 26 Ma 499 Magneiss 80◦23′00′′S q locally frayed grain boundaries,

29◦40′41′′W in matrix and inclusion in grt (Fig. 11 h)

519 sill-ksp-grt EB grt, sill, ksp, bi, GF Xenomorphic to rounded, 60–300 �m, 43 1644 ± 4 Ma 1633 Maparagneiss 80◦38′18′′S mu, q Locally frayed grain boundaries,

20◦38′15′′W in matrix, as euhedral to xenomorphicinclusion in strongly retrogressed grt(Fig. 11c)

524 sill-ksp-grt EB grt, sill, crd, ksp, GF Xenomorphic to rounded, 20–200 �m, 32 1576 ± 15 Ma 1571 Maparagneiss 80◦39′34′′S bi, plag, q in matrix and at grt rim (Fig. 11d)

19◦47′39′′W

547 grt-bearing PE grt, mu, bi, q ? Xenomorphic, 50–80 �m,quarzitic gneiss 80◦27′18′′S in matrix, optically homogeneous 3 487 ± 42 Ma 494 Ma

24◦48′03′′W

564 ky-st-grt PE grt, st, ky, sill, AF Xenomorphic, 30–170 �m, 9 508 ± 15 Ma 501 Maparagneiss 80◦20′39′′S bi, mu, plag, ksp, q 665 ± 17 ◦C/8.7 ± 1.2 kbar in matrix

24◦03′19′′W (Zeh et al., 2004)

570 ky-grt PE grt, ky, sill, AF Xenomorphic, 15–25 �m, 2 486 ± 63 Ma 485 Maparagneiss 80◦19′17′′S bi, mu, plag, q in matrix

24◦00′07′′W

680 ky-st-grt PE grt, ky, st, bi AF Xenomorphic, 30–80 �m, 7 493 ± 16 Ma 485 Maparagneiss 80◦22′18′′S plag, q 660 ± 40 ◦C/6.2 ± 1.5 kbar in matrix

21◦58′09′′W (Schubert and Will, 1994)663 ± 22 ◦C/6.8 ± 1.0 kbar(Zeh et al., 2004)

683 ky-st-grt PE grt, ky, sill, st, GF and AF Xenomorphic to angular, 90–180 �m, 11 (matrix) 489 ± 21 Ma 487 Magneiss 80◦22′18′′S bi, plag, q 660 ± 40 ◦C/6.2 ± 1.5 kbar in matrix and as inclusion in grt 3 (incl.) 1596 ± 39 1593 Ma

21◦58′09′′W (Schubert and Will, 1994) (Fig. 11e and f)663 ± 22 ◦C/6.8 ± 1.0 kbar(Zeh et al., 2004)



Tabl

e2

(Con

tinu

ed)

Sam

ple

Loca

tion

Min

eral

asse

mbl

age

Met

amor

ph

icfa

cies

Mon

azit

ech

arac

teri

stic

s(w

ith

refe

ren

ceto

Fig.

11)

#of

anal

yses

Wei

ghte

dav

erag

eag

e±

2�(M

a)Th

O* 2

vsPb

Oag

e(M

a)

684

ky-s

t-gr

tPE

grt,

ky,s

ill,

st,

GF

and

AF

Xen

omor

ph

icto

rou

nd

ed,3

0–90

�m

,9

489

±23

Ma

487

Ma

par

agn

eiss

80◦ 2

2′ 18′′ S

bi,p

lag,

q66

0±

40◦ C

/6.2

±1.

5kb

arin

mat

rix,

opti

cally

hom

ogen

eou

s21

◦ 58′ 0

9′′ W(S

chu

bert

and

Wil

l,19

94)

663

±22

◦ C/6

.8±

1.0

kbar

(Zeh

etal

.,20

04)

692

st-g

rtPE

grt,

st,b

i,A

FX

enom

orp

hic

toro

un

ded

,30–

60�

m,

54

85±

24M

a4

85M

ap

arag

nei

ss80

◦ 27′ 3

5′′ Sm

u,p

lag,

qin

mat

rix,

opti

cally

hom

ogen

eou

s22

◦ 28′ 4

8′′ W

ThO

2*

isth

esu

mof

ThO

2m

easu

red

and

the

ThO

2eq

uiv

alen

tto

the

UO

2co

nce

ntr

atio

nm

easu

red

.Th

ew

eigh

ted

aver

age

age

ofth

eap

par

ent

ind

ivid

ual

spot

ages

(wh

ich

wer

ed

eter

min

edac

cord

ing

toM

onte

let

al.,

1996

)w

ere

calc

ula

ted

usi

ng

Isop

lot/

Exce

l2.4

9(L

ud

wig

,20

01)

and

the

ThO

2*

vers

us

PbO

ages

foll

owin

gth

ep

roce

du

reou

tlin

edby

Suzu

kiet

al.(

1994

).G

ener

alab

brev

iati

ons:

NH

Hn

-Nor

ther

nH

aska

rdH

igh

lan

ds

(nor

thof

Nor

ther

nH

aska

rdFa

ult

),PE

-Pio

nee

rsEs

carp

men

t,R

M-R

ead

Mou

nta

ins,

EB-E

aste

rnB

asem

ent,

AF-

Am

ph

ibol

ite

Faci

es,G

F-G

ran

uli

teFa

cies

.Min

eral

abbr

evia

tion

s:gr

t-ga

rnet

,ksp

-K-f

eld

spar

,pla

g-p

lagi

ocla

se,b

i-bi

otit

e,q-

quar

tz,k

y-ky

anit

e,si

ll-s

illi

man

ite,

crd

-cor

die

rite

,st-

stau

roli

te,m

u-m

usc

ovit

e,sp

-sp

inel

.

