Write about Antarctica before and after the Supercontinent Gondwana

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Gondwana Research 19 (2011) 335–371 Contents lists available at ScienceDirect Gondwana Research j o u r n a l h o m e p a g e : w w w. e l s ev i e r. c o m / l o c a t e / g r GR Focus Antarctica — Before and after Gondwana Steven D. Boger School of Earth Sciences, The University of Melbourne, Victoria 3010, Australia a r t i c l e i n f o Article history: Received 12 July 2010 Received in revised form 6 September 2010 Accepted 9 September 2010 Available online 21 September 2010 Editor: M. Santosh Keywords: Antarctica Gondwana Tectonics Subduction Accretion Collision Extension Rifting a b s t r a c t The origin of the Antarctic continent can be traced to a relatively small late Archaean cratonic nucleus centred on the Terre Adélie regions of East Antarctica and the Gawler Craton region of South Australia. From the late Archaean to the present, the evolution of the proto-Antarctic continent was remarkably dynamic with quasicontinuous growth driven by accretionary or collisional events, episodically punctuated by periods of crustal extension and rifting. The evolution of the continent can be broken into seven main steps: (1) late Palaeoproterozoic to middle Mesoproterozoic accretion and collision added crust first to the Antarctic nucleus's eastern margin, then to its western margin. These events resulted in the incorporation of the Antarctic nucleus within a single large continent that included all of Proterozoic Australia, a more cryptic Curnamona–Beardsmore Craton and most probably Laurentia. (2) Rifting in the middle to late Mesoproterozoic separated a block of continental crust of unknown dimensions to form an ocean-facing margin, the western edge of which was defined by the ancestral Darling Fault in Western Australia and its unnamed continuation in Antarctica. (3) Inversion of this margin followed shortly and led to the Grenville aged collision and juxtaposition of proto-Antarctica with the Crohn Craton, a continental block of inferred Archaean and Palaeoproterozoic age that now underlies much of central East Antarctica. The Pinjarra Orogen, exposed along the coast of Western Australia, defines the orogenic belt marking this collision. In Antarctica the continuation of this belt has been imaged in sub-ice geophysical datasets and can be inferred from sparse outcrop data and via the widespread dispersal of syn-tectonic zircons. (4) Tectonic quiescence from the latest Mesoproterozoic to the Cryogenian was the forerunner to Ediacaran rifting that separated Laurentia and the majority of the Curnamona–Beardsmore craton from the amalgam of East Antarctica and Australia. The result was the formation of the ancestral Pacific Ocean. (5) The rifting of Laurentia was mirrored by convergence along the opposing margin of the continent. Convergence ultimately sutured material with Indian and African affinities during a series of Ediacaran and Cambrian events related to the formation of Gondwana. These events added much of the crust that today defines the East Antarctic coastline between longitudes 30°W and 100°E. (6) The amalgamation of Gondwana marked a shift in the locus of subduction from between the preGondwana cratons to Gondwana's previously passive Pacific margin. The result was the establishment of the accretionary Terra Australis and Gondwanide orogenies. These were to last from the late Cambrian to the Cretaceous, and together accreted vast sequences of Gondwana derived sediment as well as fragments of older and allochthonous or para-allochthonous continental crust to Gondwana's Pacific margin. (7) The final phases of accretion overlapped with the initiation of extension and somewhat later rifting within Gondwana. Extension started in the late Carboniferous, although continental separation did not begin until the middle Jurassic. Gondwana then fragmented sequentially with Africa–South America, India, Australia and the finally the blocks of New Zealand separating between the middle Jurassic and the late Cretaceous. The late Cretaceous separation of Antarctica and Australia split the original Antarctic nucleus, terminating more than 2.4 billion years of shared evolution. The slightly younger separation of New Zealand formed the modern Antarctic continent. Crown Copyright © 2010 Published by Elsevier B.V. on behalf of International Association for Gondwana Research. All rights reserved. Contents 1. 2. 3. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tectonic elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The nucleus . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . E-mail address: sdboger@unimelb.edu.au. 1342-937X/$ – see front matter. Crown Copyright © 2010 Published by Elsevier B.V. on behalf of International Association for Gondwana Research. All rights reserved. doi:10.1016/j.gr.2010.09.003 336 336 336 336 S.D. Boger / Gondwana Research 19 (2011) 335–371 4. 5. 6. 7. 8. 9. 10. Palaeoproterozoic growth through collision—Nimrod–Kimban orogenesis . . . . . . Late Palaeoproterozoic magmatism and Early Mesoproterozoic accretion . . . . . . Mesoproterozoic rifting and reconfiguration—Albany–Fraser orogenesis . . . . . . . Late Mesoproterozoic rifting—the Darling Fault . . . . . . . . . . . . . . . . . . Late Mesoproterozoic collision—the Pinjarra Orogeny . . . . . . . . . . . . . . . Neoproterozoic rifting—the formation of the Pacific Ocean . . . . . . . . . . . . . Assimilation into Gondwana—the East African–Antarctic and Kuunga Orogenies . . . 10.1. Ediacaran events (580–550 Ma) . . . . . . . . . . . . . . . . . . . . . . 10.2. Ediacaran to Early Cambrian events (550–520 Ma) . . . . . . . . . . . . . 10.3. Cambrian events (530–490 Ma) . . . . . . . . . . . . . . . . . . . . . . 11. Post-Gondwana accretionary growth—the Terra Australis and Gondwanide Orogenies 12. Rifting of Gondwana and the formation of modern Antarctica . . . . . . . . . . . 13. Some concluding remarks . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1. Introduction Syntheses of the tectonic evolution of many of the Earth's main cratons have been presented at various times in the geologic literature (e.g. Hoffman, 1988; Myers et al., 1996; Betts et al., 2002; Karlstrom et al., 2001; Cawood and Korsch, 2008). Elements of Antarctica's tectonic evolution have also been summarized, often with a particular epoch in mind (e.g. Boger and Miller, 2004), or with reference to a particular element (e.g. Payne et al., 2009) or shared history with an adjoining continent (e.g. Fitzsimons, 2003). However, a synthesis of Antarctica's geologic evolution has not been undertaken in completeness—from the birth of the Antarctic continent in the Archaean to recent times when the continent attained its modern shape and size. The present paper attempts such a compilation. It builds naturally on an abundance of excellent studies that have focused on various elements of Antarctica's geology. These studies have in general moved the understanding of Antarctica's tectonic evolution from a relatively stable model, whereby much of the continent was considered little changed since at least Mesoproterozoic times (Yoshida, 1995; Yoshida et al., 2003), to one where Antarctica is viewed as a collage of different terranes that were stitched together at some stage in the geologic past (Fitzsimons, 2000a,b; Harley, 2003). The dynamism of Antarctica's evolution can be traced right to the origins of the continent. From the late Archaean onwards the rocks that would eventually rest in Antarctica have formed parts of landmasses modified by accretionary and collisional additions and rift driven subtractions. These events have quasi-continuously modified the shape and size of the future Antarctic continent—led to its incorporation into at least two of the Earth's main supercontinents (Rodinia and Gondwana) and formed in geologically recent times the remarkable and isolated southern continent which at 14.0 × 106 km2, is the fifth-largest of the Earth's modern continents. 2. Tectonic elements Antarctica can be divided into five broad tectonic domains (Fig. 1). These divisions are based on the similarities in the age and orogenic history observed in Antarctica when compared to that preserved in Antarctica's nearest neighbours. The four domains that fall within East Antarctica, a region that lies on the continental side of the Transantarctic Mountains between longitudes 30°W and 145°E, contain rocks that are all of Precambrian age (Fig. 1). Domains 1 and 2 have close geological affinities with southeastern Africa and eastern India respectively, regions that were connected to their Antarctic equivalents within Gondwana. Domain 3 shares a common Precambrian history with rocks now found in southern Australia. Domain 4 defines the ice-covered centre of the continent and contains rocks, which have no known correlatives. Domain 5 lies on the Pacific side of the Transantarctic Mountains (e.g. Ellsworth Land) and consists mostly of Phanerozoic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 338 341 342 344 345 347 349 350 352 354 357 360 362 362 363 aged rocks that were accreted to the margin of the continent while it was a component of Gondwana. The boundary between domains 3 and 5 is reasonably well defined both geologically and geophysically (Finn et al., 2006). The boundaries between the other sectors are relatively well constrained along the coast, but become more speculative inland. 3. The nucleus Each of the domains found in East Antarctica contains rocks that date back to the Archaean (Fig. 1b). Consequently, any of these domains could in principle be taken to represent the original nucleus around which the rest of Antarctica formed. Convention however bestows this honour on the late Archaean rocks of George V and Terre Adélie Lands and their correlatives in the Gawler Craton of South Australia (Fanning et al., 1988; Flöttmann and Oliver, 1994; Oliver and Fanning, 1997, 2002; Peucat et al., 1999). This region is inferred to underlie much of the East Antarctic icesheet inland of the Wilkes and Terre Adélie coastline (Fanning et al., 1995). If correct, these rocks potentially represent one of the largest continuous pieces of Archaean to Palaeoproterozoic crust in Antarctica. In contrast, the other Archaean terranes represent fragments of much larger terranes now located in Antarctica's neighbours (Fig. 1). The Grunehogna Complex of Dronning Maud Land for example represents a small fraction of the Kaapvaal Craton of southern Africa (Halpern, 1970; Groenewald et al., 1995; Jones et al., 2003). Similarly, the Napier and Vestfold Complexes of Enderby and Queen Elizabeth Lands have affinities with more extensive exposures of Archaean rocks in India. The potential exceptions are the Archaean rocks that fall within domain 4. Albeit extremely limited, the known exposures from this domain are of Archaean or Palaeoproterozoic age (Sheraton et al., 1992; Mikhalsky et al., 2001b, 2006a; Boger et al., 2006). However the extent to which these ancient rocks underlie domain 4 remains speculative as the majority of this region lies under the Antarctic ice-cap and is thus unexposed and unexplored. Fanning et al. (1995) coined the term “Mawson Block” to describe the Terre Adélie–Gawler nucleus of East Antarctica. Described here as the Mawson Craton, this region is taken here to include the rocks exposed west of the Kalinjala Shear zone in South Australia (Fig. 2), together with the correlative rocks exposed along the Terre Adélie coast of East Antarctica. Included within the Mawson Craton are the rocks exposed in the Miller and Shackleton Ranges of East Antarctica. Although preserving different protolith rocks when compared to those in Terre Adélie Land, the Miller and Shackleton Ranges nevertheless share a common period of Late Palaeoproterozoic orogenesis with Terre Adélie Land (Will et al., 2009). This orogenic event is manifested along the eastern margin of these three regions and its presence implies they formed a coherent craton before the onset of orogenesis. The boundary between the Archaean Terre S.D. Boger / Gondwana Research 19 (2011) 335–371 b 90˚E a Casey Casey Davis Davis Mawson Mawson Na Wilkes Land Enderby Land inss uunnttaain Moo TTrraan nssaa nnta r M ic ct 0˚ + V D 2 3 R Terre Adélie Adelie Dronning Maud Land Casey Casey Davis Davis Mawson Mawson Princess Elizabelth Land Ross Ice Shelf C 4 Victoria Land Ni 1 + Tectonic domains within Antarctica 180˚ S G 5 Ellsworth EllsworthLand Land Ronne Ice Shelf Exposed rock 90˚W 337 Ice shelves 1 Rocks with African affinities 2 3 4 Rocks with Indian affinities 5 Accreted pre- and post-Gondwana sediments, arc and para-autochthonous terranes Rocks with Australian affinities Rocks with no known affinities Archaean and Early Palaeoproterozoic Complexes C Commonweath Bay Ni Nimrod R Ruker D Denman Glacier S Shackleton G Grunehogna V Vestfold Na Napier Fig. 1. Present day Antarctica. (a) Distribution of rock exposures and major geographic place names. Mawson, Davis and Casey are the three Australian research stations. (b) Tectonic domains of Antarctica differentiated on the basis of their affinities with Antarctica's correlatives within Gondwana. Adélie–Gawler rocks and the younger rocks exposed along the northern and western margins of the Gawler Craton and their unexposed but presumed correlatives in East Antarctica are taken to define the western margin of the Mawson Craton. The rocks east of the Kalinjala Shear Zone, commonly described as part of the Gawler Craton in the Australian literature (e.g. Daly et al., 1998; Reid et al., 2008), are considered likely to have been accreted to the East Antarctic nucleus in the Late Palaeoproterozoic. They are thus not considered part of the Mawson Craton, but form part of the Curnamona–Beardmore Craton (described in the following sections). The oldest rocks of the Mawson Craton are best exposed in Gawler Craton where they define the Sleaford and Mulgathing complexes (Fig. 2). These complexes crop out in the southeastern and northwestern corners of the craton respectively and consist of mostly amphibolite to granulite facies late Archaean paragneisses overlain by Late Palaeoproterozoic sedimentary rocks (Daly et al., 1998). The late Archaean paragneisses were deposited between 2560 Ma and 2510 Ma and were sourced from predominantly juvenile to moderately evolved crust of late Archaean age (Swain et al., 2005b). The deposition of these units was followed shortly by the Sleafordian Orogeny, an event dated between 2450 Ma and 2420 Ma (Daly et al., 1998; McFarlane, 2006). Peak Sleafordian pressures vary between 5 and 9 kbar as a result of differential post-Sleafordian uplift, although peak metamorphic temperatures are reportedly more consistent, and vary between 750 and 850 °C (Fanning et al., 1988; Daly et al., 1998; Tomkins and Mavrogenes, 2002). The lowest grade rocks (amphibolite facies) are exposed in the vicinity of Coffin Bay (Fig. 2) and coincide spatially with the Dutton Suite, a series of deformed intrusions with emplacement ages of around 1850 Ma (Fanning et al., 2007). The Sleafordian Orogeny was followed by the intrusion of the Miltalie gneiss and its equivalents at approximately 2000 Ma, a magmatic event that accompanied the uplift and erosion of the Sleaford–Mulgathing basement to the surface (Fanning et al., 1988, 2007; Daly et al., 1998). Thereafter these rocks were overlain by the Hutchison Group, a sequence of chemical and clastic sediments confined to the eastern margin of the craton (Daly et al., 1998). These rocks fine upwards from basal quartz-pebble conglomerate and quartzite into lower energy calcareous and aluminous metasediments (Parker, 1980). A dated rhyodacite flow found near the top of the group constrains the deposition of the Hutchison Group to N1865 Ma (Fanning et al., 2007). Younger sedimentary rocks also overlie the Sleaford–Mulgathing basement. These consist of the greenschist facies volcano-sedimentary Price–Wangary paragneiss found in the vicinity of Coffin Bay (Fig. 2). The depositional age of these rocks is taken to be approximately 1765 Ma (Oliver and Fanning, 1997). This history is by and large replicated between latitudes 138°E and 145°E along the Terre Adélie coast of East Antarctica (Fig. 2). Rocks from this region preserve identical late Archaean formation and metamorphic ages when compared to those observed in the Gawler Craton (Stüwe and Oliver, 1989; Oliver and Fanning, 2002; Ménot et al., 2005). The equivalents of the Dutton Suite are known in Antarctica as the Cape Denison orthogneiss and similar to the Dutton Suite these rocks intrude lower grade rocks, which in Antarctica are exposed west of Cape Grey (Monnier et al., 1996; Ménot et al., 2005). A similarly good correlation exists between the South Australian Price–Wangary paragneiss and the Antarctic Cape Hunter phyllite. Both units contain the same populations of detrital zircons and show similar rock compositions and metamorphic grade (Oliver and Fanning, 1997). The Hutchison Group has no confirmed equivalents in East Antarctica, although marbles exposed at Watt Bay near the Mertz Glacier may represent the continuation of these rocks (C.M. Fanning, pers. com.). The East Antarctic icesheet obscures the extent to which the rocks of the Gawler Craton and their equivalents in Terre Adélie Land continue to the south. The best guess is given by satellite magnetic data that indicate that a regional magnetic high occurs over a large portion of Terre Adélie Land and the majority of the Gawler Craton (Maus et al., 2002). This anomaly continues approximately 700 km south of the Antarctic coast and would appear to be 500–700 km wide for most of this distance (Finn et al., 2006). Surface outcrops in the Transantarctic Mountains confirm that the late Archaean Terre Adélie–Gawler crust does not extend to the Miller Range (Fig. 2). In this region the exposed rocks define the Nimrod Group and these consist of a heterogeneous mix of sedimentary and igneous rocks that include middle Archaean (2980 Ma) orthogneisses (Goodge and Fanning, 1999). A single metamorphic rim with an age of c. 2955 Ma indicates that these rocks were probably metamorphosed significantly before the Sleafordian Orogeny (Goodge and Fanning, 1999). 338 S.D. Boger / Gondwana Research 19 (2011) 335–371 a b Fig. 2. The Mawson nucleus of East Antarctica. (a) Gawler (G), Nimrod (N) and Shackleton (S) components of the Mawson Craton. Grey outline of Antarctica and Australia shows the modern, but as yet unformed extent of these continents. Numbered regions (1–5) refer to the tectonic domains shown in Fig. 1. (b) Exposures of the Mawson Craton from the Terre Adélie (Antarctica) and Gawler (Australia) regions. Note: eastern and western margins of the Mawson Block are defined by the Mertz–Kalinjala shear zone and the Karari fault zone respectively. Bracketed ages in the domains north of the Karari Fault give the inferred timing of peak metamorphism. At the South American end of the Transantarctic Mountains, basement rocks are exposed in the Read Mountains and the southern Haskard Highlands of the Shackleton Range. These differ again from those of the Miller Range and Terre Adélie–Gawler regions and are defined by Palaeoproterozoic felsic orthogneisses dated at 2330 Ma and 1830 Ma (Brommer et al., 1999; Zeh et al., 1999; Will et al., 2009). Although the Terre–Adélie–Gawler, Miller and Shackleton regions (Fig. 2) preserve different Archaean and Early Palaeoproterozoic histories, they were deformed together during the Late Palaeoproterozoic Nimrod–Kimban Orogeny (1730–1690 Ma). The continuity of this orogenic belt suggests that these three disparate parts of the Mawson Block were probably contiguous prior to 1730 Ma (Will et al., 2009). 