Fig. 4. Cathodoluminescence images of zircon grains from orthogneisses from theShackleton Range: (a) sample N4-3, (b) sample HM1, (c) sample 333, (d) sample 165,(e) sample 593. Ages in Ma. Scale bar = 100 �m.

Grenvillian, Pan-African metamorphic overprint of this rock (seediscussion below).

3.2.2. ParagneissesFifty-six analyses of internally structureless zircon grains

(Fig. 5a) from the staurolite-garnet gneiss sample G10 from theNorthern Haskard Highland (south of the Northern Haskard Fault;Fig. 2) yielded an upper intercept age of 1698 ± 7 Ma (Fig. 7a).The same age was obtained for samples from various locations:a cordierite-sillimanite-garnet gneiss (sample 172) from the ReadMountains gave an upper intercept age of 1704 ± 57 Ma anda concordia age (one grain only) of 1695 ± 12 Ma (Fig. 7b), agarnet-sillimanite-K-feldspar gneiss (sample 525) from the EasternBasement a concordia age of 1690 ± 4 Ma (Fig. 7c), and a gar-net gneiss (sample Me3-3) from the Pioneers Escarpment (MeadNunatak) a concordia age of 1687 ± 7 Ma (Fig. 7d). Very similarconcordant ages of 1707 ± 5 Ma and 1702 ± 8 Ma were obtainedfor metamorphic zircon overgrowth rims in a biotite gneiss (sam-ple 520; Fig. 8a) and a sillimanite (fibrolite)-garnet-staurolitegneiss (sample 612; Fig. 9a), respectively, from the Eastern Base-ment.

The ages obtained from the metamorphic zircon overgrowthsin samples 520 and 612 (Figs. 8a and 9a) as well as from thestructureless-diffuse zircon grains indicate clearly that these zir-cons domains formed during a metamorphic event at about1700 Ma. This is also consistent with the commonly low Th/Uratios in these domains (Table 1). The U–Pb zircon data from theparagneisses show that the central part of the Shackleton Range,extending from the Northern Haskard Highlands in the west toMeade Nunatak on the Pioneers Escarpment and the Eastern Base-ment in the east, was affected by a structural-metamorphic event atabout 1710–1680 Ma that overprinted older detrital zircon grains.This is well reflected by the zircon grains in the paragneiss samples520 and 612 from the Eastern Basement. Most of the zircon coresshow typical magmatic zoning patterns (Fig. 5e and f) and yieldearly Palaeoproterozoic to Archaean ages (Figs. 8b–d and 9b–d)that are significantly older than the 1710–1680 Ma metamorphic



Fig. 5. Cathodoluminescence images of zircon grains from paragneisses from the Shackleton Range: (a) sample G10, (b) sample 172, (c) sample 525, (d) sample Me3-3, (e)sample 520, (f) sample 612. Ages in Ma. Scale bar = 100 �m.

event. U–Pb analyses of zircon core domains from sample 520 gavea few Archaean ages of 2845 Ma, 2772 Ma and 2535 Ma (Fig. 8b)and Palaeoproterozoic ages of 2008 Ma, 1907 ± 12 Ma (Fig. 8c) and1836 ± 12 Ma (Fig. 8d). Zircon cores in sample 612 provided twoArchaean to Palaeoproterozoic ages of 2583 Ma and 2439 Ma andyounger ages of about 1900 Ma (2 grains; Fig. 9b), 1834 ± 9 Ma(Fig. 9c) and 1771 ± 9 Ma (Fig. 9d).

Several metamorphic zircon domains yielded Pb-Pb ages thatare younger than 1710–1680 Ma. Many zircon analyses of sam-ple Me3-3, which has a concordia age of 1687 ± 7 Ma, plot on adiscordia with an upper intercept at 1710 ± 24 Ma and a lowerintercept at 1602 ± 36 Ma (Fig. 7d). Additionally, Fig. 7d showsthat there are several discordant zircon analyses that pointto a multiple, post-1700 Ma Pb loss. This array indicates thatthese metamorphic zircon domains suffered Pb loss during thelate Palaeoproterozoic/early Mesoproterozoic. U–Pb zircon anal-yses of sample 172 plot on a discordia with an upper interceptage of 1704 ± 57 Ma (best constrained by a concordant zirconanalysis of 1695 ± 12 Ma) and a lower intercept at 407 ± 56 Ma(Fig. 7b). The large MSWD of 32 indicates that most of the zircondomains analysed must have experienced multiple Pb loss, proba-bly during the Pan-African orogeny at about 500 Ma and/or morerecently.

3.3. Th-U–Pb EMP monazite dating

In-situ analyses of ThO2, UO2 and PbO concentrations in mon-azite grains from 13 paragneisses (samples 167, 173, 308, 315, 519,524, 547, 564, 570, 680, 683, 684, 692) and one metagranite (sample227) from various parts of the Shackleton Range (Figs. 1 and 2) wereused to calculate ages following two different approaches. First,for each single analysis, an age was calculated using the equationsgiven by Montel et al. (1996). The error resulting from countingstatistics was typically in the order of ±20–±40 myr (1�). Usingthese apparent age data, weighted average ages for monazite pop-ulations in the samples were then calculated using Isoplot/Ex 2.49(Ludwig, 2001) and are interpreted as the age of monazite growth

or recrystallisation during metamorphism (Supplementary TableS3). Second, ‘chemical’ ages were determined using the ThO2*-PbOisochron method (CHIME) of Suzuki et al. (1994), where ThO2* isthe sum of measured ThO2 and ThO2 equivalent to the measuredUO2 (Supplementary Table S3 and Fig. 10). In all samples analysed,the model ages obtained by the two different methods agree excep-tionally well. The analytical conditions are described in AppendixC.

The granulite-facies paragneiss samples 167 and 173 from theRead Mountains yielded late Palaeoproterozoic weighted averageages of 1672 ± 10 Ma and 1667 ± 5 Ma, respectively (Fig. 11a andb; Table 2). Slightly younger ages of 1644 ± 4 Ma and 1576 ± 15 Mawere determined for the granulite-facies paragneiss samples 519and 524 from the Eastern Basement (Fig. 11c and d; Table 2). Asimilar age of 1596 ± 39 Ma was obtained from monazite inclu-sions in garnet in the paragneiss sample 683 (Fig. 11e; Table 2)from Meade Nunatak on the Pioneers Escarpment. The latter sam-ple experienced a polymetamorphic history as documented by aweighted average age of 489 ± 21 Ma for monazite grains in thematrix (Fig. 11f; Table 2). These results agree with age data pre-sented by Zeh et al. (2004), who also demonstrated that a paragneisssample from the Meade Nunatak experienced a polymetamorphichistory. Three other paragneisses from the same location (samples680, 684, 692) yielded identical Pan-African EMP monazite ages of493 ± 16 Ma, 489 ± 23 Ma and 485 ± 24 Ma, respectively, but con-tain no evidence for an older metamorphic event. All remainingsamples from the Pioneers Escarpment gave Pan-African ages rang-ing from 486 Ma to 508 Ma (Table 2 and Supplementary Table S3).These ages are identical to the 499 ± 26 Ma and 499 ± 29 Ma agesobtained on the paragneiss samples 308 and 315 from the northernHaskard Highlands (Fig. 11g and h; Table 2) and the metagranitesample 227, which yielded an age of 496 ± 20 Ma.

4. Interpretation and discussion of the results

The age data obtained in this study (420 single zircon U–PbICP-MS analyses and 181 single monazite in-situ Th-U–Pb analy-



Fig. 6. 207Pb/235U vs. 206Pb/238U diagrams for individual zircon domains in orthogneisses from the Shackleton Range: (a) sample N4-3, (b) sample HM1, (c) sample 333, (d)sample 165, (e) sample 593. The data point error ellipses are 2�. The non-shaded ellipses are not included in the age calculations.



Fig. 7. 207Pb/235U vs. 206Pb/238U diagrams for zircon grains in paragneisses from the Shackleton Range: (a) sample G10, (b) sample 172, (c) sample 525, (d) sample Me3-3. Thedata point error ellipses are 2�. The non-shaded ellipses are not included in the calculations.

Fig. 8. 207Pb/235U vs. 206Pb/238U diagrams for sample 520. (a) All data. The results obtained for metamorphic zircon overgrowth rims are indicated. (b) Apparent Pb-Pb agevs. probability diagram. (c, d) Data for magmatic cores (younger than 2000 Ma). The data point error ellipses are 2�.



Fig. 9. 207Pb/235U vs. 206Pb/238U diagrams for zircon grains in sample 612. (a) All data, with (b) apparent Pb-Pb vs. probability diagram; the results obtained for metamorphiczircon overgrowth rims are shown. (c, d) Data for magmatic cores (younger than 2000 Ma). The result shown in (d) might be geologically meaningless and could be the resultof Pb loss of the 1840 Ma old magmatic zircon domains during the 1700 Ma metamorphic event. The data point error ellipses are 2�.

ses) provide evidence that different parts of the Shackleton Rangewere affected by magmatic and metamorphic events at differ-ent times. Magmatism occurred at approximately 1850–1810 Ma,1060 Ma and 530 Ma and high-temperature metamorphism at1710–1680 Ma, 600 Ma and 510 Ma (Figs. 12–14). These Palaeo-,Meso- and late Neoproterozoic/Cambrian tectonothermal eventscompare well with the Kimban (e.g. Parker et al., 1993; Daly etal., 1998), Grenvillian and Pan-African orogenic events elsewhere.In addition, c. 2800 Ma and 2500–2400 Ma detrital zircon grainsoccur in several rocks from the Eastern Basement (Figs. 8b and9b) and indicate the presence of still older crustal componentsin the Shackleton Range. Based on our results and previouslypublished data three distinct geological units or terranes can bedistinguished.