4. Palaeoproterozoic growth through collision—Nimrod–Kimban orogenesis Nimrod–Kimban orogenesis represents the oldest regionally continuous orogenic belt recognised in Antarctica (Fig. 3). Orogenesis is constrained to between 1730 Ma and 1690 Ma and, although regional differences exist, orogenesis for the most part resulted in high-grades of metamorphism, moderate to high peak pressures, and a clockwise P–T–t path (Brommer et al., 1999; Goodge and Fanning, 1999; Talarico and Kroner, 1999; Goodge et al., 2001; Will et al., 2009). At the Shackleton Range end of the orogen, Nimrod–Kimban orogenesis reworked the Palaeoproterozoic basement reaching metamorphic conditions in the upper amphibolite to granulite facies. Peak reported P–T conditions are 8 kbar and 800 °C and the P–T–t path is described as being clockwise (Talarico and Kroner, 1999). From the Miller Range (Fig. 2), eclogite facies metamorphism accompanied Nimrod–Kimban orogenesis. Peak P–T conditions are reported as 8– 12 kbar and 700 °C and similar to the Shackleton Range, the P–T–t path described a clockwise loop (Goodge et al., 1992; Peacock and Goodge, 1995; Goodge et al., 2001). In Terre Adélie Land, Nimrod–Kimban orogenesis was more variably manifested. At Cape Hunter, Nimrod–Kimban deformation S.D. Boger / Gondwana Research 19 (2011) 335–371 339 Fig. 3. Antarctic palaeogeography in the late Palaeoproterozoic. Nimrod–Kimban collision between Mawson and contiguous North Australian, Curnamona–Beardmore and Laurentian cratons. Grey outline of Antarctica and Australia shows the modern, but as yet unformed extent of these continents. Numbered regions (1–5) refer to the tectonic sectors of Antarctica shown in Fig. 1. resulted in the development of a single generation of north–south trending folds within the Cape Hunter phyllite and reached no more than greenschist facies (Oliver and Fanning, 1997). In the underlying Archaean basement, Nimrod–Kimban reworking is thought to have reached conditions of 6–7 kbar and 500–600 °C (Stüwe and Oliver, 1989). The higher-grade equivalents of the Cape Hunter phyllite are exposed near Dumont Dúrville (Fig. 2) and from these rocks metamorphic ages of 1690 Ma have been obtained (Peucat et al., 1999). The P–T estimates from this area are 4–6 kbar and 700–750 °C and these relatively low-P and high-T conditions are interpreted to reflect a period of extension and lithospheric thinning (Monnier et al., 1996; Peucat et al., 1999). This P–T evolution is unusual for the Nimrod–Kimban Orogeny and in many respects the age and style of metamorphism more closely match that which would be expected from the western Gawler Craton where high thermal gradients associated with Ifould (1690–1670 Ma) magmatism are to be expected. Elsewhere along the Terre Adélie coast, Nimrod–Kimban structures define regional scale steeply dipping high-strain zones that show both dextral strike-slip and reverse offsets (Stüwe and Oliver, 1989; Talarico and Kleinschmidt, 2003; Ménot et al., 2005; Di Vincenzo et al., 2007). The most prominent of these structures is the Mertz Shear Zone, a feature that appears comparable in terms of orientation and kinematics with the Kalinjala Shear Zone of southern Australia (Fig. 2). The Kalinjala Shear Zone is exposed along the eastern side (Fig. 2) of the Eyre Peninsula and defines the eastern edge of the Sleaford Complex (Vassallo and Wilson, 2002; Talarico and Kleinschmidt, 2003; Di Vincenzo et al., 2007). Deformation along the Kalinjala shear zone is dated at 1690–1680 Ma and this is interpreted to date the latter stages of Nimrod–Kimban orogenesis (Swain et al., 2005b). Elsewhere on the Eyre Peninsula this event is constrained to between 1730 Ma and 1700 Ma based on the ages obtained from the syn-tectonic Middle Camp granite and the post-tectonic Moody granodiorite (Fanning et al., 2007). Peak metamorphic conditions near the western margin of the Kalinjala shear zone are estimated at 8–10 kbar and 850 °C with a decompression dominated P–T–t path (Tong et al., 2004; Dutch et al., 2010). Further to the west and away from the craton margin, Nimrod–Kimban deformation and metamorphism were more heterogeneously distributed and commonly localised into discrete high-strain zones (Swain et al., 2005a). Between the shear zones the effects of the Nimrod–Kimban orogenesis appear limited to low-grade fluid driven retrogression; an observation consistent with the preservation of pre-Nimrod– Kimban cooling ages throughout most of the Sleaford and Mulgathing complexes (Webb et al., 1986; Tomkins et al., 2004). The continuity of the Nimrod–Kimban Orogeny along the eastern margin of the Mawson Craton suggests that this margin probably formed a Palaeoproterozoic ocean–continent or continent–continent plate boundary (Goodge et al., 2001). This is consistent with the 340 S.D. Boger / Gondwana Research 19 (2011) 335–371 moderate to high pressures accompanying metamorphism and the generally decompressive P–T–t paths described from along the belt. In Antarctica there is a rapid transition from rocks deformed and metamorphosed during the Nimrod–Kimban Orogeny, to much younger sedimentary rocks deposited when Neoproterozoic rifting formed the Palaeo-Pacific margin of Gondwana. There is thus little direct evidence as to what drove Nimrod–Kimban orogenesis or whether this event ultimately terminated in ocean closure. This is not so for Australia where Archaean and Palaeoproterozoic rocks are exposed on both sides of the Nimrod–Kimban orogenic front. Defined in Australia as the Curnamona Province (Fig. 3), the rocks to the east of the Kalinjala Shear Zone do not share a common history with those of the Mawson Craton. The rocks of the Curnamona Province include in a small region of middle Archaean basement defined by 3150 Myr old felsic orthogneiss which, in addition to their emplacement ages, record evidence for 2530–2510 Ma metamorphism (Fraser et al., 2010). This event pre-dates the Sleadord Orogeny observed in the Mawson Craton. More widely the Curnamona Province is defined by the Palaeoproterozoic igneous rocks of the Donington (1850 Ma) Suite, rocks which intrude isotopically distinct sedimentary rocks that have no known correlatives west of the Kalinjala Shear Zone (Howard et al., 2009). Overlying these rocks are the Myola (1765–1735 Ma) volcanics (Daly et al., 1998). These rocks are mostly exposed along the shores of the Spencer Gulf in South Australia beyond which to the east, the preNimrod–Kimban rocks of the Curnamona Province are covered by younger sediments. The oldest of these define the 1710–1640 Ma Willyama Supergroup and the presumed equivalents defined by the Radium Creek metamorphics (Page and Laing, 1992; Page et al., 2005; Conor and Preiss, 2008). These rocks are exposed respectively in the Broken Hill–Olary and Mt Painter inliers. Although the Curnamona Province cannot be traced in the surface geology into Antarctica, evidence for Palaeoproterozoic crust east of the Nimrod–Kimban orogenic front in Antarctica is nevertheless present. Borg et al. (1990) and Borg and DePaolo (1994) reported a marked step in isotopic composition (εNd, εSr, and TDM) from the Cambro–Ordovician intrusions emplaced in the Mawson Craton when compared to those that intrude the Neoproterozoic and younger sediments exposed to the east. These latter rocks overlie the southward extension of the Curnamona Province, if these rocks continue into Antarctica. The isotopic differences were taken to reflect the derivation of the Cambro–Ordovician granites from two unrelated basement provinces (Fig. 3). The first consisted of an older inboard Mawson Craton characterised by Archaean crust, while the second was considered younger and allochthonous, and to be defined by Palaeoproterozoic crust (Borg et al., 1990). Borg et al. (1990) defined this younger terrane as the Beardmore Microcontinent. The Beardmore Microcontinent was taken to form a narrow and continuous, but unexposed, strip that extended from at least the central Transantarctic Mountains to the Australian facing margin of Antarctica (Fig. 3). If one thus assumes that the Curnamona–Beardmore rocks define part of the same Palaeoproterozoic plate, the obvious conclusion is that the combined Curnamona–Beardmore Province collided with the Mawson Craton as a result of Nimrod–Kimban orogenesis. This is contrary to the original interpretation of the Beardmore Microcontinent, which was considered to have sutured with the Mawson Craton during the Early Cambrian Ross Orogeny (Borg et al., 1990). However more recent studies dispute this conclusion on the basis of detrital zircon and sedimentological data. These datasets suggest that sedimentary rocks overlying the Beardmore Craton have an East Antarctic provenance, a finding that argues for an autochthonous rather than allochthonous origin of these sediments (Goodge et al., 2002; Myrow et al., 2002). This conclusion precludes the collision of a Curnamona–Beardmore Craton during the Cambrian, a scenario that would require the overlying sediments to also be allochthonous. Interestingly however, these data do not preclude an allochthonous origin for the Beardmore rocks, as long as they were juxtaposed with the Mawson Craton prior to the onset of deposition of the overlying strata. The juxtaposition of Beardmore rocks against those of the Mawson Craton in the Palaeoproterozoic is thus consistent with: (1) the sharp isotopic discontinuity between the Mawson and Beardmore blocks as described by Borg et al. (1990), (2) the evidence for Palaeoproterozoic orogenesis along the boundary between these two disparate rock sequences, (3) the inferred Palaeoproterozoic timing of collision between the Curnamona and Mawson cratons from the Australian sector of this margin (Myers et al., 1996; Giles et al., 2004; Betts and Giles, 2006) and, (4) the autochthonous origin of the Ediacarian and younger sediments that overlie this margin in the Transantarctic Mountains. In view of this, it is the author's opinion that Nimrod–Kimban orogenesis terminated with continental collision. The colliders were on the one side the Mawson Craton and on the other, a continuous Curnamona–Beardmore Craton. The size and shape of this Curnamona– Beardmore Craton remains to a large extent unknown. However, a number of studies have forwarded the notion that these rocks may have correlatives in the remainder of Proterozoic Australia and quite possibly Laurentia. The correlation of rocks between the North Australian Craton and those of the Curnamona Craton is relatively well established and is based on two main observations: (1) There are close temporal and structural/metamorphic similarities between rocks exposed in central Australia and the Nimrod–Kimban Orogeny (Giles et al., 2004), and (2) the cover-sequences post-dating Nimrod–Kimban orogenesis found in the Isan (North Australia) and Willyama (Curnamona) basins share a number of temporal, depositional and metallogenic features (Page et al., 2005). The fit for these terranes is after that of Giles et al. (2004) and is shown in Fig. 3. This reconstruction is consistent with the available palaeomagnetic data and aligns the structural grains of the similarly rocks exposed in the Arunta Inlier of central Australia with those of the Kimban Orogeny of the eastern Gawler Craton (Wingate and Evans, 2003). This fit also results in the Isan and Willyama basins being moved into relative proximity. Giles et al. (2004) and Betts et al. (2008) extend this model further to include Laurentia (Fig. 3). These authors propose a fit that aligns the Yavapai Orogeny of central North America with the temporally equivalent Nimrod–Kimban Orogeny of East Antarctica and Australia. In doing so, they place the western end of the Yavapai Orogeny at the Shackleton Range end of East Antarctica and argue for a continuous Yavapai–Nimrod–Kimban–Strangways Orogeny. They further argue that this belt formed a continuous accretionary margin along a large combined continent consisting of Laurentia and the combined North Australian and Curnamona cratons (Fig. 3). This model is not dissimilar to the SWEAT reconstruction (Moores, 1991), but aligns the structural grains of the Laurentian and East Antarctic–Australian belts better. The model also shows a gap between Laurentia and Australia (Betts et al., 2008 — Fig. 4), which interestingly is filled if one includes the Beardmore Microcontinent (Fig. 3). Betts et al. (2008) argue that prior to the collision of the Mawson Craton, the Nimrod–Kimban Orogeny, and its extensions along the margins of the North Australian and Laurentian cratons (Fig. 3), formed an ocean-facing margin onto which material was periodically accreted. If the Betts et al. (2008) scenario is correct, it implies that in addition to the Curnamona–Beardmore Craton, both Proterozoic Australia and Laurentia opposed the Mawson Craton in the Palaeoproterozoic. In this case the Mawson Craton formed a relatively small addition to a much larger Late Palaeoproterozoic continent (Fig. 3). This scenario is however not universally accepted and a number of studies suggest that the Mawson Craton had no tectonic interaction with the North Australian Craton until later in the Proterozoic. Both 1580–1540 Ma (Wade et al., 2006) and 1300–1150 Ma (Myers et al., 1996) are proposed alternate times of collision. Alternate relative positions for Laurentia and Australia are also implied for this interval (e.g. Burrett and Berry, 2000; Karlstrom et al., 2001). S.D. Boger / Gondwana Research 19 (2011) 335–371 With respect to the formation of Antarctica these arguments are not particularly pertinent. The important points are that: (1) the Mawson nucleus can be traced continuously from the Shackleton Range in East Antarctica to southern Australia, (2) the eastern margin of this craton underwent Late Palaeoproterozoic orogenesis and, (3) this event most probably involved a collision between the Mawson nucleus and a combined Curnamona–Beardmore craton. It is possible that this event led to the incorporation of a relatively small Mawson Craton into a much larger continent that included the rocks of North Australia, Laurentia and the intervening Curnamona–Beardmore Craton (Fig. 3). Support for such a conclusion is discussed in more depth in the following sections. Nevertheless, it is alternatively possible that Nimrod–Kimban orogenesis was of more limited extent and sutured the Mawson and Curnamona–Beardmore cratons in a collision of approximate equals. 5. Late Palaeoproterozoic magmatism and Early Mesoproterozoic accretion Subsequent to the collision of the Mawson and Curnamona– Beardmore cratons, ocean closure migrated from the eastern to the western side (Fig. 4) of the Mawson Craton (Betts and Giles, 2006). Evidence for the shift in the locus of subduction is inferred from the emplacement of extensive and commonly arc-type intrusions in the western and northern margins of the Mawson Craton (Parker, 1993; Daly et al., 1998; Swain et al., 2008). These magmatic events define the Tunkillia–Ifould (1690–1670 Ma) and St Peter (1620–1610 Ma) suites. a Mus Mus an O lar i Isa n-O ocean closure known plutons 5 Casey rifting inferred plutons C Mawson 2 ? N Mawson Craton Davis W 1490-1370 Ma trans-Laurentian igneous suite r ic en og North Australia nt F ro Cur N 1400-1330 Ma Accretion of the Warumpi Province (1640-1630 Ma) W Westward migration of the active margin with time Coompana strata sourced from Kararan Orogeny b Nimrod-Kimban orogenic belt (1730-1710 Ma) B Casey Mawson 5 2 ? i pa va Ya ov Pr inc 1 e Laurentia Mawson Craton } collision at 1720 Ma } } } collision between 1640 Ma and 1550 Ma North Australia B = Beardmore Cur = Curnamona Laurentia inc ov Pr = Warumpi c. 1640 Ma = Musgrave c. 1550 Ma = Coompana c. 1550 Ma = Nawa c. 1550 Ma = Mazatzal c. 1640 Ma i pa va Ya M W Mus C N M B 5 4 Laurentia M 5 C Davis 1 Cur Mawson Craton Kararan (1570-1530 Ma) accretion of the Nawa and Musgrave domains Both suites are exposed southeast of the Coorabie Fault Zone within the Mulgathing Complex (Swain et al., 2008, fig. 2). The Tunkillia–Ifould suite represents a relatively juvenile (εNd(t) = −6.0 to +3.0) potassic calc-alkaline suite of intrusions that formed via a mixed contribution of mantle and crustal melt (Payne et al., 2010). Betts et al. (2006) argue for an arc related origin, a conclusion consistent with relatively depleted Y and Nb concentrations. In contrast, Payne et al. (2010) argue such an equivocal tectonic classification cannot be drawn for these rocks. The geochemistry of the St Peter suite is clearer and these rocks show distinctive subduction related signatures (Hand et al., 2007; Swain et al., 2008). Evidence for deformation accompanying both suites is not widely manifested and appears mostly to have been focused along the major pre-existing shear zones (Swain et al., 2005a; Stewart and Betts, 2010). Deformation of this age may have also reoriented the Archaean structures between these lineaments, although it had little isotopic impact as the bulk of the Mawson Craton records evidence for older cooling (Tomkins et al., 2004). Subsequent Hiltiba magmatism (1600–1580 Ma) occurred throughout the Mawson Craton and in Australia resulted in the widespread emplacement of shallow plutons and the extrusion of the coeval and commonly flat-lying volcanic rocks. Temporal and geochemical equivalents of this suite are found in moraines along the coast of Terre Adélie Land, an observation that attests to the continuation of this magmatic province into Antarctica (Peucat et al., 2002). Although commonly interpreted as an anorogenic magmatic event, it is likely that some reactivation accompanied this event (McLean and Betts, 2003; Hand et al., 2007). The origin of the Hiltiba 1640-1400 Ma North Australia 341 e Fig. 4. Antarctic palaeogeography in the early Mesoproterozoic. (a) 1640–1400 Ma. Arc magmatism and accretion of terranes along the western (LHS) margin of the North Australia, Curnamona–Beardmore, Laurentia and Mawson amalgam. Trans–Laurentian igneous suite intrudes Antarctica, Laurentia, and possibly SW Australia (b) 1400–1330 Ma. Collapse of ocean between North Australia and the western (LHS) margin of the Mawson Block reorganised the North Australian–Mawson configuration. Relative motion terminated by collision along the Albany–Fraser Orogen. Grey outline of Antarctica shows the modern, but as yet unformed extent of the continent. Numbered regions (1–5) refer to the tectonic domains of Antarctica shown in Fig. 1. 342 S.D. Boger / Gondwana Research 19 (2011) 335–371 suite has on the one hand been argued to be plume related, while on the other is taken to represent back-arc intrusions emplaced distal to a subduction zone (Creaser, 1996; Wade et al., 2006; Betts et al., 2007). More recently a hybrid model has been proposed (Betts et al., 2009). The Betts hypothesis suggests that subduction associated with St Peters suite magmatism (1620–1610 Ma) overroad an oceanic hotspot. The effect was to flatten subduction and terminate arc magmatism with the greater coupling between the overriding and subducting plates driving distal deformation (Isan–Olarian) in the eastern Curnamona–North Australian Craton. Interaction of the hotspot with the overlying continental lithosphere was then responsible for the Hiltiba magmatism. Betts et al. (2009) then argue that migration of the hotspot away from the Mawson margin allowed for the reestablishment of more typical subduction, which persisted until the onset of the Kararan Orogeny. The Kararan Orogeny (1570–1540 Ma) is most widely developed north of the Karari Fault Zone, a structure that defines the northern and western margins of the Archaean nuclei of the Mawson Craton (Fig. 2). Rocks north of the Karari Fault Zone define a series of separate tectonic domains, all of which lack the Archaean protoliths observed south of this structure. The more important of these domains are the Nawa, Mabel Creek and Coober Pedy domains (Fig. 2). Largely unexposed, these rocks are understood mostly from geophysical data and samples collected from drill core. Rocks from the Nawa domain are defined by paragneisses deposited after 1740 Ma that were subsequently metamorphosed at temperatures of 950 °C and pressures of 10 kbar during an event known as the Ooldean Orogeny (Payne et al., 2006; Hand et al., 2007). Ooldean orogenesis occurred at either 1690 Ma or between 1660–1630 Ma depending on how one interprets the existing geochronological data (Daly et al., 1998; Hand et al., 2007; Payne et al., 2008). Metamorphism of Kararan age is also observed in the Nawa domain (Payne et al., 2008). In the adjacent Mable Creek and Coober Pedy domains, similarly high-grade conditions are implied, although metamorphism in these domains is attributed to Kararan rather than Ooldean orogenesis (Daly et al., 1998). The differences in geologic history preserved across the Karari Fault Zone have led a number of studies to conclude that the rocks north of the Karari Fault Zone were allochthonous with respect to those of the Mawson Craton (Fitzsimons, 2003). This conclusion is supported by the preservation of Ooldean (1660–1630 Ma) orogenesis north of this structure, and its apparent absence to the south. The Kararan Orogeny (1570–1540 Ma) is common to both the northern and southern sides of the Karari Fault Zone and it is at this time that the northern domains are argued to have sutured to the Mawson Craton (Fitzsimons, 2003; Direen et al., 2005; Payne et al., 2006). It is not possible to trace the magmatic and accretionary events observed along the western margin of the Gawler Craton across the rifted margin into Antarctica. It is nevertheless likely that some of the Palaeoproterozoic rocks underlie the Antarctic icesheet somewhere between longitudes 115°E and 140°E. Supporting evidence is derived from the Transantarctic Mountains where sedimentary rocks show large populations of detrital zircons with ages between 2000 Ma and 1400 Ma (Goodge et al., 2002, 2004). In particular, age peaks between 1580 Ma and 1520 Ma match the timing of Kararan orogenesis, while a broad spread of ages between 1700 and 1900 Ma are similar to the basement and detrital ages obtained from the Nawa domain (Payne et al., 2006). The source of these rocks is from central East Antarctica (Myrow et al., 2002), an observation consistent with elements of the western Gawler Craton continuing across the rifted margin into Antarctica. The next domain to the southwest of the Nawa domain extends to the Australian coastline and, although similarly unexposed in Antarctica, the continuity of these rocks to the rifted margin almost certainly guarantees their presence in Antarctica (Fig. 4). Known as the Coompana Block (Myers et al., 1996; Wade et al., 2007), these rocks are for the most part covered and little understood, although limited exposures are present in southern Western Australia where these rocks define Nornalup Complex (Fig. 5b). The available data from the Nornalup Complex suggest that these rocks are younger than those exposed in the Nawa and adjacent domains to the east. The protolith rocks consist of Mesoproterozoic paragneisses that in all likelihood were derived from the Kararan Orogeny (Clark et al., 2000; Fitzsimons, 2003). The timing of orogenesis in these rocks is largely unknown, although they are intruded by cross-cutting plutons with ages between 1505 Ma and 1450 Ma. The intrusion of these rocks was coincident with the reworking of shear zones in the Mawson Craton (Myers, 1995; Fraser and Lyons, 2006; Wade et al., 2007). The age and chemistry of these Mesoproterozoic intrusions i.e., A-type, εNd = 1.2– 3.3, TDM = 1.75–1.94 Ga (Wade et al., 2007), suggest strong affinities with the Trans–Laurentian igneous suite (1490–1370 Ma). These latter rocks are known from central Laurentia and are also implied to be present in central East Antarctica (Goodge et al., 2008). If one considers the reconstruction shown in Fig. 4, subduction related magmatism and deformation observed between 1690 Ma and 1540 Ma along the western margin of the Mawson Craton would have been similarly manifested along both the North Australian and Laurentian margins (Fig. 4a). The accretion of the Warumpi Province to the North Australian craton falls within this interval, as does the accretion of the juvenile crust of the Mazatzal Province to the southern margin of Laurentia (Condie, 1992; Scrimgeour et al., 2005). Similarly, Kararan orogenesis coincides temporally with the emplacement of island arc magmas into Musgrave Block (Wade et al., 2006), which may imply both subduction and micro-terrane accretion along the same margin. An alternative hypothesis suggests that the 1690–1540 Ma accretionary history observed along the western margin of the Mawson Craton was terminated by collision between the North Australian and the Mawson cratons (Wade et al., 2006). This scenario is supported by the overlap of ages in both cratons from about 1580 Ma onwards (Fig. 4a). However, by placing the North Australian and Mawson cratons into their modern configuration by 1500 Ma this model implicitly implies that the Albany–Fraser Orogen (see below) was of intra-plate rather than inter-plate origin. This conclusion is at odds with the majority of studies from the Albany–Fraser Belt (Myers et al., 1996; Clark et al., 2000; Bodorkos and Clark, 2004a,b; Cawood and Korsch, 2008). With respect to the Mesoproterozoic relationship between Laurentia and the Mawson Craton, the presence of the 1490– 1370 Ma Trans–Laurentian igneous suite within both continents implies that they were likely connected (Goodge et al., 2008). These rocks are not directly observed in East Antarctica, but are identified from large quantities of 1470–1370 Ma detrital zircon and clasts of 1440 Ma granite found in sediments and glacial moraines derived from central East Antarctica (Myrow et al., 2002; Goodge et al., 2002, 2004, 2008). There are thus a number of geological similarities between Laurentia and North Australia/Mawson Craton that can be traced from at least the late Palaeoproterozoic to the early Mesoproterozoic (Betts et al., 2008; Goodge et al., 2008). It therefore seems reasonable to assume the configuration of the continents, or some variant thereof, shown in Fig. 3 continued unchanged into the epoch shown in Fig. 4. 6. Mesoproterozoic rifting and reconfiguration—Albany–Fraser orogenesis The middle Mesoproterozoic tectonic evolution of Antarctica was one of reorganisation of the existing crustal elements (Fig. 4b). Assuming the connectivity between the North Australia and Curnamona–Beardmore cratons illustrated in Fig. 4a, Giles et al. (2004) suggest that between 1500 and 1350 Ma the North Australian Craton rotated via slab retreat along a subduction zone west of Mawson S.D. Boger / Gondwana Research 19 (2011) 335–371 1140-1080 Ma a b 116˚E 343 120˚E 118˚E 122˚E n ge Cur N 34˚S Albany Mobile Belt Fr as er M ob i lt Darling Fau 5 32˚S 34˚S Esperance 65˚S C Davis Mawson Mawso Ma wson n 2 Ma Mawso wso nn Crat Craton on Al ba n y-F rase rO ro Yilgarn Craton (North Australia) Perth 32˚S North Australia le Be lt Darling Fault Albany Dunsbourough Fault B Casey Windmill Islands Obruchev Hills ? S 65˚ 4 ˚E ˚E 105 Mawson Craton Nornalup Complex } Corumup Complex Crohn Craton Biranup Complex 0˚E 200 km deformed western margin of Mawson Craton Fraser Complex (c. 1300 Ma) 10 Rifting produced the sharp high-angle contact between the Yilgarn and Albany-Fraser rocks and those of the subsequent Pinjarra Orogeny Mirny ˚E Rifted component of North Australia and proto-Antarctica of unknown size Mt Amundsen Mt Barr-Smith 95 ce vin ro iP pa va Ya M cier Gla an nm De Laurentia ˚E 110 Sco ttGla cier 115 5 David Island 1 Mawson Craton Bunger Hills Hill Yilgarn Craton Leeuwin Complex } } mostly exotic material southern margin of north Australian Craton Perth basin Fig. 5. Antarctic palaeogeography in the middle Mesoproterozoic. (a) Rifting along the ancestral Darling Fault. Grey outline of Antarctica shows the modern, but as yet unformed extent of this continent. Numbered regions (1–5) refer to the tectonic domains shown in Fig. 1. Domains signified with letters are: B = Beardmore, C = Coompana, Cur = Curnamona, M = Mazatzal and N = Nawa. Stars denote the presence of the known (black) and inferred (grey) plutons of the Trans–Laurentian igneous suite. (b) Correlation of units between the Wilks and Queen Elizabeth regions of East Antarctica and southwestern Australia. The continuation of the Darling Fault in Antarctica approximately underlies the Scott Glacier and separates the Obruchev and Bunger Hills. Craton. This resulted in rifting between the Curnamona and North Australian cratons and the collapse and closure of the ocean between the southern margin of the North Australian Craton and the western margin of the Mawson Craton (Fig. 4b). The result was collision between the southern margins of the North Australian Craton and the western margin of the Mawson Craton (Myers, 1993; Clark et al., 2000). Collision occurred along the Albany–Fraser Orogen of Western Australia, an orogenic belt that can be traced along strike into both the Musgrave region of Central Australia and the Wilks Land region of East Antarctica (Black et al., 1992b; Sheraton et al., 1993; Clarke et al., 1995; Myers et al., 1996; White et al., 1999). Two stages of orogenesis are reported. These occurred between 1340 Ma and 1260 Ma (stage I) and between 1215 Ma and 1140 Ma (stage II; Clark et al., 2000). The central part of this orogenic belt is exposed in southwestern Western Australia (Fig. 5b) where is subdivided into four major units. From north to south these units are defined by: (1) the thermally and structurally reworked margin of the North Australian Craton defined by the Yilgarn Craton (Beeson et al., 1988; Jones, 2006); (2) the Biranup Complex, a narrow mostly continuous domain defined by Late Archaean and Palaeoproterozoic protoliths exposed along the margin of the Yilgarn Craton. These rocks may represent either reworked elements of the Yilgarn Craton, or younger and exotic material that was accreted to the margin of the Yilgarn Craton during Albany–Fraser orogenesis (Nelson et al., 1995); (3) the Fraser Complex, comprised mostly of high-grade Mesoproterozoic mafic rocks of island arc affinity (Clark et al., 1999; Condie and Myers, 1999) and; (4) the western margin of the Mawson Craton defined by the Nornalup Complex (Myers, 1995; Clark et al., 2000; Wade et al., 2007; Cawood and Korsch, 2008). The oldest structures are developed in the western margin of the Mawson Craton. Deformation began slightly before 1330 Ma and was accompanied high-T and low-P (750 °C and 4 kbar) metamorphism (Clark et al., 2000). Geochemical and isotopic data from the syntectonic Recherche granitoids indicate a substantial mantle component (Nelson et al., 1995) and these data, together with the absence of pre-1315 Ma intrusions within Biranup or Fraser complexes, are taken to suggest that subduction prior to Albany–Fraser collision occurred beneath the Mawson Craton (Clark et al., 2000; Bodorkos and Clark, 2004b). This is consistent with the long-term subduction polarity and westward younging of orogenic and plutonic events observed along the western margin of the Mawson Craton. Subsequent deformation was localised entirely within the Corumup and Fraser complexes—rocks that lie effectively in the suture between the North Australian and Mawson cratons (Fig. 5). Orogenesis commenced at approximately 1310 Ma and initially occurred at relatively high-T and low-P metamorphic conditions (800–850 °C and 5–7 kbar, Clark et al., 1999; Bodorkos and Clark, 2004a). Subsequent tectonism occurred between 1295 Ma and 1285 Ma and involved 7 to 14 km of near isothermal burial (Nelson et al., 1995; Clark et al., 1999; Bodorkos and Clark, 2004a). The up-pressure part of the P–T–t path is considered to have resulted from the tectonic loading of the Fraser Complex as a result of stacking during collision (Bodorkos and Clark, 2004b). The absence of evidence for 1290–1280 Ma metamorphism within the Nornalup Complex was taken to suggest that these rocks 344 S.D. Boger / Gondwana Research 19 (2011) 335–371 a Warakurna large igneous province b 1080 Ma Sodruzhestvo Group (PCM) n=195 Denman Region Orthogneiss emplacement: 3005 Ma & 2640 Ma Metamorphism: 2890 Ma 35 30 North Australia 25 20 North Australia <1050 Ma 40 15 10 5 5 500 Cur N Another continent ? 1000 1500 2000 2500 3000 Adelaidean Group Kanmantoo Group (SEA) 5 n=560 60 50 40 SE SEA 30 20 AF 500 2000 2500 OH 3000 Skelton Group (SVL) n=118 30 SV SVL 25 20 10 5 ? 500 1000 1500 2000 2500 3000 Koettlitz Group (SVL) n=64 10 s iv e m a r g i n 15 PCM 5 en 1000 1500 eO 2500 n=538 ro g Laurentia 3000 Beardmore Group Byrd Group (TAM) e v i ll 2000 SR 70 Hannah Ridge and Patuxent Fms (PM) n=137 60 50 30 25 20 in c ov Pr Gr 500 5 PM i pa va Ya Sedimentary cover sequences: Stinear <2700 Ma Ruker <2450 Ma Blake <1800 Ma Sodruzhestvo <1040 Ma M in c ov Pr Orthogneiss emplacement: 3170 Ma Metamorphism: 2780 Ma Laurentia i pa va Ya Southern PCM AM TAM . 5 1 Vos os pa s g eny 2 1500 B Mawson Craton O ro Crohn Craton 1000 r ot r ra Davis Mawson C n ja Neop Pi 10 e 40 15 30 eny 10 20 5 10 500 1000 1500 2000 2500 3000 500 1000 1500 2000 2500 3000 Fig. 6. Antarctic palaeogeography in the late Mesoproterozoic. (a) Pinjarra collision between the Crohn Block and the North Australia, Curnamona–Beardmore, Laurentia, Coompana– Nawa and Mawson amalgam. Domains signified with letters are: AF = Albany–Fraser, B = Beardmore, C = Coompana, Cur = Curnamona, M=Mazatzal and N=Nawa. Stars denote the known (black) and inferred (grey) plutons of the Trans–Laurentian igneous suite. (b) Dispersal of Pinjarran derived sediments. Large arrows indicate inferred sediment transport away from a Pinjarran highland. Locations of sedimentary strata with large populations of Pinjarra zircons (grey stars) are: PCM = Prince Charles Mountains, PM = Pensacola Mountains, SEA = South Eastern Australia, SVL = Southern Victoria Land and TAM = Transantarctic Mountains. Detrital zircon data for each region shown in bar graphs with grey zone highlighting the age range attributable to Pinjarra orogenesis. White triangles indicate Antarctic locations where Pinjarran Orogenesis is either known from the surface geology or inferred from geophysics. OH = Obruchev Hills, SR = Shackleton Range and Vos = Lake Vostok. Grey outline of Antarctica shows the modern, but as yet unformed extent of this continent. Numbered regions (1–5) refer to the tectonic domains shown in Fig. 1. constituted the upper plate during collision (Myers et al., 1996; Clark et al., 2000; Bodorkos and Clark, 2004b). The second stage of Albany– Fraser orogenesis occurred from 1215 to 1140 Ma (Clark et al., 2000). This event is generally inferred to be of intra-plate origin (Clark et al., 2000; Fitzsimons, 2003), although metamorphic grade reached granulite facies in the extreme west of this belt and deformation is widely recognised in the Biranup Complex (Beeson et al., 1988; Black et al., 1992a; Duebendorfer, 2002). The propagation of the Albany–Fraser Orogen into East Antarctica is well established (Lovering et al., 1981; Sheraton et al., 1993; Fitzsimons, 2003). The two stages of orogenesis recognised in Western Australia are both observed along the Antarctic coast between longitudes 100°E and 115°E. Rocks from the Windmill Islands, exposed in vicinity of Casey (Wilkes Land, Fig. 5b), record both stages of Albany–Fraser orogenesis (Paul et al., 1995; Duebendorfer, 2002). The older event occurred at amphibolite facies and is dated between 1340 Ma and 1310 Ma. The second overprinting event occurred at granulite facies and is dated to between 1210 Ma and 1180 Ma (Post et al., 1997). The suturing phase of Albany–Fraser orogenesis (1310– 1270 Ma) is not observed in the Windmill Islands, suggesting that this region may correlate most closely with the rocks of the Nornalup Complex. If correct, the Windmill Islands probably formed part of the leading edge of the proto-Antarctic plate prior to collision (Fig. 5). This is consistent with TDM model ages that for both the Nornalup Complex and Windmill Islands show a common range of ages between 3.2 Ga and 1.9 Ga (Möller et al., 2002). Four hundred kilometres west, the rocks of the Bunger Hills (Fig. 5) are defined by intercalated felsic orthogneisses, two pyroxene mafic gneiss and pelitic and semi-pelitic gneisses. The orthogneisses are of granodioritic composition and were emplaced at approximately 1700 Ma (Sheraton et al., 1992). Subsequent deformation and highgrade metamorphism slightly pre-dated widespread plutonism dated between 1170 Ma and 1150 Ma (Stüwe and Powell, 1989; Sheraton et al., 1992, 1993). The age of the pre-orogenic orthogneisses and the predominance of Stage II Albany–Fraser orogenesis show close parallels between the Bunger Hills and the Biranup Complex of southern Western Australia (Fig. 5). 7. Late Mesoproterozoic rifting—the Darling Fault The westward continuation of the Albany–Fraser Orogeny, both in Antarctica and Australia, is terminated sharply against the Darling Fault, a lineament that can be traced along the entire length of the West Australian margin and that separates the Archaean rocks of the Yilgarn Craton from the Meso- and Neoproterozoic rocks of the Northampton, Mullingarra and Leeuwin complexes (Fig. 5). It is a steeply dipping crust-piecing structure that defines a sharp gravity boundary (Dentith et al., 1993) between the thick crust of the Yilgarn Craton and the younger and thinner Proterozoic complexes exposed to the west (Fig. 5). Evidence exists for both dip-slip and strike-slip movements, although the most pronounced offset was the result of sinistral strike-slip motion that reoriented by 90° the structures in the S.D. Boger / Gondwana Research 19 (2011) 335–371 western Albany–Fraser Orogen (Fig. 5). This deformation post-dates tectonism within the Albany–Fraser Orogen and has been attributed to events at around 750 Ma and 550 Ma (Fitzsimons, 2003). Its unnamed equivalent in Antarctica underlies the Scott Glacier and separates the Bunger Hills from the Obruchev Hills (Fig. 5b). An alternate possibility is that the Darling Fault was active in the Mesoproterozoic and resulted in the formation of a Late Mesoproterozoic ocean-facing margin along the western margin of Australia and proto-Antarctica (Fig. 5a). This inference is based firstly on the observation that the rocks of the Yilgarn Craton and the Albany–Fraser Belt terminate abruptly against the Darling Fault. Given that the westward continuation of Yilgarn or Albany–Fraser equivalents have not been observed directly west of the Darling Fault (although may be present in the offshore Naturaliste Plateau, Halpin et al., 2008), one must assume the complete removal of these rocks via rift and subsequent drift tectonics (Fig. 5a). This need not have occurred in the Mesoproterozoic, although this timing is considered likely as subsequent convergent orogenesis occurred at 1080 Ma. This event mostly likely occurred along or near a plate boundary (see the following section). Supportive of this notion is that the precursor rocks of the Northampton and Mullingarra complexes, rocks found west of the Darling Fault, contain detrital zircons with ages between 1900–1600 Ma and 1450–1150 Ma (Bruguier et al., 1999). These zircons are considered to have been derived from either the Capricorn or Albany–Fraser Orogens (Bruguier et al., 1999; Fitzsimons, 2003), a conclusion that implies that these rocks were derived from Australia and deposited west of the Darling Fault in the Late Mesoproterozoic. They could also have been sourced from the continuation of these terranes west of the Darling Fault, although the accumulation of sediment in what now are the Northampton and Mullingarra complexes implies that these regions were low-lying and thus probably floored by oceanic or thinned continental crust. The implication of this scenario is that a terrane of unknown size and defined by the westward continuation of the Yilgarn Craton and rocks of the Albany–Fraser Orogeny as well as their extensions into Antarctica, was likely removed from the western margin of the Australo–Antarctic plate between the Albany–Fraser (1340–1140 Ma) and subsequent Pinjarra (1080–1060 Ma) orogenies. 8. Late Mesoproterozoic collision—the Pinjarra Orogeny Rock exposures west of the Darling Fault are limited to three basement inliers—the Northampton, Mullingarra and Leeuwin complexes. Of these, the Northampton and Mullingarra complexes record orogenic ages between 1080 and 1060 Ma and contrast with the Leeuwin Complex that records a generally younger Cryogenian to Ediacaran history. Both the Meso- and Neoproterozoic histories of this region have been described as the Pinjarra Orogeny. The overlapping use of this nomenclature has arisen from differing interpretations from the region. Wilde (1999) recognised the importance of 1080–1060 Ma orogenesis within the Northampton and Mullingarra complexes and argued that this event represented a collisional orogeny formed along the margin of Western Australia. He applied the term “Pinjarra Orogeny” to describe this event. Fitzsimons (2003) alternatively suggested that the rocks of the Northampton and Mullingarra complexes were allochthonous with respect to the West Australian margin and were translated into their present position during the Neoproterozoic (750 Ma). On this basis Fitzsimons (2003) argued that it is more logical to use the term “Pinjarra Orogen” for the Neoproterozoic event. This paper follows the nomenclature of Wilde (1999) as it is the author's opinion that the 1080–1060 Ma event most likely represents a collisional orogeny formed in situ along the Western Australian margin as originally envisaged. Although evidence for this collision is not evident in Western Australia, it is more obviously manifested in East Antarctica where both sides of the suture are still present. 