The Southern Terrane (Unit I) extends for some 200 km in aneast-west direction and is exposed in the Read Mountains, theEastern Basement, the Haskard Highland (south of the North-ern Haskard Fault, Fig. 1) and Meade Nunatak on the PioneersEscarpment. Magmatic rocks intruded at about 1850–1810 Ma andwere affected by a tectonometamorphic overprint at 1710–1680 Ma.Rocks from Meade Nunatak record a later Pan-African metamor-phic overprint at c. 500 Ma, which is otherwise absent in Unit I.Palaeoproterozoic magmatic activities are well constrained by U–Pbintrusion ages of 1850 ± 13 Ma (sample 165 from the Read Moun-tains) and 1810 ± 2 Ma of the Mt. Weston Gneiss in the HaskardHighlands (Zeh et al., 1999). Most likely, the different emplace-ment ages reflect that Unit I experienced a protracted magmatichistory from 1850–1810 Ma. Prolonged magmatic activity is alsosupported by magmatic zircon cores that were found in the parag-neiss samples 520 and 612 from the Eastern Basement. The zirconcore domains in these gneisses provide evidence for magmaticactivities at about 1836 ± 12 Ma, 1834 ± 9 Ma and 1771±9 Ma (Figs.8d and 9c and d). The ages of these detrital zircon grains clearly indi-cate that the paragneiss precursors must have been deposited afterthe emplacement of the Palaeoproterozoic granitoids, but prior totheir metamorphic overprint at 1710–1680 Ma. Detrital zircon coreage data also constrain the presence of Archaean magmatic rocks,as old as 2850 Ma (Fig. 8b) in the provenance area of the Unit Iparagneisses. Rocks of this age are otherwise unknown from the

Shackleton Range itself and are most likely derived from the hin-terland, i.e. the East Antarctic Craton, prior to the 1710–1680 Maold tectonometamorphic event. So far the oldest reported age forthe Shackleton basement was 2328 ± 47 Ma (Brommer et al., 1999)and was determined for a gneiss from the La Grange Nunatak(Fig. 1).

All para- and orthogneisses of the Southern Terrane expe-rienced a high-temperature, low-pressure upper amphibolite-to granulite-facies metamorphism at 1710–1680 Ma. The meta-morphic conditions are reflected by the petrological record(Tables 1 and 2), and the age of metamorphism is constrained by theLA-SF-ICP-MS U–Pb zircon data from the Northern Haskard High-lands to the south of the Northern Haskard Fault (sample G10),the Read Mountains (sample 173) and the Eastern Basement (sam-ples 520, 525, 612). Furthermore, numerous EMP monazite agedata from various locations shown in Fig. 1 (paragneiss samples167, 173, 519) confirm this early Palaeoproterozoic metamorphicevent (Table 2, Supplementary Table S3 and Fig. 10). Additionally,this is supported by a U–Pb age of 1715 ± 6Ma for metamorphiczircon overgrowth rims in a gneiss from the La Grange Nunatak(Brommer et al., 1999) as well as U–Pb zircon and Sm–Nd garnet-whole rock data of 1665 ± 60 Ma from the Haskard Highlands (Zehet al., 1999). Taken together these ages indicate that the time spanfor the Palaeoproterozoic metamorphic event was less than 25 Ma.

Th–U–Pb EMP dating of monazite grains in sample 683 fromMeade Nunatak (Fig. 1) provides evidence for a polymetamor-phic history in this area, with high-temperature, low-pressuregranulite-facies metamorphism at c. 1700 Ma and a Barrovian-type amphibolite-facies event at c. 510 Ma, corroborating earlierresults by Zeh et al. (2004). As this is the only location on thePioneers Escarpment where a polymetamorphic history can bedemonstrated, it is impossible to assess the spatial extent of thearea that was affected by both Palaeoproterozoic Kimban- andPan-African-age metamorphism. For example, EMP monazite dat-ing of other rocks from Meade Nunatak (samples 680, 684, 692)yielded only Pan-African ages of 500 ± 10 Ma. Possibly, the earlyhigh-temperature sillimanite-cordierite-garnet assemblages mayhave been largely erased at many locations by the Pan-African meta-morphism that produced the garnet-staurolite-kyanite domains,



Fig. 10. ThO2* versus PbO isochron plots for monazite analyses obtained by in-situ EMP monazite dating. The age is determined by the slope of the regression line (Suzuki etal., 1994). Sample 227 is a metagranite, the remaining samples are paragneisses.

which are characteristic for the textures of the polymetamorphicrocks. Despite the fact that the Pan-African event dominates thepetrological record, the Palaeoproterozoic ID-TIMS data of Zeh et al.(2004) and our LA-SF-ICP-MS and EMP data (Figs. 7d and 11e) canbe taken as tentative evidence that Meade Nunatak is an extensionof the Southern Terrane.

Unit II or the Eastern Terrane is situated in the north-easternmost area of the Eastern Basement. Granitoid intrusionsoccurred at 1059 ± 9 Ma as constrained by the zircon LA-SF-ICP-MS analyses of sample 593 (Fig. 6e). Minor zircon overgrowthrims yielded an age of 602 ± 11Ma, indicating that these rockswere affected by a metamorphic event during the Neoproterozoic.This metamorphic overprint occurred some 100 myr earlier thanthe tectonothermal event that affected the Northern Terrane (seebelow) of the Shackleton Range at some 500 Ma. Grenvillian andsubsequent Neoproterozoic events have never been demonstratedfor the Shackleton Range. Our data provide therefore the first evi-dence that Grenvillian age tectonism occurred in the ShackletonRange.

The Northern Terrane or Unit III extends from the North-ern Haskard Highlands (north of the Northern Haskard Fault)to the Herbert Mountains and the Lord and Baines Nunatakson the Pioneers Escarpment (Fig. 1). This unit comprisesmetapelitic and metapsammitic rocks that are closely associatedwith mafic and ultramafic rocks. Magmatism occurred at about530 Ma, as reflected by the U–Pb zircon ages obtained fromthe metadiorite and metagranite samples N4-3, HM1 and 333(Fig. 6a–c) from the Northern Haskard Highlands. Subsequently,these granitoids and the associated country rocks experiencedan upper amphibolite- to granulite-facies metamorphic event(Tables 1 and 2), with a peak pressure of some 13 kbar (Schubertand Will, 1994). Alpine-type ultramafic rocks from the NorthernHaskard Highlands, however, provide unambiguous evidence forhigh-temperature, high-pressure eclogite-facies metamorphism of710–810 ◦C/20–23 kbar at the same time (Schmädicke and Will,2006). This was recently corroborated by Romer et al. (2009), whodetermined Sm–Nd garnet-whole rock ages of the ultramafic rocksof 525 ± 5 Ma and 520 ± 14 Ma.