345 Where exposed in the Northampton and Mullingarra complexes, Pinjarran orogenesis reworked the middle Mesoproterozoic paragneissic protolith rocks. Deformation occurred at granulite facies conditions and is dated to between 1080 Ma and 1060 Ma, with late- to posttectonic intrusive activity continuing until 990 Ma (Bruguier et al., 1999). More regionally, Pinjarran orogenesis may also have a causal link to the coeval Warakurna large igneous province, an extensive and predominantly mafic intrusive and extrusive suite exposed throughout large areas of central and western Australia (Wingate et al., 2004). South of the Mullingara Complex, the Pinjarra Orogen lies buried beneath the cover rocks of the Perth Basin. From this region these rocks are known only from drill core, but can nevertheless be differentiated from the Leeuwin Complex on the basis of differing Sm– Nd model ages. For the Pinjarra Orogen, these lie between 2.0 and 2.2 Ga, while those of the Leeuwin Complex are consistently b1.5 Ga (Fletcher et al., 1985). In southwestern Australia the Pinjarra Orogen narrows to no more that 70 km where it is fault bound between the eastern margin of the Leeuwin Complex and the western margin of the Yilgarn Craton (Wilde, 1999, Fig. 5). The significance of Pinjarran orogenesis remains contentious. It has been inferred to mark a Grenville age suture between Western Australia and another craton. India was initially proposed as the likely colliding continent (Myers et al., 1996; Bruguier et al., 1999). More recent studies have forwarded both the Kalahari and Congo cratons as alternate possibilities (Powell and Pisarevsky, 2002; Pisarevsky et al., 2003). These options however remain debated. Jacobs et al. (2008) have for example argued that a number of geologic inconsistencies exist with the Kalahari–Western Australia fit. Similarly the most recent synthesis of Rodinia has argued for the more tradition position of the Congo Craton adjacent to the West African and Amazonian cratons at Grenville times (Li et al., 2008). In contrast to the Australia sector where the colliding continent, if one was ever present, has since rifted leaving the present exposures of the Pinjarra Orogen ocean-facing, both sides of the Pinjarra Orogen are in place in East Antarctica (Fig. 6). On the eastern side lie the rocks that have dominated the discussion so far—the rocks of domain 3 with their distinct affinities with southern Australia (Fig. 1). The western side of the suture is defined by rocks of sector 4, those that have no clear affinity with Antarctica's present nearest neighbours and are grouped here into the “Crohn Craton” (Fig. 6a). The name given to this craton is in honour of Peter Crohn, a geologist and member of the 1954 Australian expedition that established the Australian Station at Mawson. He was also one of the team of expeditioners to traverse overland in the winter of 1954 and visit for the first time the Prince Charles Mountains. This range represents one of the principle regions where rocks of the “Crohn Craton” are exposed. The majority of the Crohn Craton presently lies under the Antarctic icesheet, and as such its size, shape, and composition remain largely unconstrained. Nevertheless, if one assumes that the boundaries of this craton are defined in the east by the linear projection of the Pinjarra Orogen through East Antarctica (argued for below) and, in the west by the positions of subsequent Early Cambrian sutures identified in both the Shackleton Range and Prince Charles Mountains, then the Crohn Craton likely defines approximately one third of subglacial East Antarctica (Fig. 6). The principle outcrop localities of the Crohn Craton consist of the limited coastal exposures located west of the Scott Glacier in Princess Elizabeth Land (Fig. 5b) and the southernmost outcrops in the southern Prince Charles Mountains (Fig. 6b). The Scott Glacier outcrops consist of two temporally distinct suites of Archaean orthogneiss. The older of these suites (3005 Ma) records a period of Archaean metamorphism (2890 Ma) that pre-dates the emplacement of the younger (2640 Ma) suite (Sheraton et al., 1992; Black et al., 1992b). In the southern Prince Charles Mountains, the Crohn Craton is defined by both a deformed Archaean basement complex (Ruker Complex), and a series of younger unconformable cover-sequences 346 S.D. Boger / Gondwana Research 19 (2011) 335–371 (Boger et al., 2006; Phillips et al., 2006; McLean et al., 2009). The Ruker Complex consists of the Mawson Suite and Menzies Group, which respectively define a suite of Middle Archaean felsic orthogneisses (3170 Ma) and terrigenous sediments of similar age (b3150 Ma). These rocks were deformed and metamorphosed at mid- to upper-amphibolite facies during the late Archaean (2780 Ma, Mikhalsky et al., 2001b, 2006a; Boger et al., 2006). The Stinear, Ruker, Blake and Sodruzhestvo Groups define the cover-sequences to the Ruker Complex. These rocks were deposited episodically from the early Palaeoproterozoic to the Tonian and were not deformed until the Cambrian (Mikhalsky et al., 2001b; Phillips et al., 2006). The exposed rocks of the Crohn Craton thus suggest that this craton is of Archaean and Palaeoproterozoic age. Nevertheless, the Scott Glacier and Prince Charles Mountains preserve distinct geologic histories and this indicates that the Crohn Craton is composite. These pieces were nevertheless amalgamated prior to Pinjarra orogenesis. Although the majority of the Crohn Craton is unexposed, those parts that do crop out, do not share a common geologic history with the rocks of domain 3. For example, the intrusion of the Mawson Suite in the southern Prince Charles Mountains (3170 Ma) and the two orthogneiss suites in the Scott Glacier region (3005 Ma and 2640 Ma) substantially pre-date the deposition and intrusion of the oldest rocks (2550–2450 Ma) found in the Mawson Craton. Similarly, the long history of Late Palaeoproterozoic and Mesoproterozoic collision and terrane accretion recorded along the western margin of the Mawson Craton is unknown in the Crohn Craton. The inference that the Crohn and Mawson cratons collided is made on the assumption that one can propagate the Pinjarra Orogeny (1080–1060 Ma) across the rifted Australia–Antarctica margin and into central East Antarctica (Fig. 6). Exposed evidence for this orogenic belt in Antarctica is unfortunately scant. The single known Pinjarran age comes from an Archaean orthogneiss exposed in the Obrushev Hills. This rock is exposed immediately west of the continuation of the Darling Fault in Antarctica and has a lower intercept zircon age of 1040 ± 53 Ma (Sheraton et al., 1992). This is the only Pinjarran age presently reported from this area, which otherwise shows substantially more evidence for Ediacaran to Early Cambrian plutonism and structural reworking (570–510 Ma, Black et al., 1992b). This is nevertheless similar to the situation in southwestern Western Australia where the Pinjarra Orogen is limited geographically to a narrow zone between the Leeuwin Complex in the west and the Albany–Fraser Belt in the east (Wilde and Murphy, 1990; Collins, 2003). In Antarctica, the Cambrian overprinting observed west of the Denman Glacier is equivalent to that observed in the Leeuwin Complex, while the rocks exposed in the Bunger Hills are equivalents to those exposed in the Albany–Fraser Belt (Fig. 6b). Projecting the Pinjarra Orogen further into Antarctica would suggest that this belt may may pass close to Lake Vostok (Fig. 6b). This is a region of no outcrop, but one where magnetic and gravimetric data have been collected. These data have been interpreted to reflect the juxtaposition of two crustal blocks of differing crustal thickness and magnetic property (Studinger et al., 2003). Forward modelling of the Vostok gravity data suggests that this boundary was constructed via the east directed underthrusting/subduction of the passive continental margin of the eastern plate (Crohn Craton?) under the western margin of the Mawson Craton (Studinger et al., 2003). Although no age can be ascribed to this duplexing, its position and tectonic style are consistent with the collisional tectonism envisaged for the Pinjarra Orogeny. Beyond Lake Vostok, the along strike continuation of the Pinjarra Orogen brings this belt out on the opposite side of Antarctica in the vicinity of the Shackleton Range (Fig. 6). Recent work in this region has identified granitoid intrusions with ages of 1060 Ma (Will et al., 2009). This age is again comparable to those known from the Pinjarra Orogeny, although admittedly the context of these rocks is presently not well understood. The exposed evidence for the continuation of the Pinjarra orogeny in Antarctica is unquestionably limited. However additional evidence supporting the continuity of this belt comes from the late Mesoproterozoic and Neoproterozoic sedimentary basins that lie on both sides of the inferred trace of the Pinjarra Orogen (Fig. 6b). These basins were shed from the interior of Antarctica and crop out in the southern Prince Charles Mountains and in both the Transantarctic Mountains and the Adelaide Hills region of southern Australia. The Sodruzhestvo Group of the southern Prince Charles Mountains represents the oldest of these deposits. These rocks consist of a predominantly clastic sequence of both carbonaceous and siliceous sediments (Kamenev et al., 1993; Mikhalsky et al., 2001b; Phillips et al., 2005b, 2006). On the basis of microfossil data, these sediments are interpreted to have deposited in the late Meso- or early Neoproterozoic (Iltchenko, 1972), an interpretation consistent with zircon data. The dominant detrital zircon population in these rocks has an age between 1090 Ma and 1040 Ma, identical to the timing of deformation and metamorphism inferred for the Pinjarra Orogeny. The remaining spectra show a broad lower amplitude distribution of ages between 2800 Ma and 2200 Ma, together with a tail of data extending to the Late Archaean (Phillips et al., 2006). The depositional environment of these rocks remains conjectural. Phillips et al. (2006) argued that these rocks were deposited in a continental basin deposited within a united East Gondwana. However, the concept of a united a pre-Gondwana amalgam of India, Antarctica and Australia is not widely accepted. This is due to palaeomagnetic data from the Indian plate that suggest these rocks were separated from East Antarctica until significantly after the Sodruzhestvo Group was deposited (Torsvik et al., 2001; Gregory et al., 2009). Considering that the Sodruzhestvo Group overlies the basement rocks of the Crohn Craton and crops out within 50 km of the first definite exposures of rocks with clear Indo–Antarctic affinities, it is clear that these rocks are exposed close to the proposed early Cambrian suture between domains 2 and 4 (Boger et al., 2001; Boger et al., 2008). With this in mind, the Sodruzhestvo Group was thus more likely deposited near to, or along, the early Neoproterozoic western ocean-facing passive margin of the Crohn Craton (Fig. 6b). The other Antarctic basins were deposited on the opposite side of the continent from the Cryogenian to Cambrian (Fig. 6b). In the sedimentary strata of southeastern Australia, zircons with ages between 1090 and 1060 Ma form a relatively minor component within the Lower Adelaidean Group, but become more volumetrically significant both in the upper part of this group and in the overlying Kanmantoo Group (Ireland et al., 1998). Zircons of this age are also present as a variably important population within middle Cambrian and younger sediments exposed east of the Kanmantoo Group in southeastern Australia (Squire et al., 2006). The temporal correlatives of the upper Adelaidean Group in Antarctica include the Skelton and Koettlitz Groups of southern Victoria Land, the Beardmore Group of the central Transantarctic Mountains and the Hannah Ridge Formation in the Pensacola Mountains (Goodge et al., 2004; Wysoczanski and Allibone, 2004). Both the Skelton and Koettlitz Groups are dominated by 1090– 1060 Ma detrital zircons while in the Hannah Ridge Formation, zircons of this age are second only in importance to a population with ages between 640 and 600 Ma (Goodge et al., 2004; Wysoczanski and Allibone, 2004). The Beardmore Group is exceptional in the respect that it does not contain zircons of this age (Goodge et al., 2002, 2004). Temporal equivalents to the Cambrian Kanmantoo Group in Antarctica are not known from Victoria Land, but are exposed in the central Transantarctic Mountains where they define the lowermost Byrd Group and in the Pensacola Mountains where they define the Patuxent Formation (Rowell et al., 2001; Goodge et al., 2002, 2004). Similar to the Kanmantoo Group, 1090–1060 Ma detrital zircons form one of the more significant fractions within both of these formations (Fig. 6b). S.D. Boger / Gondwana Research 19 (2011) 335–371 9. Neoproterozoic rifting—the formation of the Pacific Ocean The differences in outcrop geology across the proposed trace of the Pinjarra Orogen, the geophysical evidence for collisional tectonics along this belt, and the abundance of late Mesoproterozoic detrital zircons sourced from inland Antarctica, all support: (1) the inference that the Pinjarra Orogen can be extended across the Australo– Antarctic margin and likely transects East Antarctica and (2) likely represents a collisional orogeny that sutured the rocks of the Crohn Craton to western margin of the Mawson Craton in the late Mesoproterozoic. Assuming the above scenario is generally correct, consider the configuration of Laurentia and Antarctica illustrated originally in Fig. 3 and propagated through the subsequent figures. The configuration shown in Fig. 3 was proposed by Betts et al. (2008) and was discussed by these authors for the interval between 1800 and 1600 Ma. It would appear that this reconstruction could have persisted until at least 1400 Ma given the inferred continuation of the Trans–Laurentian igneous suite into Antarctica (Goodge et al., 2004, 2008). Although Goodge et al. (2004, 2008) support a SWEAT configuration, the Betts et al. (2008) arrangement of the continents appears to better align much of the geological data across the Antarctic–Laurentia margin and remains consistent with the findings of Goodge et al. (2004, 2008). Although it is not a goal of this paper to discuss configurations for Rodinia, it is noteworthy that the continuation of the Pinjarra Orogen through East Antarctica links the exposed late Mesoproterozoic of Western Australia with the Grenville Orogen of southern and eastern Laurentia (Fig. 6a). Ages from both the Pinjarra Orogeny and the Ottawan pulse of the Grenville Orogeny are identical (McLelland et al., 2001), while parallels in pre-Grenville geologic history appear consistently manifested in Antarctica, Australia and Laurentia from the Late Palaeoproterozoic to the Late Mesoproterozoic. a Curnamona-Isa SWEAT 347 The long-lived and semi-continuous tectonism observed along the western margin of the Mawson Craton ceased with the collision and suturing of the Crohn Craton. Thereafter tectonic quiescence settled over the proto-Antarctic continent for at least the next 300 million years. During this time Antarctica was presumably a component of Rodinia (e.g. Li et al., 2008; Rino et al., 2008). Tectonism began again in the Cryogenian and in contrast to previous epochs, it was extension rather than compression that was predominant. Evidence for the onset of extension within the Antarctic side of Rodinia is best preserved in Australia and Laurentia. Extension in Australia is marked by the intrusion of the 830 Ma Gairdner dyke swarm and the coeval extrusion of continental flood basalts in the Adelaide Fold Belt (Wingate et al., 1998; Preiss, 2000). Sedimentation occurred within the broad epicontinental Centralian Basin and the fault bounded Adelaide Geosyncline (Walter et al., 1995; Preiss, 2000). In both basins the first 50 million years of sedimentation resulted in the deposition of terrestrial and shallow marine sediments. The first major marine incursion occurred at about 750 Ma in the Adelaidean basin and between 730 and 700 Ma in the Centralian Basin (Walter et al., 1995; Preiss, 2000). Sedimentation was then quasi-continuous from Late Cryogenian until the early to middle Cambrian after which time the Adelaidean basin was flooded by Gondwana derived detritus of the Kanmantoo Group. Although the first marine incursion in the Adelaidean basin dates to the Cryogenian (Preiss, 2000), the appearance of mafic rocks along the eastern margin of Australia (marked by the Tasman Line, Fig. 7) did not occur until the Ediacaran (580 Ma, Crawford et al., 1997; Meffre et al., 2004). These rocks are interpreted to represent the c Tasman Line Centralian Basin i. 830 Ma onset of extension and sediment deposition marked by mafic magmatism ii. NE-SW directed extension iii. Quasi-continuous sediment deposition into the Cambrian iv. Continental separation achieved by 580 Ma? B A Mawson Craton Neoproterozoic (830-550 Ma) sedimentary basins Beardmore A = Adelaide Group NAB = Neoproterozic Antarctic Basins (includes the Skelton, Koettlitz and Beardmore Groups) W = Windermere Supergroup b rifted component of the Curnamona-Beardmore Craton Crohn Craton 2 North American sector i. 780 Ma onset of extension and sediment deposition marked by mafic magmatism ii. Early extension potentially driven by the arrival of a mantle plume between Laurentia, East Antarctica and the Curnamona-Beardmore Craton iii. Quasi-continuous sediment deposition into the Cambrian iv. Continental separation achieved by 580 Ma ? AUSWUS Mawson Australian sector NAB B W Gunbarrel plume 1 W Meosoproterozoic A (1.65-1.55 Ga) Archaean (>2.5 Ga) Meosoproterozoic B (1.5-1.3 Ga) Palaeoproterozoic A (2.3-2.0 Ga) Meosoproterozoic C Trans-Laurentian suite (1.4 Ga) Palaeoproterozoic B (2.0-1.8 Ga) Meosoproterozoic D (1.35-1.15 Ga) Palaeoproterozoic C (1.8-1.65 Ga) Meosoproterozoic E (1.1-1.05 Ga) Belt-Purcell Group Laurentia B = Belt-Purcell Group Fig. 7. Antarctic palaeogeography in the middle Neoproterozoic. (a) SWEAT reconstruction (Moores, 1991) of East Antarctica, Australia and Laurentia prior to the onset of Rodinia rifting. (b) AUSWUS reconstruction (Burrett and Berry, 2000) of East Antarctica, Australia and Laurentia prior to the onset of Rodinia rifting. (c) Proposed rifting of Laurentia, Curnamona–Beardmore and Australo–Antarctic Blocks. Grey outline of Antarctica and Australia shows the modern, but as yet unformed extent of these continents. Numbered regions (1–2) refer to the tectonic sectors of Antarctica shown in Fig. 1. S.D. Boger / Gondwana Research 19 (2011) 335–371 ia Au lia n C r a t o n 5 Mawson Craton Mawson s tra Casey A u s tra lo C ra t o n Davis Azania rth t ic c r a t o n r In d Craton. Admittedly the size of the rifted component of this plate is not constrained, nor is it known where these rocks presently reside. The inferred mode of formation of the Pacific margin of East Antarctica is illustrated in Fig. 7c. The model presented assumes: (1) near orthogonal extension between Laurentia and both East Antarctica and the Curnamona–Beardmore Craton, and; (2) oblique extension between Australia–East Antarctica and the Curnamona– Beardmore Craton. The orientation and timing of extension between Laurentia and East Antarctica/Curnamona–Beardmore is based on the model of Park et al. (1995), which argues that the arrival of the Gunbarrel plume provided the impetus for the onset of extension. The position of the plume is inferred to have been off the present west coast of North America which, during the Cryogenian, is inferred to have been below the approximate triple point between East Antarctica, Laurentia and the Curnamona–Beardmore Craton. These cratons are assumed to have moved radially away from the plume head. Oblique extension between Australia and the Curnamona– Beardmore Craton is inferred from the NW trend of the Gairdner dyke swarm which implies orthogonal SW directed extension (Fig. 7c). The onset of extension in both regions can be constrained from the age of Gairdner dyke swarm (830 Ma, Wingate et al., 1998) in Australia and the age of the earliest phases of magmatism related to the Gunbarrel plume in Laurentia (780 Ma, Harlan et al., 2003). The timing of continental separation is however less well constrained. Between Laurentia and East Antarctica break-up is constrained to have occurred before 570 Ma, after which time tectonic subsidence curves indicate that passive margin sedimentation was underway (Bond et al., 1984; Colpron et al., 2002). A similar timing of continental separation can be implied for the Australian margin a rc -Ant G re a te ar a rw Dh I n d o - A n ta r c tic c r a to n earliest formation of sea floor and imply continental separation was probably not achieved until the end of the Ediacaran (Direen and Crawford, 2003). This is consistent with the observation that the oldest sedimentary rocks overlying seafloor basalts east of the Tasman Line date to the early Cambrian (Squire et al., 2006). The Skelton and Koettlitz Groups of southern Victoria Land and the Beardmore Group exposed in the central Transantarctic Mountains define the Antarctic equivalents Adelaidean Group (Fig. 7). These rocks consist of thick sequences of unfossiliferous sandstone, shale and carbonate, intercalated with glaciogenic diamictite and minor volcanic rocks (Laird, 1991; Stump, 1982, 1995). The constraints on the timing of deposition are provided by detrital zircon data that suggest these rocks were deposited after 650 Ma (Stump et al., 2007). This late Cryogenian maximum depositional age is similar to that implied for the nearby Koettlitz Group as well as for the more southerly Beardmore Group (Goodge et al., 2002, 2004). Similar to the sedimentary rocks in South Australia, these rocks are interpreted to represent both pre-rift and passive margin deposits. What rifted from the margins of Australia and Antarctica depends on ones view of the pre-rift continental configuration. Most commonly this is inferred to be Laurentia, although the present paper would argue that the Australo–Antarctic margin more likely faced the continuation of the Curnamona–Beardmore Craton. This conclusion generally supports that of Li et al. (1995) who have argued that a continent separated Australia and Laurentia within Rodinia, although it is not advocated here that the continuation of the Curnamona–Beardmore Craton formed part of the south China Block (Li et al., 1995). The Australo–Antarctic conjugate is inferred simply to have been the eastward continuation of the Curnamona–Beardmore No 348 2 SVL SV Crohn Craton 5 Kalahari Craton (domain 1) 4 Rayner Belt (990-900 Ma) Undifferentiated Indian Archaean and Palaeoproterozoic rocks 1 G r i c ra Indo-Antarctic Craton (domain 2) ha K a la Kaapvaal and Grunehogna cratons to n Maud-Natal Belt ( 1130-1060 Ma) Coats Land Block (domain 4) Kaapvaal craton Coats Land Block Folded basement with overlying and undeformed Mesoproterozoic volcanic rocks (1110 Ma) Fig. 8. Pre-collision components of Gondwana. Grey outline of Antarctica and Australia shows the modern, but as yet unformed extent of these continents. Numbered regions (1–5) refer to the tectonic domains shown in Fig. 1. S.D. Boger / Gondwana Research 19 (2011) 335–371 given that ocean floor basalts east of the Tasman line did not exist prior to 580 Ma (Direen and Crawford, 2003). Rifting between East Antarctica and rocks of the Curnamona–Beardmore Craton is considered to have occurred prior to 680 Ma (Goodge et al., 2002), although this could arguably have occurred later as there is limited data from this section of the margin. prior to the formation of Gondwana was that of Meert et al. (1995). These authors introduced the name “Kuunga Orogen” and this paper follows their nomenclature. In addition to numerous names there are a number of alternate routes proposed for the Kuunga Orogen. The three main hypotheses are shown in Fig. 9. The earliest of these models is that of Meert and coauthors who envisaged a two-phase assembly of eastern Gondwana (Meert et al., 1995; Meert and Van der Voo, 1997; Meert, 2003). The first phase involved collision between north and central East Africa and what was described as the Slamin terrane, a continent consisting of Madagascar, India and the Antarctic rocks that fall within domain 2. The suture between these blocks is defined by the arc terranes of the East African Orogen along which collision is thought to have occurred between 650 Ma and 620 Ma (Stern, 1994). It was then suggested that subsequent collision (570–530 Ma) merged the Africa–Slamin cratons with Australo-Antarctica (Fig. 9a). Suturing between these blocks occurred along the Kuunga Orogen, an orogenic belt that linked events in the Damara–Zambezi Belt of Africa with those observed in the Prince Charles Mountains–Prydz Bay region of East Antarctica and the Leeuwin Complex of Western Australia (Fig. 9a). In this model the Kalahari Craton was merged into Gondwana the same time as collision along the Kuunga Orogen (Meert, 2003). Although there is a good correspondence of ages along the route proposed by Meert (2003), there remain a number of geographic problems with this model. The proposed route for example transects southern Madagascar and in doing so cross-cuts at high angles the structural grain of the rocks in this region—the inferred route of the Kuunga Orogen trends east, while the structures in southern Madagascar trend north. The north trending structures in Madagascar are attributable to Kuunga age orogenesis and these are considered likely to mark collision between at least Madagascar and IndoAntarctica but also potentially between Madagascar and Africa (Paquette et al., 1994; Kröner et al., 1999; Martelat et al., 2000; Collins et al., 2003; Collins and Pisarevsky, 2005; Boger and Miller, 2004; Tucker et al., 2007). Secondly, to link the events observed along the Damara–Zambezi Belt with those in the Prince Charles Mountains–Prydz Bay region, the route of the Kuunga Orogen must cross a long section of the Antarctic coastline between Lützow–Holm Bay and Prydz Bay, a region that does not preserve evidence for widespread Pan-African reworking (e.g. Boger et al., 2000, 2002; Carson et al., 2000; Kelly et al., 2002; Halpin et al., 2007a). This region is instead 10. Assimilation into Gondwana—the East African–Antarctic and Kuunga Orogenies The break-up of Rodinia and the formation of the Pacific margin along the eastern edge of Australia and Antarctica set in motion a series of events that ultimately lead to convergence and collision along the opposing Crohn margin of the continent (Grunow et al., 1996). It is along this margin that the final additions to East Antarctica were made during the Ediacaran and early Cambrian. These additions include rocks that define domains 1 and 2 (Fig. 1), as well as the rocks of the Coats Land Block, the final component of domain 4. Each of these regions is separated from the others, and from the remainder of East Antarctica, by Ediacaran to early Cambrian aged orogenic belts. These are widely described as part of the “Pan-African” series of events that assembled Gondwana (McWilliams, 1981; Cawood and Buchan, 2007). The rocks of domains 1 and 2 formed parts of larger lithospheric blocks that respectively are defined as the Kalahari and Indo–Antarctic cratons (Fig. 8). The Coats Land Block represents an independent cratonised domain (Kleinschmidt and Boger, 2009) consisting of a low-grade folded basement of unknown age overlain by undeformed volcanic rocks dated at 1110 Ma (Gose et al., 1997). The two most important collisions related to the formation of East Antarctica are those manifested along the Antarctic coast between longitudes 15°W and 40°E, and between longitudes 60°E and 80°E. These belts respectively define the East African–Antarctic Orogen or the Mozambique Belt (e.g. Grunow et al., 1996; Jacobs et al., 1998; Meert and Van der Voo, 1997) and the Kuunga Orogen (e.g. Meert et al., 1995; Meert and Van der Voo, 1997; Meert and Lieberman, 2008; Boger et al., 2001; Fitzsimons, 2000b, 2003). The East African– Antarctic Orogen defines the boundary between the African and Coats Land plates while the Kuunga Orogen defines the boundary between the Indo–Antarctic and Australo–Antarctic plates. For this latter orogen, a number of names are used to describe this belt. The earliest study to recognise the likely separation of India and East Antarctica a 349 b c Rayner Complex AUS S ana ndw ANT N E ast A fri c a n O r o g e n AFRICA SAM I Go ana ndw Kuunga Suture I Go ndw Go AFRICA ANT e tur Su ANT EAO rr a I AUS Ku u a Suture ng P i nja ana AUS DZ SAM E ast A fri c a n O r o g e n AFRICA DZ SAM Fig. 9. Proposed routes for Ediacaran and early Cambrian sutures within Antarctica and southern Africa. (a) Route of the Kuunga suture as proposed by Meert (2003). Kuunga orogenesis post-dates collision between Africa and India along the East African Orogen (EAO). (b) Possible routes of the Pinjarra (= Kuunga) suture as proposed by Fitzsimons (2003). (c) Route of the Kuunga suture as proposed by Boger et al. (2001) and Boger and Miller (2004). DZ = Damara–Zambezi Belt. 350 S.D. Boger / Gondwana Research 19 (2011) 335–371 defined by rocks of clear India affinity (Mezger and Cosca, 1999) and must therefore lie on the Indian side of the Kuunga Suture. Fig. 9b illustrates the possible routes of the Kuunga Orogen as envisaged by Fitzsimons (2003). He proposed two alternative routes, and for each used the south trending boundary of the Mesozoic Lambert Graben (Boger and Wilson, 2003; Lisker et al., 2003; Phillips and Läufer, 2009) to define the western edge of the Kuunga Orogen. Both routes were considered to traverse the Gamburtsev Subglacial Mountains and to terminate either to the north or south of the Shackleton Range. The route to the south of the Shackleton Range intersects the Transantarctic Mountains somewhere between the Miller and Shackleton Ranges. This now seems improbable given the implied continuity of the Nimrod–Kimban Orogeny along the entire eastern margin of the Mawson nucleus (Will et al., 2009). The route running to the north of the Shackleton Range intersects the East African Orogen in western Dronning Maud and is not grossly dissimilar to the route proposed by Boger and Miller (2004). The Boger and Miller (2004) route (Fig. 9c) nevertheless differs in the respect that it honours the structural grain of the rocks in the southern Prince Charles Mountains. Whereas the proposed route of Fitzsimons (2003) trends south, structures in the southern Prince Charles Mountains trend west to southwest, an orientation known from surface observations (Boger and Wilson, 2005; Phillips et al., 2005a,b; Corvino et al., 2008) as well as from airborne geophysical data (McLean et al., 2009). The Boger and Miller (2004) route, although correct for the Prince Charles Mountains region, nevertheless fails to adequately incorporate the similarly aged orogenesis observed in the Shackleton Range. Kleinschmidt and Boger (2009) highlighted the likely continuity of orogenesis observed in both of these regions. The following paragraphs outline my preferred interpretation of how the final components East Antarctica were added to the Australo–Antarctic core of the continent. Events are divided into Ediacaran, Ediacaran–early Cambrian and Cambrian, events that are considered likely to represent spatially separate phases of collision. Within the resolution of the presently available geochronology, the ages from these events overlap to varying degrees. However, when the mean ages from each of these belts are considered, an older to younger progression is present. It is likely that many of the described orogenic belts were active at the same time, with the terminal stages of the older events overlapping with the early stages of the younger events. Although the evolution described below may in time be proved more or perhaps less correct, the key points are that: (1) the rocks of domains 1, 2 and the Coats Land part of domain 4 were sutured into their modern positions in Antarctica as a result of the formation of Gondwana, (2) these events were all closely spaced in time and, (3) these rocks were first sutured to the margins of West Gondwana (Africa and South America) before ocean closure along the Kuunga Orogen sutured these rocks to the Crohn margin of the Australo–Antarctic plate. 10.1. Ediacaran events (580–550 Ma) The process of Gondwana amalgamation began with the collision of the rocks of domain 1 with those of the Coats Land Block (Fig. 8). The rocks that define domain 1 crop out in Dronning Maud Land between longitudes 15°W and 30°E and include the Archaean Grunehogna Province as well as the Mesoproterozoic rocks of the Maud Belt (Fig. 1). The Grunehogna province consists of a poorly exposed granitic basement of Archaean age that, on the basis of Rb–Sr and 207Pb/204Pb data, is thought to be between 2960 Ma and 2820 Ma old (Halpern, 1970; Barton et al., 1987). These rocks are overlain by the Mesoproterozoic Ritscherflya Supergroup, a flat-lying sequence of both sub-aqueous and sub-aerial sediments that are intercalated with basaltic to andesitic lavas and lesser volcaniclastic sediment (Watters et al., 1991). The lavas are geochemically similar to the shallowly emplaced and temporally equivalent Borgmassivet intrusions that form large and widespread sills within the Ritscherflya Supergroup (Krynauw et al., 1988). A tuff intercalated within the Ritscherflya Supergroup has yielded a U–Pb zircon age of 1135 Ma (unpub. data of Knoper and Tucker cited by Powell et al., 2001 and Frimmel, 2004), similar to the detrital zircon ages obtained from the intercalated terrigenous sediments (Perrit, 2001). Very similar geochemistry and palaeomagnetic poles for both the well-dated Umkondo dolerites of southern Africa (1110 Ma) and the Borgmassivet intrusions of the Grunahoga Province (Jones et al., 2003; Basson et al., 2004; Gose et al., 2006), together with similarities in basement ages from the Grunehogna Province and that of the Kaapvaal Craton, have led to the conclusion that these regions were continuous from at least the middle Mesoproterozoic and formed part of a combined Grunehogna– Kaapvaal Craton (Groenewald et al., 1995). This is referred to more generally as the Kalahari Craton. The margin of this craton, both in southern Africa and in Antarctica, is defined by the Maud–Natal Belt (Jacobs et al., 1993, Fig. 9). The protolith rocks from the Maud–Natal Belt are defined by predominantly juvenile calc-alkaline rocks of both volcanic and intrusive origin that were emplaced between 1160 Ma and 1130 Ma and inferred to have formed in a volcanic arc (Arndt et al., 1991; Jacobs et al., 1996, 1998, 2003b; Bauer et al., 2003; Paulsson and Austrheim, 2003; Board et al., 2005). The formation of this arc is interpreted to have occurred either along the margin of the Kalahari Craton (Bisnath et al., 2006; Grosch et al., 2007), or as an intra-oceanic arc that was accreted to the margins of this craton during subsequent orogenesis (Groenewald et al., 1995; Bauer et al., 2003). The former of these interpretations is based on the similarity in age of detrital zircons found within the Ritscherflya Supergroup and the volcanic rocks exposed in the Natal–Maud arc (Frimmel, 2004). These data were argued to require the Ritscherflya Supergroup to have been proximal to the Maud–Natal arc at the time of sediment deposition (Basson et al., 2004). The alternate interpretation arguing that the Natal–Maud arc formed outboard of the Kalahari Craton was based on the absence of 1160–1130 Ma intrusions in the Kalahari Craton (Groenewald et al., 1995; Bauer et al., 2003). Subsequent to the formation of the Maud–Natal arc, these rocks underwent high-grade metamorphism at 1090–1030 Ma and again at 580–520 Ma (Arndt et al., 1991; Jacobs et al., 1998, 2003a,b; Bauer et al., 2003; Board et al., 2005; Bisnath et al., 2006). The earlier of these events is generally regarded as having resulted in collision and ocean closure between the Kalahari Craton and what is generally implied to be the Coats Land Block (Bauer et al., 2003; Grosch et al., 2007; Jacobs, 2009). However, the palaeomagnetic poles from these regions show a mismatch of approximately 30° at 1100 Ma (Gose et al., 1997), a time that does not significantly precede the implied onset of collision (1090 Ma). It thus seems more likely that although the Maud–Natal Belt was tectonically active between 1090 and 1030 Ma, orogenesis did not lead to ocean closure. The subsequent Ediacaran to Early Cambrian event occurred in two pulses. The first is marked by the localised intrusion of anorthosite and charnockite at approximately 605 Ma, followed by deformation and metamorphism between 580 Ma and 540 Ma (Jacobs et al., 1998, 2003a; Board et al., 2005). This event was the result of crustal convergence and was accompanied by high-temperature metamorphism characterised by decompression from moderate to high pressures (Jacobs et al., 2003a; Board et al., 2005). The second pulse was mostly of magmatic character and was marked by the intrusion of large volumes of post-tectonic granite between 530 Ma and 490 Ma (Ohta et al., 1990; Mikhalsky et al., 1997; Jacobs et al., 1998, 2003a; Roland, 2002; Paulsson and Austrheim, 2003). These intrusions show distinctive within-plate geochemistry and accompanied a period of extensional deformation (Jacobs et al., 2003c) associated with the southward extrusion of a number of microplates which now define the Falkland, Ellsworth–Haag, and Filchner Blocks (Jacobs and Thomas, 2004). S.D. Boger / Gondwana Research 19 (2011) 335–371 manifested in the Coats Land Block either. Ediacaran collision between the Coats Land Block and the Maud–Natal margin of the Kalahari craton is supported by the increasing intensity of Ediacaran deformation and metamorphism toward the margin of the Coats Land Block (Jacobs et al., 2003a,b), while the assumption that these rocks were not already part of East Antarctica can be inferred from the younger phase of orogenesis manifested on the opposite southern and western side of Coats Land exposed in the Shackleton Range. Deformation here involved subduction of oceanic lithosphere and eclogite facies metamorphism (Schmädicke and Will, 2006; Romer et al., 2009) indicating that this region likely forms a plate boundary. Orogenic ages similar to those observed in Dronning Maud Land are also observed in southern and western Madagascar and in parts of southern India. In Madagascar, Ediacaran to Cambrian age convergent deformation occurred between 580 Ma and 530 Ma and overprints at various metamorphic grades most rocks in the island (Paquette et al., 1994; Kröner et al., 1999; Martelat et al., 2000; Tucker et al., 2007). With the exception of the Vohibory domain of southwestern Madagascar (de Wit et al., 2001; Jöns and Schenk, 2008), orogenic ages N600 Ma are absent in Madagascar. These ages (N600 Ma) equate with events recognised in the juvenile Neoproterozoic arc terranes of The significance of these Ediacaran and Early Cambrian events in Dronning Maud Land remains debated. Grosch et al. (2007) for example imply that the eastern margin of the Kalahari craton, marked by the Natal–Maud Belt, was ocean-facing prior to the onset Ediacaran orogenesis. They further suggest that this belt marks the suture between the Kalahari Craton and the combined Australo–Antarctic continent, the leading edge of which they implied to be defined by the Coats Land Block. Other studies suggest that the Coats Land Block had already docked with the Maud–Natal Belt (Bauer et al., 2003; Jacobs, 2009) and place the suture further to the east within the Shackleton Range. Alternatively the Coat Land Block collided with the Maud– Natal margin of the Kalahari Craton in the Ediacaran, but at the time represented a separate and unrelated micro-continent unattached to the remainder of East Antarctica (Kleinschmidt and Boger, 2009). This latter interpretation is supported by a number of observations. Firstly, by assuming that the Maud–Natal Belt remained ocean-facing in the Mesoproterozoic accounts for the large latitudinal separation that existed between the Coats Land Block and the Kalahari craton just prior 1090–1030 Ma orogenesis. It also explains the lack of deformation of this age observed in the Coats Land Block—although this is not unique to the Mesoproterozoic event, Ediacaran deformation is not rt h s tr a lia n C r a t o n 5 a rc t Casey Mawson Craton A u s tra lo - A n t Dha rw a rC d ia vis Da on ws Ma Long-lived continental arc within the Madurai Block some evidence for metamorphism between 580-550 Ma on ra t Au No ic c r a t o n n ta ra to n a t e r In Continuing ocean closure between West Gondwana (Precambrian Africa & South America) and: (i) Indo-Antarctica (ii) the combined Kalahari-Coats Land craton A o- cc rc ti G re In d 580-550 Ma 351 2 Crohn Craton Collision: Azania and Precambrian Africa (~560 Ma) Collision: Coat Land Block and Kalahari Craton (~560 Ma) 5 Azania 4 1 Coats Land Block G Precambrian Africa ri c ra to n Kaapvaal craton a ah K al Kalahari Craton (domain 1) Maud-Natal Belt ( 1130-1060 Ma) Kaapvaal and Grunehogna cratons 5 Precambrian South America 5 Indo-Antarctic Craton (domain 2) Rayner Belt (990-900 Ma) Undifferentiated Archaean and Palaeoproterozoic rocks Coats Land Block (domain 4) Folded basement with overlying and undeformed Mesoproterozoic volcanic rocks (1110 Ma) Fig. 10. Antarctic palaeogeography in the Ediacaran. Collisions and/or subduction highlighted with large white stars. Collision unites Azania with West Gondwana (pre-600 Ma Africa and South America) and the Coats Land Block with the Kalahari Craton. Continental arc (Madurai Block) was active offshore from the Indo–Antarctic margin. Grey outline of Antarctica and Australia shows the modern, but as yet unformed extent of these continents. Numbered regions (1–5) refer to the tectonic domains shown in Fig. 1. 352 S.D. Boger / Gondwana Research 19 (2011) 335–371 East Africa (Stern et al., 2006), which are separated from the Archaean Antananarivo domain of central and northern Madagascar by rocks tectonised between 580 Ma and 530 Ma. With this in mind, 580– 530 Ma orogenesis likely marked the suturing of central and northern Madagascar with the remainder of Africa (Fig. 10). Arguably central and northern Madagascar was at this time part of a micro-continent called Azania, which included parts of Somalia, Ethiopia and Arabia (Collins and Pisarevsky, 2005). This conclusion however remains contentious with other authors instead arguing that the Late Archaean rocks of Madagascar were at the time contiguous with the Indo– Antarctic plate (e.g. Tucker et al., 1999). In southern India, a long-lived continental arc defined by the Madurai Block (Fig. 10) is considered to have been active offshore from the margin of the Indo–Antarctic plate during the Ediacaran (Santosh et al., 2009). Some plutonism and associated metamorphism within this arc was contemporaneous with the first phase of collisions recognised between the Kalahari–Coats and Azania–Africa blocks (Bartlett et al., 1998; Santosh et al., 2003). 