Fig. 12. Time-space diagram for various geographical parts of the Shackleton Range. O-Ordovician, C-Cambrian. Mt. Prov.-Mount Provender area.

ID-TIMS U–Pb ages of 509 ± 1 Ma and 510 ± 1 Ma (Fig. 3)obtained on zircon and monazite grains in samples N4-3a andSC from the Northern Haskard Highlands and the Shaler Cliffin the Herbert Mountains, respectively, and several EMP mon-azite ages (samples 227, 308, 315 from the Northern HaskardHighlands and samples 547, 564 and 570 from the Freshfield,Lord and Baines Nunataks, respectively, Table 2, SupplementaryTable S3 and Fig. 10) constrain the high-grade metamorphicevent at 510–500 Ma. Additional support is provided by a Sm–Ndgarnet-whole rock age of 506 ± 6 Ma obtained for a metadior-ite from the Northern Haskard Highlands by Zeh et al. (1999),monazite U–Pb TIMS and Sm–Nd garnet-whole rock isochronages of 514 ± 1 Ma and 518 ± 5 for a paragneiss from the LordNunatak (Zeh et al., 2004). K-Ar and Rb–Sr cooling ages ofc. 500 Ma (Kleinschmidt and Buggisch, 1994; Zeh et al., 2004,respectively) and Ar–Ar plateau ages pf 500–490 Ma (own unpub-lished data) indicate that the rocks in the Northern Terrane,including the high-pressure eclogite-facies rocks, must have beenexhumed rapidly from depths of 60–70 km within a few mil-lion years after their peak metamorphism and deformation. K-Arages of c. 500 Ma from the Herbert Mountains (Brommer andHenjes-Kunst, 1999; Talarico et al., 1999) also conform to ourdata.

4.1. Correlation of the Shackleton Range basement provinces withother terranes

The Southern Terrane contains many geological features thatare typical of the Australo-Antarctic Mawson Continent, as initiallydefined by Fanning et al. (1996), which comprises the Gawler Cra-ton in South Australia, Terre Adélie and King George V Land in thesouth-eastern sector of the East Antarctic Craton and, most likely, asdiscussed by Goodge and Fanning (1999) and Goodge et al. (2001),the Miller Range in the central Transantarctic Mountains (Fig. 14).The 1850–1810 Ma magmatism in the Southern Terrane is matchedby the 1850 Ma Donning Suite Granitoids (e.g. Fanning et al., 2007),exposed at the east coast of the Eyre Peninsula in South Australia.Moreover, the Southern Terrane as well as the Mawson Continentexperienced Kimban-age tectonism with very similar high-grade,low-pressure metamorphism. This event caused partial melting ofgarnet-cordierite bearing metapelitic rocks in all of these Palaeo-proterozoic basement complexes and is dated at c. 1710–1680 Ma(e.g. this study; summary in Fitzsimons, 2003; Peucat et al., 1999;Di Vincenzo et al., 2007). Despite the fact that the Southern Terraneis separated by several 1000 km’s from the other exposed basementprovinces of the Mawson Continent, the similarities in the geolog-ical record between these terranes are striking. This was already

Fig. 11. Backscattered electron images (BSE) showing the textural position of monazite grains used for in-situ EMP dating: (a) sample 167, (b) sample 173, (c) sample 519, (d)sample 524, (e, f) sample 683, (g) sample 308, (h) sample 315. Samples 308 and 315 are HT amphibolite- to granulite-facies paragneisses, the remaining samples are HT/LPgranulite-facies paragneisses. In addition to the HT/LP metamorphism, sample 683 experienced a subsequent MT/MP metamorphic event. The weighted average age ± 2�(Ma) obtained for the respective domain is indicated; the age given in (b) applies to monazite grains in the matrix and the inclusion in garnet. Scale bar = 100 �m. For mineralabbreviations see Table 2, in addition: mzt-monazite, ilm-ilmenite.



Fig. 13. Map showing the distribution of crustal provinces and their age of tectonism in the Shackleton Range. The stippled lines indicate the possible locations of the sutures.LGN-La Grange Nunatak, LN-Lord Nunatuk, MN-Meade Nunatak.

noted by Tingey (1991) and Fitzsimons (2000, 2003), but could notbe substantiated because of the lack of reliable data for the Shackle-ton Range. Our new geochronological data may indeed corroborateearlier ideas and indicate that the Mawson Continent extends fromSouth Australia to Terre Adélie and King George V Land, the MillerRange in the central Transantarctic Mountains and further acrossthe entire East Antarctic Shield to the southern Shackleton Range(Fig. 14).

The 1060 Ma magmatic and 600 Ma metamorphic events doc-umented for the Eastern Terrane have not been reported for theShackleton Range before. However, such ages are known from sev-eral Antarctic crustal provinces such as Dronning Maud Land to thenorth of the Shackleton Range (e.g. Jacobs et al., 2003; Harley, 2003;Board et al., 2005; Bisnath et al., 2006), the Ellsworth-WhitmoreMountains (e.g. Flowerdew et al., 2007) and the Prydz Belt (e.g.Kelsey et al., 2008). It seems highly likely that the Eastern Terraneis the southward extension of Dronning Maud Land into the Shack-leton Range. The Maud Belt itself is considered to be the easternand southern extension of the Mozambique Belt into Antarctica(e.g. Jacobs et al., 1998; Boger et al., 2001; Fitzsimons, 2003; Meert,2003; Boger and Miller, 2004). Thus, by implication, the EasternTerrane would be the southernmost part of the Mozambique Beltin Antarctica (Fig. 14). Consequently, a ∼600 Ma suture related tothe amalgamation of the Indo-Antarctic plate with West Gondwana

(the Mozambique suture sensu Boger and Miller, 2004) shouldoccur somewhere in the easternmost Shackleton Range and maybe delineated by several mylonitic shear zones that are exposed inthat area.

The geological record of the Northern Terrane preservesextensive evidence for Pan-African magmatism, amphibolite-to granulite-facies metamorphism and deformation. Magmatismoccurred at ∼ 530 Ma, peak metamorphic conditions were reachedshortly afterwards at c. 510 Ma. The occurrence of eclogite-faciesalpine-type ultramafic rocks in the Northern Haskard Highlands(Schmädicke and Will, 2006; Romer et al., 2009) clearly tes-tifies to subduction of oceanic crust and collisional tectonicsduring the Pan-African orogeny. Most likely, dehydration melt-ing of the mantle wedge above the subducting slab producedascending melts that underplated the Palaeoproterozoic crust andcaused its partial melting at 530 Ma. Continued convergence causedcomplete subduction of the oceanic lithosphere and continent-continent collision at approximately 510 Ma, which produced thehigh-temperature, high-pressure rocks. These rocks and the ophi-olite relics found in the Herbert Mountains (Talarico et al., 1999),together with our new geochronological data, testify to the pres-ence of a 530–510 Ma old, roughly E-W trending suture in thenorthern Shackleton Range. Suturing of that age was referred toas Kuungan by Boger and Miller (2004) and the associated Kuunga

Fig. 14. Distribution of crustal provinces in East Gondwana as modified after Fitzsimons (2003) and Harley (2003). The location of the Tasman Line marking the easternmostextent of Precambrian rocks in Australia is taken from Cawood and Korsch (2008). The probable continuations of the orogenic sutures related to the assembly of the variousShackleton Range basement blocks are indicated as grey bands. CC-Curnamona Craton, GC-Gawler Craton, GhC-Grunehogna Craton, IC-Indian Craton, KC-Kalahari Craton. BH-Bunger Hills, cDML, wDML-central and western Dronning Maud Land, DG-Denmar Glacier, EG-Eastern Ghats, LHB-Lützow Holm Bay, MR-Miller Range, NP-Napier Complex,PB-Prydz Bay, PCM-Prince Charles Mountains, SR-Shackleton Range, SRM-Sør Rondane Mountains, TA-Terre Adélie and King George V Land.



suture, which is inferred to join and overprint the Mozambiquesuture somewhere in central Dronning Maud Land (Boger et al.,2001), represents the final collision site between the combinedIndo-Antarctic and West Gondwana plates (that were amalgamatedat ∼ 600 Ma) and East Gondwana (East Antarctic Craton and Aus-tralia) during the early Cambrian between 530 Ma and 510 Ma.Based on monazite EMP age dating, the presence of a suture ofthe same age was already inferred by Asami et al. (2005) forthe Sør Rondane Mountains (Fig. 14). As documented by theseauthors, granulite-facies rocks in the Sør Rondane Mountains andsurrounding areas record early Cambrian metamorphism. Contem-poraneous events were also demonstrated for high-grade rocksfrom the Lützow Holm Bay area to the east of the Sør RondaneMountains by Shiraishi et al. (1994). It is therefore conceivablethat the Neoproterozoic/Cambrian Kuunga suture continues fromthe Shackleton Range northwards to the Sør Rondane Mountainsand the Lützow Holm Bay area from which it continues to SriLanka and further into India (Fig. 14). In this scenario, the Kuungasuture must overprint the older Mozambique suture in Dron-ning Maud Land as suggested by Boger et al. (2001). However,as this idea is solely based on the grade and age of peak meta-morphism in these respective areas it should be treated withcaution.

5. Conclusion

New U–Pb zircon and Th-U–Pb monazite age data presentedin this study reveal that the Shackleton Range represents acomposite terrane that was affected to various extents by thePalaeoproterozoic Kimban orogeny, the Mesoproterozoic Grenvil-lian orogeny and the late Neoproterozoic/Cambrian Pan-Africanorogeny (Figs. 13 and 14). Separate crustal terranes could be delin-eated.

The Palaeoproterozoic Southern Terrane is interpreted to be thewesternmost extension of the Mawson Continent. If true, the spatialextent of the Australo-Antarctic Mawson Continent is in the orderof 4000 km, a distance that is similar to or even longer than thelength of the Transantarctic Mountains. Furthermore, the possibleextension of the Mawson Continent across the entire East Antarc-tic Craton has implications for the position of younger orogens. Forexample, the continuation of the Neoproterozoic Pinjarra or KuungaOrogen that extends from westernmost Australia to Prydz Bay(Boger et al., 2001) is debated. Fitzsimons (2000, 2003) discussedseveral possibilities how this orogen, which separates different Pro-terozoic basement provinces in East Antarctica, might continue intoAntarctica: (i) extension of the Pinjarra Orogen to an area northof the Shackleton Range, where it intersects the Mozambique Beltsomewhere in western Dronning Maud Land or, (ii) extension toa location between the Miller Range in the central TransantarcticMountains and the Shackleton Range. If our interpretation of a con-tinuous Mawson Continent across the East Antarctic Shield is truethe latter option must be ruled out and the first possibility seemsmore likely.