10.2. Ediacaran to Early Cambrian events (550–520 Ma) The accretion of the Coats Land Block onto the margin of the Kalahari Craton assembled the coastal margins of East Antarctica between Dronning Maud and Princess Elizabeth Lands (Fig. 1). However, these rocks had not as yet attained their modern configuration, as equally significant events were to follow from the end of the Ediacaran to the early Cambrian. Orogenesis during this interval was manifested in two unrelated orogenic belts that resulted in the incorporation of both the Kalahari–Coats and Indo–Antarctic plates into West Gondwana. These belts are defined by the Damara–Zambezi Orogen of southern Africa, and by an as yet unnamed orogenic belt that extends from the Lützow– Holm region of East Antarctica, through Sri Lanka, southern India and, assuming the existence of the Malagasy Azania plate, into eastern Madagascar. These belts respectively mark the sutures along which the Kalahari–Coats Land block amalgamated with the northern half of Africa (Barnes and Sawyer, 1980; Meert, 2003; Johnson et al., 2005) and where the Indo–Antarctic plate collided with central and northern Madagascar and with the Maud–Natal margin of the Kalahari Craton (Fig. 11). Ocean closure along the Damara–Zambezi Orogen is indicated by relic ocean floor found within the orogen. These rocks were metamorphosed locally to eclogite facies, an observation that points to the deep underthrusting consistent with plate margin tectonics (John et al., 2003, 2004a; John and Schenk, 2003). The timing of highpressure metamorphism is constrained to between 660 Ma and 585 Ma (John et al., 2004b) and is interpreted to reflect pre-collision subduction beneath the Kalahari Craton (John et al., 2004a). Collisional orogenesis is argued to be somewhat younger and is constrained broadly to between 615 Ma and 520 Ma with the bulk of age data Fig. 11. Antarctic palaeogeography in the Ediacaran and early Cambrian. Collisions highlighted with large white stars unite: (1) the Indo–Antarctic plate with West Gondwana (pre600 Ma Africa, South America and recently accreted Azania) and, (2) the Kalahari Craton with West Gondwana. Grey outline of Antarctica and Australia shows the modern, but as yet unformed extent of these continents. Numbered regions (1–5) refer to the tectonic domains shown in Fig. 1. S.D. Boger / Gondwana Research 19 (2011) 335–371 yielding a mean between 550 and 520 Ma (Johnson et al., 2005). It is this latter interval that is interpreted to date suturing (Fig. 11). Evidence for the broadly coeval collision between the Indo– Antarctic and West Gondwana plates is to be found in the Lützow– Holm region of East Antarctica, peninsula India and in eastern Madagascar (Fig. 11). The Lützow–Holm region includes rocks of the Sør Rondane Mountains and Lützow–Holm Bay exposed between latitudes 25°E and 40°E. Rocks in these two regions show remarkably consistent geologic histories. Deformation and metamorphism are constrained to between 550 and 510 Ma with a mean age of around 530 Ma. Metamorphism accompanying this event occurred at mostly granulite facies with peak P–T conditions of around 800–900 °C and 11 kbar (Shiraishi et al., 1994, 2003; Fraser et al., 2000; Asami et al., 2005; Kawasaki et al., in press). Secondary mineral assemblages indicate post-peak decompression (Shiraishi et al., 1994; Motoyoshi and Ishikawa, 1997; Fraser et al., 2000). 530-500 Ma Although metamorphic events in Dronning Maud Land are commonly linked with those in Madagascar via the rocks exposed in Lützow–Holm Bay (Grunow et al., 1996; Wilson et al., 1997; Jacobs et al., 1998; Jacobs and Thomas, 2004), correlation of these areas is, in the author's opinion, still uncertain. Whereas orogenesis in Dronning Maud Land was bimodal, consisting of a collisional phase between 580 Ma and 540 Ma and an extensional phase between 530 Ma and 490 Ma (Jacobs et al., 2003a,c), orogenesis in the Lützow–Holm region was unimodal and marginally younger (530 Ma). Of equal significance is that the Antarctic hinterlands to these regions differ. In Dronning Maud Land the Coats Land Block defines the colliding plate. In the Lützow–Holm region, the collider is defined by the Rayner Complex— rocks that have clear affinities to the pre-Gondwana Indo–Antarctic plate (Fig. 11). Orogenesis in Dronning Maud Land is thus interpreted to mark the collision of the Kalahari Craton and the Coats Land Block, while b Prince Charles Mountains-Prydz Bay Kuunga Suture 70 ˚S a 353 70˚ E r a l ia n C r a t o n ia GM Convergence and collision Prince Charles Mountains and Shackleton Rangge 5 SR ie r E 100 200 km Leading edge of the Australo-Antarctic Plate 60˚E Sodruzhestvo Group - Late Mesoproterozoic sediments Rayner Complex - major Cambrian overprint Common orogenic and intrusive history between 990 and 900 Ma Fisher Complex Ruker and Blake Groups Palaeoproterozoic sediments Ruker Complex - Archaean magmatic and orogenic history Lambert Complex Kaapvaal craton 70˚ 65˚E Rayner Complex - minor Cambrian overprint G Precambrian Africa G la c 0 Leading edge of the Indo-Antarctic Plate 4 1 75˚ E 74˚S 60˚E 70˚S 2 SVL SV 72˚S a to n ar C r Crohn Craton Mawson e rt b PCM Mawson Lam Davis a rw 80 ˚E Casey 3 Dh Amery Ice Shelf 65˚E 68˚S r In d 74˚ S Prydz Bay 5 Mawson Craton G re a te 72 ˚S st Au No rt h Davis First common history 530-490 Ma c Lateral extrusion and within plate magmatism (Dronning Maud Land) Precambrian South America Shackleton Range Leading edge of the Australo-Antarctic Plate Basement Read Group Pioneers and Stratton Groups Pan-African Orogenic Belts Neoproterozoic to Cambrian rocks Common (Kimban) orogenic history between 1710 and 1680 Ma Ophiolite Complex - major Cambrian deformation Early Cambrian Wastts Neddle and Mt Wegener Formations Eastern terrain - 1060 plutonism (Pinjarran magmatism) Cambrian collisional belts (530-490 Ma) 80˚45’S Basement terranes Indo-Antarctic Craton (sector 2) Ediacaran collisional belts (580-550 Ma) Rayner Belt (990-900 Ma) Basement terrains Kalahari Craton (sector 1) Maud-Natal Belt ( 1130-1060 Ma) Kaapvaal and Grunehogna cratons Kuunga Suture Ediacaran-Cambrian collisional belts (550-520 Ma) 80˚15’S Undifferentiated Archaean and Palaeoproterozoic rocks Coats Land Block (sector 4) Folded basement with overlying and undeformed Mesoproterozoic volcanic rocks (1110 Ma) Leading edge of the Coats Land Block Northern Gneiss Complex - major Cambrian overprint 20˚W 0 22˚W 24˚W 26˚W 28˚W 20 km 30˚W Fig. 12. Antarctic palaeogeography in the early to middle Cambrian. (a) Collision along the Kuunga Suture unites Gondwana. White stars highlight regions where the Kuunga Suture is exposed. PCM = Prince Charles Mountains, SR = Shackleton Range. Grey outline of Antarctica and Australia shows the modern, but as yet unformed extent of these continents. Numbered regions (1–5) refer to the tectonic domains shown in Fig. 1. (b) Enlargement of the Prince Charles Mountains highlighting the geology of the opposing Indo–Antarctic and Australo–Antarctic plate margins. (c) Enlargement of Shackleton Range highlighting the geology of the opposing Australo–Antarctic and Coats Land plate margins. 354 S.D. Boger / Gondwana Research 19 (2011) 335–371 orogenesis in the Lützow–Holm region likely represents collision between the western margin of the Rayner Complex and the already joined Kalahari–Coats Land Craton. The suture between these latter regions lies between the Sør Rondane Mountains and Lützow–Holm Bay and is argued by Asami et al. (2005) to also include a suite of Neoproterozoic sediments deposited on one or both of the passive margins prior to collision. The trace of the Lützow–Holm suture is implied to continue through Sri Lanka and into eastern Madagascar via southern India (Fig. 11). The Sri Lankan crust is divided into three domains on the basis of differing Nd crustal residence and U–Pb intrusion ages. Of these, the northern and western Wanni Complex consists of protolith rocks with crustal residence ages between 2.0 Ga and 1.0 Ga that were deformed and metamorphosed between 1010 Ma and 890 Ma (Milisenda et al., 1988; Kröner et al., 2003). These data show strong similarities to the orogenic and model age datasets from the Rayner Complex of East Antarctica, a finding that suggests these two regions correlate (Black et al., 1987; Young and Black, 1991; Shiraishi et al., 1994; Young et al., 1997; Zhao et al., 1997a; Willbold et al., 2004; Corvino and Henjes-Kunst, 2007). The Sri Lankan Wanni Complex thus likely defines the leading edge of the Indo–Antarctic plate. The adjacent Highland Complex consists of older rocks which have TDM ages that range from 3.4 to 2.0 Ga (Milisenda et al., 1988). No model age data has been published from the Lützow–Holm Complex, although petrological similarities suggest that these regions may correlate (Ogo et al., 1992; Shiraishi et al., 1994; Osanai et al., 2006). The extension of the Highland Complex toward India is not certain, although the Highland Complex shows a number of similarities with the Madurai Block. Both regions have for example similar Nd model ages (3.2–2.1 Ga) and detrital zircon spectra, and appear to record slightly older ages of metamorphism (610–550 Ma) when compared to the neighbouring terranes (Kröner et al., 1987; Braun and Kriegsman, 2003; Santosh et al., 2009). The southern and eastern Vijayan Complex of Sri Lanka is similarly considered to be arc related, but differs in the respect that it is defined by Late Mesoproterozoic (1030–1020 Ma) mostly juvenile metaigneous rocks (Kröner et al., 2003). These rocks are interpreted to have formed in an intra-oceanic arc, a conclusion consistent with this terrane having formed on the oceanic plate between the Indo– Antarctic and African continental margins prior to Gondwana amalgamation. The continuation of the Lützow–Holm suture within Sri Lanka thus most likely lies along the boundary between the Wanni and Highland complexes. Within southern India this structure is interpreted to lie between the Madurai Block and the southern margin of the Dhawar Craton where it is defined by the Palghat–Cauvery Shear Zone (Clark et al., 2009a,b; Santosh et al., 2009). The geology of southern India and Sri Lanka is however complicated, and it is probable that the broad collision between the West Gondwana and Indo–Antarctic plates involved the incorporation of a number of microplates. Consequently, a number of subordinate sutures are also likely to be present and these potentially separate the Highland and Vijayan complexes in Sri Lanka and the Madurai and Trivandrum blocks within southern India. These latter rocks are exposed at the southern tip of India and consists mostly of now metamorphosed continental margin sediments that have detrital zircon characteristics common with similarly metamorphosed sedimentary rocks exposed in southeastern Madagascar (Collins et al., 2007). The trace of the Lützow–Holm suture is arguably narrower in eastern Madagascar where it is described as the Betsimisaraka suture (Collins, 2006; Collins and Windley, 2002; Collins et al., 2003; Collins and Pisarevsky, 2005). This suture defines the eastern limit of the Malagasy 810–770 Ma Imorona–Itsindro suite (Handke et al., 1999), contains podiform ultramafic rocks of potentially ocean floor affinity, and marks the boundary between sedimentary rocks with East African and Indian affinities (Collins and Windley, 2002; Collins et al., 2003). The opposing view is presented by Tucker et al. (1999, in press) who highlight the age and isotopic similarities between the Antananarivo domain of central and northern Madagascar and the Dhawar craton of western India—rocks which are exposed on either side of the Betsimisaraka suture. They argue that these regions were contiguous since Late Archaean times. If correct, this conclusion implies that the Lützow-Holm suture probably continues to the west of the Antananarivo bock through southern and western Madagascar where there is also widespread evidence for Latest Neoproterozoic and Cambrian aged deformation and metamorphism. 10.3. Cambrian events (530–490 Ma) The youngest and final phase of collision merged the Australo– Antarctic plate with West Gondwana (Fig. 12). It is this event that ultimately sutured the rocks that now define domains 1, 2 and the Coats Land Block section of domain 4 into their modern configurations. This collision occurred along the Kuunga Orogen, which can be traced from the Denman Glacier region of Princess Elizabeth Land through the Prince Charles Mountains to the Shackleton Range (Boger et al., 2001; Kleinschmidt and Boger, 2009). In the Prince Charles Mountains the Kuunga suture separates rocks with Indian affinities (domain 2) from those of the Crohn Craton (domain 4). These latter rocks define the Ediacaran margin of the Australo–Antarctic plate (Fig. 11). In the Shackleton Range, the Kuunga suture separates rocks of the Mawson and Coats Land blocks, regions that respectively represent the parts of the Australo–Antarctic and Afro-Antarctic plates (Fig. 11). In the Prince Charles Mountains rocks which lie on the Indo– Antarctic plate include those exposed in the Rayner, Fisher and Lambert complexes (Fig. 12). Of these the Rayner Complex is the one most clearly of Indian affinity, having correlative rocks exposed in the Eastern Ghats of India (Mezger and Cosca, 1999). It is also the most extensive and can be traced from the Lützow–Holm region at longitude 45°E in the west, eastwards and inland of the Archaean Napier Complex and into the Lambert Basin at longitude 70°E. The protolith rocks of the Rayner Complex, although undated, are likely to be of Mesoproterozoic age based on the prevalence of Nd model ages of between 2.2 Ga and 1.5 Ga (Black et al., 1987; Young et al., 1997; Zhao et al., 1997a; Corvino and Henjes-Kunst, 2007). These rocks were deformed during a single high-grade orogenic event that was accompanied by widespread charnockitic and granitic magmatism dated to 990–900 Ma (Grew et al., 1988; Young and Black, 1991; Kinny et al., 1997; Boger et al., 2000; Carson et al., 2000; Kelly et al., 2002; Halpin et al., 2007a). The metamorphic evolution of this belt is variable, with clockwise P–T–t paths with peak pressures of 9–10 kbar and 850–900 °C described in the north and west adjacent to the Napier Complex, and counter-clockwise P–T–t paths with lower peak pressures (5–7 kbar) but similar peak temperatures (850– 900 °C) prevalent throughout the rest of the belt (Clarke et al., 1989; Thost and Hensen, 1992; Fitzsimons and Harley, 1992; Stephenson and Cook, 1997; Boger and White, 2003; Kelly and Harley, 2004; Halpin et al., 2007b). The Rayner Complex has commonly been terminated against the Lambert Graben and overlying glacier of the same name. This feature geographically separates the Neoproterozoic granulites of the Rayner Complex from the rocks from the Prydz Bay coast. These latter rocks differ in the respect that they record widespread evidence for Early Cambrian age orogenesis and high-grade metamorphism (Zhao et al., 1992; Carson et al., 1995, 1996; Fitzsimons et al., 1997; Kelsey et al., 2003), an event that in the Rayner Complex is limited to the development of localised shear zones and the emplacement of pegmatitic dykes under relatively low-temperature conditions (Grew, 1978; Black et al., 1983; Clarke, 1988; Manton et al., 1992; Carson et al., 2000; Boger et al., 2002). Although strongly overprinted, the protolith rocks of the Prydz Bay coast would nevertheless appear S.D. Boger / Gondwana Research 19 (2011) 335–371 to correlate with those observed in the Rayner Complex. The Prydz Bay protoliths are defined by Mesoproterozoic orthogneisses that were intruded sporadically between 1380 Ma and 1020 Ma (Carson et al., 2007; Liu et al., 2007b, 2009; Wang et al., 2008), an age range similar to that which can be implied for the Rayner Complex protoliths. Model ages (2.2–1.5 Ga) from both areas overlap (Sheraton et al., 1984; Zhao et al., 1995), while Rayner age metamorphism and plutonism (990–900 Ma) is also described from the Prydz Bay (Hensen and Zhou, 1995; Zhao et al., 1995; Tong et al., 2002; Carson et al., 2007; Kelsey et al., 2007; Wang et al., 2008; Liu et al., 2009). The only significant lithologic difference between the two regions is that the Prydz Bay region consists of a felsic composite orthogneissic basement (Søstrene Orthogneiss) of probable Rayner origin and an overlying and younger cover-sequence (Brattstrand Paragneiss). These latter rocks have no known equivalents of the west of the Lambert Glacier. On the basis of detrital zircon data the Brattstrand Paragneiss is taken to have been deposited in the Ediacaran and are inferred to have an Indo–Antarctic origin, an interpretation consistent with their deposition close to the margin of the Indo–Antarctic plate (Zhao et al., 1995; Kelsey et al., 2008). The Mesoproterozoic protolith age and bimodal orogenic history observed in Prydz Bay is replicated in the rocks of both the Grove Mountains and central Prince Charles Mountains (Fig. 12). In the Grove Mountains, the oldest presently known rocks are of Rayner age and consist of 930–900 Ma felsic and mafic orthogneisses with TDM model ages between 2.5 Ga and 1.6 Ga (Liu et al., 2007a). These are in turn intruded by widespread syn-tectonic to late-tectonic charnockite and granite plutons with ages between 540 Ma and 500 Ma (Mikhalsky et al., 2001a; Liu et al., 2006). Metamorphism was at granulite facies and occurred at nearly identical P–T conditions (6.1– 6.7 kbar and 850 °C) to those observed at Prydz Bay (Liu et al., 2003; Zhao et al., 2003). The rocks of the central Prince Charles Mountains (e.g. Clemence and Shaw Massifs) are characterised by widespread 1090–1050 Ma pre-orogenic orthogneisses, similar in age to those observed in southern Prydz Bay. The TDM ages of these rocks are also between 2.2 Ga and 1.5 Ga and within the range observed in the Rayner Complex (Corvino and Henjes-Kunst, 2007). The zircon systematics from this region preserves evidence for both 990– 900 Ma and at 550–500 Ma metamorphism (Corvino et al., 2005; Maslov et al., 2007). Given the isotopic similarities among the Grove Mountains, central Prince Charles Mountains, Prydz Bay and the Rayner Complex senso stricto, it is the author's opinion that these regions all form parts of the same variably reworked Early Neoproterozoic orogenic belt. Within the Lambert basin the Rayner Complex can thus be traced to least 72°S, and is exposed on both sides of the Lambert Glacier (Fig. 12b). The only significant regional difference lies in the degree of Cambrian reworking, which is most intense in the south and east. The Fisher and Lambert complexes (Fig. 12b) are geologically distinct. The Fisher Complex is a relatively narrow belt of greenschist to amphibolite faces mafic to intermediate volcanic and minor metasedimentary rocks intruded by a variety of subvolcanic plutons. Ages for these rocks are from 1300 Ma to 1280 Ma (Beliatsky et al., 1994; Kinny et al., 1997; Mikhalsky et al., 1999). Both the plutonic and volcanic rocks have chemistries suggestive of their formation in a Mesoproterozoic island arc (Mikhalsky et al., 1996), while amphibolite facies metamorphism and associated granite intrusion is dated at 1020–990 Ma ages that can be broadly equated with those observed in the Rayner Complex (Mikhalsky et al., 1999). The Lambert Complex is defined by markedly older crust that dates mostly to the Palaeoproterozoic. The oldest component of this complex is defined by middle Archaean orthogneiss dated at 3520 Ma (Boger et al., 2008). These rocks are tectonically interleaved with metasedimentary rocks and felsic orthogneisses of early Palaeoproterozoic age. The oldest of these are of granodioritic to synogranitic composition and were emplaced between 2490 Ma and 355 2410 Ma (Mikhalsky et al., 2006a; Corvino et al., 2008). TDM model ages for these rocks are between 3.4 Ga and 3.0 Ga (Mikhalsky et al., 2006b). The second suite of felsic orthogneisses is of granitic composition and was emplaced between 2180 Ma and 2110 Ma (Boger et al., 2008; Corvino et al., 2008). Limited model age data (TDM = 2.87) suggest that these rocks were sourced from younger protoliths when compared to the older orthogneiss suite, but this source was nevertheless substantially older than that observed in the Fisher or Rayner complexes. Deformation and metamorphism in the Lambert Complex was bimodal. In the north, zircon and monazite data suggest that the main phase of deformation occurred between 