Two Pan-African sutures must be present in the ShackletonRange. The ∼600 Ma suturing affected the easternmost part of theRange, which is part of the Mozambique Belt that extends viaDronning Maud Land into East Africa. The younger, 530–510 MaKuunga suture in the northern Shackleton Range probably extendsto the Sør Rondane Mountains and the Lützow-Holm Bay area. Theolder Pan-African suturing event is related to the amalgamation ofWest Gondwana with the Indo-Antarctic plate, the younger eventcorresponds to the final assembly of Gondwana. Our results sup-port earlier investigations (e.g. Boger et al., 2001; Meert, 2003;Harley, 2003; Fitzsimons, 2003; Boger and Miller, 2004; Boardet al., 2005; Boger and Wilson, 2005) that East Antarctica wasnot a single coherent crustal block within East Gondwana or

Rodinia, but was assembled only during the Neoproterozoic to Cam-brian.

Acknowledgments

Financial support from the Deutsche Forschungsgemeinschaft(grant Fr 2183/1-1, 1-2) is gratefully acknowledged. W. Schubert isthanked for providing the samples, which he collected during theGerman GEISHA expedition in 1987/88. P. Späthe prepared superbthin sections, and U. Schüssler is thanked for his help with themicroprobe work at the University of Würzburg. V. v. Seckendorffand B. Schulz helped greatly with the monazite electron microprobedating at the University of Erlangen. J. Schastok is thanked for herassistance with the mineral separation at the University of Frank-furt. M. Fanning and G. Gibson are also thanked for their thoroughand helpful reviews as well as R. Parrish for editorial assistance.

Appendix A. ID-TIMS zircon and monazite analyses(Supplementary Table S1)

Zircon and monazite were prepared and analysed at the NERCIsotope Geosciences Laboratory in Keyworth, Nottingham, UK,using standard crushing and heavy mineral separation techniques.All zircon fractions were abraded (Krogh, 1982) and then washedin 4N HNO3 and H2O to remove possible traces of pyrite. Monazitegrains were washed in 1N HNO3 and subsequently in H2O. U and Pbwere extracted using the methods of Krogh (1973) and Corfu andAyres (1984). Fractions were spiked with a mixed 205Pb and 235Uisotopic tracer prior to digestion and chemistry (Krogh and Davis,1985). U and Pb were loaded together onto outgassed single Re fil-aments with silica gel. Isotope analyses were performed on a VG354 mass spectrometer using a Daly photomultiplier ion countingsystem, supplemented in part by multiple Faraday cup collectionof the larger ion beams. Pb isotope ratios were corrected for ini-tial common Pb in excess of laboratory blank using the common Pbevolution model of Stacey and Kramers (1975). The laboratory blankduring the time of analysis was 3 pg. Ages were calculated using thedecay constants of Steiger and Jäger (1977) and data reduction wascarried out using PBDAT (Ludwig, 1989).

Appendix B. LA-SF-ICP-MS U–Pb isotope analyses(Supplementary Table S2)

Zircon concentrates were prepared using standard crush-ing, sieving, magnetic and heavy mineral separation techniques.Selected grains were separated under alcohol and set in epoxyresin on round disks. Subsequently, the zircon grains were pol-ished to reveal their centres. Prior to analysis all grain mountswere photographed and zircon grains were imaged by SEMcathodoluminescence (CL) to identify different growth domainsor zoning patterns. This was done using the JEOL JSM-6400 elec-tron microprobe at the University of Frankfurt. Uranium, thoriumand lead isotopes were analyzed using a Thermo-Scientific Ele-ment 2 sector field ICP-MS coupled to a New Wave ResearchUP-213 ultraviolet laser system at Frankfurt (Gerdes and Zeh,2006, 2009). Data were acquired in time resolved – peak jump-ing – pulse counting mode over 810 mass scans, with a 19second background measurement followed by 30 second sam-ple ablation. The size of the ablated spots was 30 �m with15–20 �m crater depth. A teardrop-shaped, low volume laser cellenables precise monitoring of variations in the 207Pb/206Pb and206Pb/238U ratios during sequential depth profiling (see Janouseket al., 2006; Frei and Gerdes, 2009). Such variations can becaused for instance by drilling into mineral inclusions, domainsof different ages (e.g. older cores) or domains affected by Pbloss.



Raw data were corrected offline for background signal, commonPb, laser induced elemental fractionation, instrumental mass dis-crimination, and time-dependent elemental fractionation of Pb/Uusing an in-house MS Excel© spreadsheet program (Gerdes and Zeh,2006, 2009). A common Pb correction based on the interference-and background-corrected 204Pb signal and a model Pb compositionwas carried out, where necessary. The necessity of the correctionwas usually based on the 206Pb/204Pb (<10000). However, in casethe interference corrected 204Pb could not be precisely detected(e.g. <30 cps, counts per second), this was only applied when thecorrected 207Pb/206Pb was outside of the internal errors of the mea-sured ratios and yielded more concordant results. The interferenceof 204Hg (mean = 178 ± 18 cps) on mass 204 was estimated using a204Hg/202Hg ratio of 0.2299 and measured 202Hg and was correctedonline for each individual ratio (n=32). The data were acquired andprocessed in individual sequences consisting of 12 analyses of thereference zircon GJ-1 and that of 35 unknown zircons. Reporteduncertainties (2�) are propagated by quadratic addition of theexternal reproducibility (2 s.d., standard deviation) obtained fromthe GJ-1 zircon of the respective sequence (n = 12) and the within-run precision of each analysis (2 s.e., standard error). The offset(i.e. inter-elemental fractionation) of the measured 206Pb/238U ratiofrom the ‘true’ ID-TIMS value (0.0981 ± 0.004; ID-TIMS in-housevalue) of the GJ-1 zircon slowly decreased during the two-day ses-sion from +18% to −21% (Supplementary Fig. S1 and SupplementaryTable S4). Such changes, if observed, typically occur relative con-tinuously over several hours (Fig. S1) and can be corrected forperiods of ∼2 hours (= one sequence) by assuming a linear drift.This results typically in a within sequence precision (2 s.d.) of thedrift-corrected 206Pb/238U of 1.2–1.6%, as it was the case for the 11sequences of this study. It is worth noting that this inter-elementalfractionation is strongly dependent on the gas flow parameter andusually changes during retuning of the system (see Fig. S1). Pre-vious studies have shown that this approach yields precise andaccurate 207Pb/206Pb and 206Pb/238U ages on Palaeozoic to Protero-zoic grains that agree within 1% with results determined by theID-TIMS method (Gerdes and Zeh, 2006, 2009; Finger et al., 2008;Slama et al., 2008; Frei and Gerdes, 2009). For the 207Pb/206Pb ratioonly a minor correction of 0.2–0.3% (mass bias) was necessary. Asignal dependent 207Pb/206Pb uncertainty propagation was appliedfollowing the procedure outlined by Gerdes and Zeh (2009). Thereproducibility (2�) obtained from 113 analyses of the referencezircon GJ-1 are 1.1% and 1.3% for 207Pb/206Pb and 206Pb/238U, respec-tively. Concordia diagrams (2� error ellipse), concordia ages andupper intercept ages (95% confidence level) were calculated usingIsoplot/Ex 2.49 (Ludwig, 2001).

Appendix C. Th-U–Pb electron microprobe monaziteanalyses (Supplementary Table S3)

The in-situ analysis of ThO2, UO2 and PbO concentrations inmonazite grains with the electron microprobe (EMP) in thin sec-tion was used to calculate model ages following the proceduresoutlined by Suzuki et al. (1994) and Montel et al. (1996), respec-tively. The analytical conditions are summarised below; for moredetailed information on the technique see Schulz et al. (2007).

The in-situ ‘chemical’ Th-U–Pb dating of monazite, (LREE,Th)PO4, by EMP analysis is based on the observation that the con-centration of common lead in monazite is negligible compared tothe radiogenic lead resulting from the decay of Th and U (Parrish,1990; Cocherie and Albarede, 2001). Analysis of Th, U and Pb forcalculation of model ages, and of Ca, Si, Fe, Al, LREE, Yb and Y forcorrections, were carried out on a JEOL JXA 8200 (University ofErlangen, Germany) at 20 kV, 100 nA and a beam size of 5 �m. M�1lines were chosen for Th and Pb; the M�1 line for U; L�1 for La,

Y, Ce; L�1 for Pr, Sm, Nd, Gd and K�1 for Si, P, Fe, Al and Ca. Thecounting times for Pb, Th and U were 330 s, 30 s and 100 s per anal-ysis on peak and 2 x 120 s, 2 x 15 s and 2 x 50 s on background,respectively. Resulting errors (1�) are typically 1.20 %, 0.20 % and1.30 % for Pb, Th and U, based on counting statistics. La, Ce and Ndwere determined with 20 s (2 x 10 s), the other REE with 50 s (2 x25 s) counting time. Orthophosphates from the Smithsonian Insti-tute were used as standards for REE analysis (e.g. Donovan et al.,2003), PbO was calibrated using a vanadinite standard, and for Uan appropriate glass standard was employed. A Madagascar peg-matite monazite crystal termed ‘Madmon’ was used for calibrationof some LREE and ThO2. A small Y interference on the Pb M� line wascorrected by linear extrapolation after measuring several Pb-freeyttrium glass standards with 5 % and 12 % Y2O3 (Montel et al., 1996).An interference of ThM� on U M� was also empirically corrected,and a Gd-interference on U M� requires correction if Gd2O3 in mon-azite is larger than 5 wt%. For each single analysis a chemical ageand its respective error was calculated. To control the quality of theobtained model ages, the Madmon standard was measured togetherwith the monazite grains from our samples. The chemical analysesof the Madmon standard are given in Supplementary Table S5. Assummarised by Schulz et al. (2007), the age of the Madmon standardis well known and was determined by various techniques that giveages of 496±10 Ma (concordant SHRIMP U–Pb), 498 ± 2 Ma (TIMSPb-Pb evaporation age) and 502±6 Ma and 503 ± 6 Ma, respectively(EMP chemical model ages). Weighted average ages for several anal-yses within one grain and for monazite populations in the sampleswere calculated after Ludwig (2001).

Appendix D. Supplementary data

Supplementary data associated with this article can be found, inthe online version, at doi:10.1016/j.precamres.2009.03.008.

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palaeoproterozoic to palaeozoic magmatic and metamorphic events in the shackleton range, east...

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