930 Ma and 900 Ma, with a Cambrian (530–500 Ma) overprint becoming more prevalent in the south (Boger et al., 2001, 2008; Corvino et al., 2008; Phillips et al., 2009). Importantly both the Fisher and Lambert complexes record evidence of the early Neoproterozoic (1000–900 Ma) event common to the Rayner Complex. This shared early Neoproterozoic history implies that these three terranes were amalgamated by this time. Although the process by which this occurred has not been described, their common early Neoproterozoic history ties the Rayner, Fisher and Lambert complexes to the Indo–Antarctic plate prior to Kuunga age collision. The position of the Kuunga Suture within the Lambert Basin is thus defined by the juncture between the most southerly exposures of rocks which record Rayner aged (990–900 Ma) orogenesis, and those of the Crohn margin of the Australo–Antarctic plate, which do not record this event. This boundary is exposed in the southern Prince Charles Mountains along the southern margin of the Lambert Complex (Fig. 12b). Here the Palaeoproterozoic Lambert Complex is juxtaposed against the older and unrelated rocks of the Ruker Complex (Boger et al., 2001). The Ruker Complex forms the oldest component of the Crohn margin of the Australo–Antarctic plate and consists of an intercalated sequence of Middle Archaean felsic orthogneisses and similarly aged metasediments that were deformed and metamorphosed in the Late Archaean (Mikhalsky et al., 2001b, 2006a; Boger et al., 2006). Cover rocks to the Ruker Complex include the Stinear, Ruker, Blake and Sodruzhestvo Groups that were deposited from the early Palaeoproterozoic to early Neoproterozoic (Mikhalsky et al., 2001b; Phillips et al., 2006). With the exception of some detrital zircon ages, there is no overlap in the geologic histories of rocks of the Indo– Antarctic (Lambert–Fisher–Rayner complexes) and Australo–Antarctic (Ruker Complex plus its cover-sequences) margins prior to the Cambrian (Boger et al., 2008; Corvino et al., 2008). From the Cambrian onwards however, rocks from both sides of the Kuunga Suture record a common geologic history (Fig. 12b). Based on the U–Pb zircon and monazite age data from the Lambert Complex, orogenesis on the Indo–Antarctic side of the Kuunga Suture is constrained to between 530 Ma and 490 Ma (Boger et al., 2001, 2008; Mikhalsky et al., 2006a; Corvino et al., 2008; Phillips et al., 2009). Peak pressures and temperatures are estimated at 6–7 kbar and 600–700 °C with retrograde decompression describing a clockwise P–T–t path (Boger and Wilson, 2005; Phillips et al., 2009). Similar pressures, but markedly higher temperatures were recorded more distal from the plate margin in Prydz Bay (Fitzsimons and Harley, 1992; Thost et al., 1994; Carson et al., 1997; Zhao et al., 1997b). This may reflect higher pre-orogenic heat flow inboard from the plate margin, an observation that potentially places the Prydz Bay coastline with a back-arc setting prior to collision (Kelsey et al., 2008). This would imply that subduction polarity was beneath the Indo–Antarctic plate, an inference consistent with the significantly larger volume of Cambrian aged intrusions within the Indo–Antarctic plate when compared to the Australo–Antarctic plate, as well as the mantle derived and subduction related geochemical signatures obtained from some of these intrusions (Liu et al., 2006; Boger et al., 2008). Within the Australo–Antarctic plate, Cambrian orogenesis resulted in the formation of generally NW trending folds within the Ruker and 356 S.D. Boger / Gondwana Research 19 (2011) 335–371 Gondwana break-up (Collins, 2003; Halpin et al., 2008). Elements of this belt may be exposed as deformed basement in the Himalaya where remarkably similar ages are found. As part of greater India prior to Indo–Eurasian collision, these rocks lie along strike of the projection of the Kuunga Orogen in Antarctica (Fig. 12a). In the opposite direction, the Kuunga suture passes beneath the Antarctic icesheet and likely reappears along the Antarctic coast in the Shackleton Range (Kleinschmidt and Boger, 2009). The oldest rocks of the Shackleton Range are of Palaeoproterozoic (2330–1830 Ma) age and form part of the Australo–Antarctic plate (Brommer et al., 1999; Zeh et al., 1999, 2004; Will et al., 2009). These rocks are exposed in the south of the range and include the Read, Pioneers and Stratton Groups (Fig. 13c). In addition to their Palaeoproterozoic history, these rocks also preserve evidence for Cambrian reworking at approximately 510 Ma (Zeh et al., 1999; Will et al., 2009). Deformation at this time is supported by the stratigraphic relationships preserved in the overlying strata. The Watts Needle and Mt. Wegner formations were deposited in the Ediacaran and early Cambrian and deformed and metamorphosed at low grades (Buggisch and Henjes-Kunst, 1999; Sodruzhestvo cover-sequences and more localised shear zones within the basement rocks of the Ruker Complex (Phillips et al., 2005b, 2007). Metamorphism was also at generally lower grades when compared to that observed in the Lambert Complex. Peak pressures and temperatures are estimated at 4.0–5.0 kbar and 565–580 °C. Ar– Ar ages are between 530 Ma and 480 Ma from both the Ruker basement and the overlying cover-sequences (Phillips et al., 2007). These ages are consistent with some earlier Rb–Sr and U–Pb lower intercept ages obtained from the same area (Mikhalsky et al., 2001b). The continuation of the Kuunga Suture to the east of the Prince Charles Mountains is mostly unconstrained, although the intrusion of a 515 Ma syenitic pluton to the immediate west of the Denman Glacier in Princess Elizabeth Land, as well as 600–500 Ma lower intercepts observed from the surrounding Archaean country rocks (Black et al., 1992b) indicate that the Kuunga suture may pass nearby (Fig. 12a). This is supported by similar metamorphic ages (530– 510 Ma) obtained from the Leeuwin Complex of Western Australia and the adjacent offshore Naturaliste Plateau, regions which were both proximal to the Denman region of East Antarctica prior to East Antarctic basement terranes Kalahari Craton (domain 1) NEFB Kuunga Suture rt h Maud-Natal Belt ( 1130-1060 Ma) s tr a lia n C r a t o n + No Kaapvaal and Grunehogna cratons Au ia 4 2 + TAM Crohn Craton EM Kaapvaal craton Pinjarra Belt (1330- 1140 Ma) Precambrian Africa + + + Central Antarctic Craton (domain 4) Precambrian South America + + + + Devonian and Carboniferous intrusions + + AP(e) + + D + + + aA PF + P + + + + Ch + + Ediacaran-Cambrian collisional belts (550-520 Ma) + + + + Cambrian collisional belts (530-490 Ma) CP(e) + + u st NP Ter r CFB SFB Pan-African Orogenic Belts Ediacaran collisional belts (580-550 Ma) CP(w) RBT MB(a) AP(cw) G on G Albany-Fraser Belt (1330- 1140 Ma) + MB(r) o g e n 5 1 0-3 0 0 M a 4 1 Ch ChP NZ(w) + + s Or Nawa-Coompania Block (undifferentiated Palaeo & Mesoproterozic) Crohn Craton (Middle to Late Archaean ± Proterozoic cover sequences) Folded basement with overlying and undeformed Mesoproterozoic volcanic rocks (1110 Ma) + + PM Curnamona-Beardmore Block (undifferentiated Palaeoproterozoic) + Ma Mawson Mawson Craton a to n ar C r a rw dw on Davis + VL r a li Mawson Craton (Late Archaean) G Australo-Antarctic Craton (domain 3) 3 Dh ana Rayner Belt (990-900 Ma) Undifferentiated Archaean and Palaeoproterozoic rocks LH + + AF Casey rog e n 3 0 0 -1 0 0 r In d + id e O Indo-Antarctic Craton (domain 2) + T B dwan G re a te + Cu Cu West Antarctic (and related) terranes Reworked passive margin sediments (Ediacaran to Permian) Terrains and/or Gondwana derived turbiditic sedments accreted 510-300 Ma Terrains accreted 110 Ma Fig. 13. Antarctic palaeogeography from the middle Cambrian to the late Carboniferous. Westward subduction under the Pacific margin of Gondwana marks onset of Terra Australis orogenesis. Black stars mark exposures of the passive margin sequences: AF = Adelaide Fold Belt (Adelaidean and Kanmantoo Groups), CFB=Cape Fold Belt, EM = Ellsworth– Whitmore Mountains, NP = North Patagonia Terrane, PF = Puncoviscan Formation, PM = Pensacola Mountains (Hannah Range Formation), SFB=Sierra de la Ventana Fold Belt, TAM = Transantarctic Mountains (Beardmore Group and lower Byrd Group), VL = Victoria Land (Koettlitz Group). Terranes accreted from the Late Cambrian to the Late Carboniferous include: AP(e) = eastern domain Antarctic Peninsula, B = Bowers Terrane, Ch = Chilenia Terrane, ChP = Challenger Plateau (western domain New Zealand (NZ(w))), CP(w) = Campbell Plateau (western domain New Zealand), Cu = Cuyania Terrane, D = Deseado Terrane, LH = Lord Howe Rise, MB(r) = Ross Province Marie Byrd Land, NEFB = New England Fold Belt, P = Pampean Terrane, RBT = Robertson Bay Terrane and T = Tasmanides. Outboard terranes: AP(cw) = central and western domains of the Antarctic Peninsula, CP (e) = Campbell Plateau (eastern domain New Zealand) and MB(a) = Amundsen Province Marie Byrd Land. Numbered regions (1–5) refer to the tectonic domains shown in Fig. 1. S.D. Boger / Gondwana Research 19 (2011) 335–371 Buggisch and Kleinschmidt, 2007). Deformation in these rocks is bracketed by their depositional ages (N540 Ma) and by the age of the overlying and undeformed red beds of the Ordovician (N490 Ma) Mt. Provender Formation (Buggisch and Kleinschmidt, 2007). Similar radiometric ages are obtained from an ophiolite complex exposed north of the Australo–Antarctic plate margin (Fig. 12c). The ophiolitic rocks are interpreted to represent part of a relic ocean that separated the Australo–Antarctic Craton from the Coats Land Block (Tessensohn et al., 1999). These rocks are of peridotitic composition and preserve garnet-olivine bearing assemblages indicative of subduction to pressures of 20–25 kbar and peak metamorphic temperatures of 700–850 °C (Schmädicke and Will, 2006; Romer et al., 2009). Metamorphism is constrained to between 530 Ma and 490 Ma based on Sm–Nd garnet-whole rock and K–Ar amphibole ages (Talarico et al., 1999; Romer et al., 2009). Identical ages are known from the Northern Gneiss Complex. These rocks extend from the Northern Haskard Highlands to the Herbert Mountains and the Lord and Baines Nunataks on the Pioneers Escarpment (Will et al., 2009). This region is dominated by paragneisses of unconstrained origin and dioritic to granitic intrusions that were emplaced between 530 Ma and 520 Ma (Will et al., 2009). Metamorphism in the host rocks was contemporaneous with granite intrusion (Brommer and Henjes-Kunst, 1999; Brommer et al., 1999; Zeh et al., 1999, 2004). The age of orogenesis observed in the Shackleton Range thus shows a remarkable correspondence to that observed in the Prince Charles Mountains. With mean ages for deformation in both areas spanning the early and middle Cambrian (530–500 Ma) these belts represent the youngest of the “Pan-African” belts observed in Antarctica and indeed within Gondwana generally. Both belts also lie the furthest from the components of West Gondwana. On the basis of these geographic and chronologic similarities these two regions are thought to lie along the same structure (Kleinschmidt and Boger, 2009). Although argued to represent a suture by some (Boger et al., 2001; Boger and Miller, 2004; Kleinschmidt and Boger, 2009), this view is not agreed upon by all (Phillips et al., 2006; Squire et al., 2006). Despite the conjecture about the significance of the Kuunga Orogen, features typical of intercontinental sutures are present. In the Prince Charles Mountains, there is clear evidence for the juxtaposition of terranes with distinct and unrelated geological histories. Although subduction related features are largely absent in the Prince Charles Mountains, they are preserved in the Shackleton Range where, conversely, the opposing margins of the suture are not directly observed. The effects of collision along the Kuunga Orogen were manifested more regionally within both the Afro-Antarctic and Indo–Antarctic plates. In Dronning Maud Land (Afro-Antarctic plate), the compressional structures (580–550 Ma) associated with the collision of the Coats Land and Kalahari blocks are overprinted by locally developed extensional and strike-slip shear zones and by the intrusion of gabbros and A-type granitoids (Jacobs et al., 2003b,c). The timing of this overprint occurred contemporaneously with collision along the Kuunga Suture. The plate scale extrusion of material and the fracturing of the crust into what were to become a series of discrete microplates (Jacobs and Thomas, 2004) as well as the widespread intrusion of post-tectonic granitoids were, in the author's opinion, driven by the nearby collisional tectonics observed along the Kuunga Suture. Similarly, intra-plate stresses associated with Kuunga collision resulted in the formation of shear zones and associated pegmatite emplacement within parts of the Indo–Antarctic Rayner Complex (Carson et al., 2000; Boger et al., 2002; Wilson et al., 2007). 11. Post-Gondwana accretionary growth—the Terra Australis and Gondwanide Orogenies The suturing of the West Gondwana and Australo–Antarctic plates along the Kuunga Orogen (Fig. 12a) brought to an end the long-lived 357 process of convergence between the pre-collision components of Gondwana. The result was a reconfiguration of the early to middle Cambrian plate system and the consequent transfer of ocean floor consumption from between the pre-Gondwana cratons to the outboard Pacific margin of newly formed Gondwana supercontinent (Boger and Miller, 2004; Cawood, 2005; Foden et al., 2006; Cawood and Buchan, 2007). This led to the establishment of the accretionary Terra Austrais Orogen (Cawood, 2005), a general name given to the orogenic belt that stretched continuously from northern South America to northern Australia and which began in the early to middle Cambrian and lasted until the late Carboniferous (Fig. 13). Although Terra Australis orogenesis mostly post-dated the formation of Gondwana, there is nevertheless evidence that subduction began, at least locally, along the Pacific margin of South America and Antarctica before Gondwana was completely formed. The first such location is known from South America where subduction related plutonism dates to between 555–525 Ma. This phase of magmatism slightly pre-dated the Pampean Orogeny and was accompanied by some deformation (Lira et al., 1997; Sims et al., 1998; Rapela et al., 1998b; Stuart-Smith et al., 1999; Schwarz et al., 2008). The second area is found along the Antarctic margin between the central Transantarctic Mountains and Victoria Land where magmatism and some reworking of the East Antarctic margin can similarly be traced back to at least 530 Ma (Black and Sheraton, 1990; Goodge et al., 1993; Rowell et al., 1993; Encarnación and Grunow, 1996; Allibone and Wysoczanski, 2002). More generally however Terra Australis orogenesis began between 520 Ma and 510 Ma—slightly after the terminal suturing of Gondwana. The onset of this event was responsible for the termination of passive margin sedimentation along the majority of the Pacific margin of Gondwana and marks the onset of widespread and broadly coeval deformation and arc-type plutonism. It also began a long-lived process of accretion that added much of the crust that defines eastern Australia, West Antarctica (domain 5) and western South America (Cawood, 2005; Cawood et al., 2009). In Australia, Terra Australis (Delamerian) orogenesis is best described from South Australia where it reworks the late Neoproterozoic and early Cambrian sedimentary rocks of the Adelaidean and Kanmantoo Groups. The effects of this orogeny are nevertheless traceable, albeit more obscurely, along the entire Precambrian margin of Australia. In southern Australia, Delamerian deformation resulted in the formation of west-verging folds and thrusts coincident with variable grades of metamorphism. High geothermal gradients are described from close to the craton margin in the west while lower temperatures at similar pressures are described more distally in the east (Flöttmann et al., 1994; Sandiford et al., 1995; Miller et al., 2005). This metamorphic anisotropy is consistent with west-dipping subduction beneath the Australian margin (Finn et al., 1999). Syntectonic and post-tectonic plutons were derived from both mantle and crustal sources, although they are not clearly subduction related (Foden et al., 2002). The age of the oldest of these intrusions is interpreted to constrain the onset of orogenesis to approximately 515 Ma (Foden et al., 2006), the same time that subduction and obduction of ocean floor took place in Tasmania (Turner et al., 1998). Similar to slightly younger ages are argued for the onset of Delamerian orogenesis in northern Australia (Nishiya et al., 2003; Fergusson et al., 2007). In Antarctica Terra Australis (Ross) orogenesis also deformed and variably metamorphosed the pre-Gondwana passive margin (Fig. 13). The Antarctic passive margin, in order of increasing distance from the Australian margin, is defined by the Wilson Terrane, then the Skelton, Koettlitz, Beardmore and Hannah Ridge Formations (e.g. Rowell et al., 2001; Goodge et al., 2004; Wysoczanski and Allibone, 2004; Tessensohn and Henjes-Kunst, 2005). The highest metamorphic grades are found in the Wilson Terrane. Deformation and metamorphism in this region correlate well with that observed in the Delamerian Orogeny of southern Australia (Flöttmann et al., 1993, 358 S.D. Boger / Gondwana Research 19 (2011) 335–371 1994; Finn et al., 1999). The protolith rocks of the Wilson Terrane are mostly of metasedimentary origin. They record multiple generations of folding, and show a zonation in metamorphic conditions such that a low-pressure high-temperature internal belt and an external highpressure low-temperature belt are present (Ricci et al., 1997; Federico et al., 2006). The earliest phases of magmatism in the Wilson Terrane date to approximately 540–530 Ma, some 20 Ma earlier than is recognised in Australia (Allibone and Wysoczanski, 2002; Giacomini et al., 2007). Nevertheless, the vast majority of plutons were emplaced between 510 Ma and 480 Ma, an interval identical to that observed in southern Australia (Encarnación and Grunow, 1996; Cox et al., 2000; Foden et al., 2006). In Antarctica, these rocks are more clearly of volcanic arc origin (Allibone and Wysoczanski, 2002). In the central Transantarctic Mountains (Fig. 13), the onset of deformation is dated to between 540 Ma and 520 Ma (Goodge et al., 1993) and is similar in timing to the implied onset of plutonism observed in Victoria Land. However, the effects of this early phase of deformation appear relatively limited, and in common with elsewhere along the Australo–Antarctic margin, the main phase of Terra Australis orogenesis is constrained to the early to middle Cambrian (515–505 Ma). In this region the onset of orogenesis is most clearly recorded in the sedimentary record where a distinct carbonate-toclastic transition is interpreted to reflect a shift from passive sedimentation to active uplift and erosion associated with orogenesis (Myrow et al., 2002; Goodge et al., 2004). Similar stratigraphic relationships are described from the Pensacola Mountains (Fig. 13) where an unconformity between the Hannah Ridge Formation and the overlying Patuxent Formation suggests that deformation slightly predated the late middle Cambrian, a finding confirmed by a more recent isotopic data which ties deformation to the interval 510–500 Ma (Rowell et al., 2001; Goodge et al., 2004; Curtis et al., 2010). The extension of the Gondwana passive margin beyond the Pensecola region of Antarctica includes rocks that now define the Cape Basin of South Africa as well as correlative rocks found in both the Ellsworth–Whitmore Mountains of West Antarctica and in the Sierra de la Ventana of southern South America (Fig. 13). These sedimentary successions are defined by thick (8–13 km) volcanosedimentary sequences that were derived mostly from the erosion of nearby African and South American sources and which depositionally spanned the Middle Cambrian to Permian (Webers et al., 1992; Curtis et al., 1999; Flowerdew et al., 2007). The evolution of these regions is dominated by extension and subsistence, punctuated in South Africa and South America by the localised early to middle Cambrian (550– 510 Ma) granitoid emplacement (Rozendal et al., 1999; Scheepers and Armstrong, 2002; Rapela et al., 2003). Within the Sierra de la Ventana magmatism occurred during ongoing extension, while in the Cape Basin mild deformation accompanied plutonism. Plutonism is not known from the Ellsworth–Whitmore Mountains, although early to middle Cambrian volcanism is common and correlates temporally magmatism observed elsewhere (Curtis, 2001; Rapela et al., 2003). This sector of the Gondwana margin is thus distinguished from elsewhere in the respect that ongoing deposition overlapped with convergent orogenesis observed elsewhere. This may suggest that the deposition of the Cape, Ellsworth–Whitmore and the Sierra de la Ventana strata were localised either in a continental back-arc basin or in a setting whereby subduction was not present immediately outboard of this section of the Gondwana margin (Curtis, 2001; Rapela et al., 2003). Beyond this tectonically quiet zone, Terra Australis subduction was active along the remainder of the Precambrian margin of South America (Fig. 13). At the time this was defined by western edge of the Amazonian and Rio de La Plata Cratons where, similar to the majority of the Australo–Antarctic margin, Terra Australis (Pampean) orogenesis reworked the Ediacaran to early Cambrian passive margin. In central South America the passive margin is defined by the turbidite dominated Puncoviscan Formation (Jezek et al., 1985; Rapela et al., 1998a). The latter stages of deposition within this formation were followed closely by the onset of deformation which is constrained to between 525 Ma and 510 Ma based on the widespread intrusion of peraluminous syn-tectonic granites, Ar–Ar data from locally formed ultramylonites, and an angular unconformity that separates the deformed Puncoviscan Formation from the overlying middle to late Cambrian (b510 Ma) platform sediments of the Mesón Group (Rapela et al., 1998a,b; Schwarz et al., 2008; Mulcahy et al., 2010). Deformation and metamorphism occurred at conditions varying from greenschist to granulite facies, with the higher temperature rocks recording P–T–t paths characterised by decompression (Rapela et al., 1998b; PiñánLlamas and Simpson, 2006). Pampean orogenesis arguably terminated with the collision of the Pampean terrane (Fig. 13), a relatively large continental block consisting of the Arequipa–Antofalla and the Western Sierras Pampeanas cratons (Rapela et al., 1998a, 2007). Collision at this time is in contrast to other segments of the Terra Australis Orogeny that remained ocean-facing. Following the onset of Terra Australis orogenesis, the shift in the locus of convergence from within Gondwana to the continent's Pacific margin set in motion a long-lived phase of semi-continuous accretion (Cawood et al., 2009). Along the Australian margin this began in the early Ordovician and included the vast oceanic province defined by the Tasmanides (Fig. 13). These rocks consist of Gondwana derived detritus that were deposited mostly as turbidity flows onto a substrate of ocean floor basalt (Fegusson and Coney, 1992; Spaggiari et al., 2004; Squire et al., 2006). These rocks were semi-continuously deposited along the Australian margin and then stepwise accreted from the early Ordovician to early Carboniferous (490–340 Ma, Foster and Gray, 2000; Foster et al., 2009). This process incorporated a number of arc terranes, the youngest of which crops out in the New England Fold Belt (Cawood and Leitch, 1985). More regionally the Tasmanides correlate with parts of the mostly submerged continental blocks defined now by the Lord Howe Rise and the Challenger and Campbell plateaus (Fig. 13). Small portions of the latter plateaus are exposed in New Zealand where they crop out along the west coast of the south island (Sutherland, 1999; Adams, 2008). These rocks are in turn correlated with rocks exposed in both the Bowers Terrane and Marie Byrd Land (Ross province), regions now found in Antarctica (Adams, 2008; Bradshaw et al., 2009). Terrane accretion in Antarctica began at much the same time as it did in Australia. The first addition to the Gondwana margin was the Bowers Terrane, a sequence of rocks that consist of a mafic substrate of island arc or back-arc origin (Sledgers Group) overlain by a cover of mostly terrigenous sediments known as the Mariner and the Leap Year groups (Weaver et al., 1984; Federico et al., 2006). The two cover-sequences are of middle to late Cambrian age and consist of fossiliferous fine-grained mudstone with lesser limestone and conglomerate. These grade up sequence into coarser-grained quartzose sandstones and conglomerate. Collision with the Gondwana margin occurred along the Lanterman fault, a major tectonic lineament that exposes mafic rocks metamorphosed to eclogite facies (Di Vincenzo et al., 1997). Underthrusting and metamorphism are dated to between 505 Ma and 490 Ma, an interval taken to date collision with the Wilson Terrane margin of Gondwana (Di Vincenzo et al., 1997; Federico et al., 2006). The adjacent and outboard Robertson Bay Terrane (Fig. 13) is similarly defined by a thick sequence of continentally derived late Cambrian to early Ordovician flysch, but differs in the respect that it appears to be underlain by somewhat older continental crust (Borg et al., 1987; Fioretti et al., 2005; Gemelli et al., 2009). These basement rocks are exposed on Surgeon Island where paragneisses of probably Neoproterozoic age are intruded by Cambrian granite (Fioretti et al., 2005). Xenocrystic zircons recovered from the granite commonly have Palaeo- and Mesoproterozoic (1800–1500 Ma) ages. Zircons of this age are also widely recovered from the Neoproterozoic sedimentary and intrusive rocks exposed western Tasmania, a region S.D. Boger / Gondwana Research 19 (2011) 335–371 359 East Antarctic basement terranes Kalahari Craton (domain 1) NEFB Maud-Natal Belt ( 1130-1060 Ma) rt h Kaapvaal and Grunehogna cratons s tr a l ia n C r a t o n No Indo-Antarctic Craton (domain 2) Au Rayner Belt (990-900 Ma) T 132 G re a te r In d ia + 96 + KP ar a rw Dh Somali Basin Albany-Fraser Belt (1330- 1140 Ma) LG D 4 M CP(e) MB(r) 5 Pinjarra Belt (1330- 1140 Ma) DFZ Central Antarctic Craton (domain 4) Crohn Craton (Middle to Late Archaean ± Proterozoic cover sequences) Coats Land Block - folded basement and undeformed Mesoproterozoic volcanic rocks ( 1110 Ma) Pan-African Orogenic Belts Riiser-Larsen Sea + MB(a) + + + + AP(e) Precambrian Africa + + EM 4 1 Ch ChP + NZ(w) CP(w) + Crohn Craton 2 + 83 165 3 132 C ra t o n Mawson Craton (Late Archaean) Nawa-Coompania Block (undifferentiated Palaeo & Mesoproterozic) + + + C Mawson Craton Australo-Antarctic Craton (domain 3) Curnamona-Beardmore Block (undifferentiated Palaeoproterozoic) LH + 83 Undifferentiated Archaean and Palaeoproterozoic rocks 165 AP(cw) + + + Kaapvaal craton Cretaceous suture (120- 110 Ma) Cambrian collisional belts (530-490 Ma) + Ediacaran-Cambrian collisional belts (550-520 Ma) + + + Ediacaran collisional belts (580-550 Ma) + + Precambrian South America + + + + + + West Antarctic (and related) terranes + + + + Devonian and Carboniferous intrusions Permo-Triassic strata + + + + P + + Reworked passive margin sediments (Ediacaran to Permian) Terrains and/or Gondwana derived turbiditic sedments accreted 510-300 Ma Terrains accreted 110 Ma Fig. 14. Rifting of Gondwana. Boxed ages between continents give age of first appearance of seafloor between continents. DFZ = Davey Fracture Zone, KP = Keguelen Plume (black star), LG = Lambert Graben. Labelled domains in western South America, West Antarctic and eastern Australia are the same as those in Fig. 13. Numbered regions (1–5) refer to the tectonic domains shown in Fig. 1. C,D and M = Australian research stations Casey, Davis and Mawson. that crops out along strike from the Robertson Bay terrane (Turner et al., 1998; Black et al., 2004). In turn, these Tasmanian rocks are further considered to lie, largely unexposed, under a sizeable area in central Victoria (Cayley et al., 2002; McLean et al., 2010). If the rocks from these regions are comparable then together they could define a sizable continental slither of either para-autochthonous or exotic origin. The approach of these rocks potentially drove the late Cambrian to early Ordovician accretion of both the Bowers and Robertson Bay terranes (Federico et al., 2006), although deformation in the Robertson Bay Terrane arguably occurred somewhat after that observed along the Lanterman fault, a conclusion that would suggest a more stepwise series of accretionary events that began in the west and migrated east with time (Tessensohn and Henjes-Kunst, 2005). The next significant terrane along the Ordovician margin of Gondwana is defined by the Ross Province of Marie Byrd Land (Fig. 13). Similar to the Robertson Bay Terrane, this region is dominated by a thick sequence of early Ordovician turbidites. Known as the Swanson Formation, these rocks show very similar detrital zircon populations to temporally equivalent sediments found within the Tasmanides (Pankhurst et al., 1998b; Siddoway et al., 2004b; Squire et al., 2006) and, similar to the Tasmanides, were deformed and metamorphosed at low grades during the late Ordovician (440 Ma, Adams et al., 1995). If the model for the development of the Tasmanides during this interval can be applied to Marie Byre Land, then the Swanson Formation was probably deposited as a thick turbidite blanket in a back-arc or marginal oceanic basin, and was thickened and imbricated shortly thereafter as a result of accretion to the continental margin (Foster et al., 2009). If this scenario is correct it implies that the Ross Province of Marie Byrd Land was accreted Gondwana's Pacific margin in the late Ordovician or early Silurian (450–430 Ma), slightly after the Bowers and Robertson Bay terranes. Along the South American sector of the Gondwana margin, Ordovician and younger accretion more commonly involved the collision of often allochthonous continental fragments, as opposed to the off-scraping and imbrication of the Gondwana derived accretionary wedge. Although not all agree, this potentially began with the late Cambrian accretion of the Pampean terrane (Rapela et al., 1998a, 2007). This was followed by the establishment of the early to middle Ordovician (495–460 Ma) Famatanian magmatic arc (Rapela et al., 1998a; Pankhurst et al., 1998a, 2000, 2006; Ramos et al., 1998; Sims et al., 1998; Dahlquist et al., 2008; Chernicoff et al., 2010). Famatanian 360 S.D. Boger / Gondwana Research 19 (2011) 335–371 plutons intrude the Arequipa–Antofalla rocks of northern Chile and Peru and define a semi-continuous belt of intrusions traceable into northern Patagonia (Loewy et al., 2004; Pankhurst et al., 2006). Arc magmatism was manifested along the continental margin of Gondwana, behind which a wide back-arc basin formed and was filled by coeval siliciclastic and volcaniclastic sediments (Astini and Dávila, 2004; Bahlburg et al., 2006). Subduction along the Famatanian arc progressively closed the ocean basin between Gondwana and the Cuyania (Argentine Precordillera) terrane, rocks which are widely viewed as being of Laurentian origin (Fig. 13). This conclusion is based on: (1) the presence of late Cambrian North American trilobites in Cuyanian strata, (2) similarities between these strata, their tectonic subsidence curves and their detrital zircon spectra, and those observed from the Laurentia passive margin and, (3) Cuyania basement rocks with Mesoproterozoic ages that overlap with those observed from Grenville Province (Bond et al., 1984; Kay et al., 1996; Astini, 1998; Thomas et al., 2004; Ramos, 2004, 2010). Collision of the Cuyania Terrane inverted the Famatanian back-arc in the middle Ordovician (470 Ma). Deformation and metamorphism of the same age is also observed on the Cuyania plate (Astini et al., 1995; Ramos et al., 1998; Voldman et al., 2009). The final two terranes observed along the South American margin are the Deseado and Chilenia terranes, both of which are widely considered to have collided in Devonian or Carboniferous times (Ramos et al., 1986; Pankhurst et al., 2006). These two terranes today crop out respectively south of latitude 47° near the southern tip of South America and in western Argentina between 29°S and 33°S. The Chilenia Terrane consists of broadly Grenville age basement rocks (Ramos, 2010) that are tectonically overlain/juxtaposed with a Neoproterozoic to Cambrian age metasedimentary–ophiolitic complex (López de Azarevich et al., 2009). Known as the Guarguaraz Complex, the ophiolitic rocks are generally interpreted to have defined the oceanic floor between the Chilenia and Cuyania plates. The tectonically intercalated sedimentary rocks are implied to have been deposited on the western margin of the Cuyania Terrane (Davis et al., 1999; López de Azarevich et al., 2009). Deformation of these units, which is dated by Ar–Ar geochronology and stratigraphic constraints to between 390 Ma and 350 Ma, is interpreted to mark the collision of the Chilenia Terrane (von Gosen, 1995; Davis et al., 1999). Metamorphism accompanying collision was up to lower-amphibolite facies with garnet bearing metapelitic rocks preserving evidence for burial to pressures of 13 kbar (T = 500 °C). Subsequent tectonic unroofing and rapid decompression is argued to reflect collision driven crustal thickening and subsequent isostatic uplift (Massonne and Calderón, 2008). Inboard of the zone of collision, deformation associated with the accretion of the Chilenia Terrane was manifested via the development of greenschist to amphibolite facies shear zones (Sims et al., 1998). The similarly timed collision of the Deseado Terrane is perhaps more controversial as this terrane contains many features similar to the rocks of the North Patagonia Terrane (Pankhurst et al., 2003), which itself is variably argued to be both autochthonous (Pankhurst et al., 2006) and allochthonous (Ramos, 2008) with respect to the Gondwana margin. Pankhurst et al. (2006) argued that the North Patagonia Terrane probably formed part of the Gondwana margin as the protolith rocks of this region consist of Cambro–Ordovician metasediments of Gondwana origin that are intruded by Middle Ordovician granites analogous to those of the Famatinian arc. Working from this base, the accretion of the Deseado Terrane was then argued to have occurred in the middle Carboniferous after the establishment of Devonian (400–370 Ma) and then Carboniferous (330–315 Ma) phases of magmatism (Pankhurst et al., 2006). There is however a significant overlap in plutonic ages from both the Deseado and North Patagonia terranes from about 400 Ma onwards (Pankhurst et al., 2003) and this may imply an early Devonian rather than middle Carboniferous timing of collision. Temporal equivalents of the Devonian and Carboniferous intrusions observed in both the Deseado and North Patagonia terranes are also widespread in the South American and Australo–Antarctic plates. Intrusions of this age are found within the Pampean Terrane (Ramos et al., 1998) but more widely include: (1) the Target Hill granitoids (390–320 Ma) of the Antarctic peninsula (Eastern domain, Vaughan and Storey, 2000; Millar et al., 2002), (2) the Ford granodoirite suite and subsequent intrusions (380–330 Ma) of Marie Byrd Land (Ross Province, Pankhurst et al., 1998b; Muksaka and Dalziel, 2000; Siddoway and Fanning, 2009; Korhonen et al., 2010), (3) the Admiralty suite (370– 350 Ma) of the Robertson Bay Terrane (Stump, 1995), (4) the Karamea– Paringa (370–360 Ma) and Ridge–Tobin (355–340 Ma) suites of western and southern New Zealand (Tulloch et al., 2009), (5) the 400–350 Ma granites of the central Victorian magmatic province (Arne et al., 1998; Beirlein et al., 2001; Black et al., 2005) and finally (6) the magmatic rocks of the Retreat Batholith (380–370 Ma) and the silicic magmatism associated with the Campwyn Volcanics (360–350 Ma) from the New England Fold Belt of northeastern Australia (Bryan et al., 2004). These intrusions by and large are interpreted to have been emplaced in the Devonian to Carboniferous continental margin of Gondwana, a conclusion that implies their host terranes were accreted to the margin of Gondwana by 400 Ma. The late Carboniferous is taken to mark the end of Terra Australis orogenesis and the onset of the Gondwanide Orogeny (Cawood, 2005). Gondwanide orogenesis coincides the formation of the supercontinent Pangaea that was formed via the collision of Gondwana and the combined Laurentia and pre-Permian Eurasia. Gondwana–Laurentia collision occurred along the northern and western margins of South America and Africa and was thus distal from the Pacific margin of Gondwana. Nevertheless, collision at this time more than likely led to a significant kinematic reorganisation of plate motions (Cawood and Buchan, 2007). This was manifested along the Gondwana margin by the inversion of the previously passive section of the margin defined by the Cape, Ellsworth–Whitmore and the Sierra de la Ventana basins (Curtis, 2001). The folding of these rocks coincided with widespread plutonism within Patagonia and in both the central domain of the Antarctic Peninsula and the Amundsen province of Marie Byrd Land (Vaughan and Storey, 2000; Pankhurst et al., 1998b, 2006). These latter two terranes are interpreted to have remained separate from the remainder of Antarctica until the Cretaceous (Di Venere et al., 1995; Pankhurst et al., 1998b; Vaughan et al., 2002), an inference that suggests Permo–Triassic subduction was not in all places directly beneath the Gondwana margin. Along the Australian sector of the margin, Gondwanide orogenesis resulted in widespread deformation within the New England (265–230 Ma) Fold Belt (Cawood et al., 2009). 12. Rifting of Gondwana and the formation of modern Antarctica The final phase of Antarctica's tectonic evolution was one of crustal extension that ultimately led to the disintegration of Gondwana. This occurred in three main phases beginning in the Jurassic with the separation of Africa and South America from the Dronning Maud Land margin of Antarctica. Rifting then progressed in a clockwise fashion such that India and then Australia separated in the early Cretaceous before New Zealand separated from the West Antarctica in the late Cretaceous (Reeves and de Wit, 2000). The break-up of Gondwana was however preceded by the widespread deposition of Devonian to Triassic strata. Within Antarctica, these rocks include the foreland deposits of the Beacon Supergroup (Collinson et al., 1994) as well as the intracontinental deposits of the Amery Group (McKelvey and Stephenson, 1990) and the temporally equivalent strata exposed in Dronning Maud Land (Bauer, 2009). The Beacon Supergroup is exposed widely within the Transantarctic Mountains (Fig. 14) and is partly derived from, and deposited inboard of, the Gondwanide continental arc (Elliot and Fanning, S.D. Boger / Gondwana Research 19 (2011) 335–371 2008). The Amery and Dronning Maud Land deposits were deposited within Gondwana following intra-continental extension and sagging that began at the end of the Carboniferous (Veevers, 1988). These rocks were sourced mostly from an inferred Antarctic highland centred on the Gamburtsev Subglacial Mountains, the detritus from which was deposited principally within the African Karoo and the Indian Gondwanide basins, as well as forming a widespread cover over much of Australia (Veevers and Saeed, 2007, 2008). The Dronning Maud Land strata define cold climate fluvimarine sediments that represent the relics of the Karoo basin within Antarctica (Bauer, 2009). The Amery Group defines the upslope equivalents of the Gondwana strata in India and is mostly defined by fluvial sandstone and coal deposits (Fielding and Webb, 1995; McLoughlin and Drinnan, 1997a,b; Holgate et al., 2005). Sediment deposition was quite localised as increased denudation along the margins in both basins slightly preceded or was contemporaneous with sediment deposition (Lisker et al., 2003; Emmel et al., 2004, 2007). This implies structural control of both basins, a feature attributed to the accommodation of distal within-plate stresses associated with ongoing tectonism along the Pacific margin of Gondwana (Lisker et al., 2003; Schandelmeier et al., 2004; Phillips and Läufer, 2009). Approximately 100 Ma after the onset of sedimentation, the intraplate Karoo–Ferrar magmatic province developed in the hinterland of Gondwana's Pacific margin. Consisting mostly of mafic intrusive and extrusive rocks, the Karoo–Ferrar Province follows closely the foreland basin in which the Beacon Supergroup was deposited. Karoo–Ferrar rocks can be traced from southern Africa, through Antarctica, and into southern Australia. They were emplaced over a short interval in the middle Jurassic (185–172 Ma) with melting and magmatism variably argued to be either related to an extensive mantle plume or linked to subduction along the Gondwana margin and the formation of the Beacon basin (Cox, 1978; Hergt et al., 1991; Storey, 1995; Storey and Kyle, 1997; Encarnación et al., 1996; Duncan et al., 1997; Zhang et al., 2003; Jourdan et al., 2005). It remains debated as to what extent Karoo–Ferrar magmatism was related to the onset of Gondwana break-up. Some studies link these events closely (Vaughan and Storey, 2007), while others argue that, although closely spaced in time, these events were nevertheless independent (Jokat et al., 2003). The eventual separation of Africa from Antarctica was achieved in the middle Jurassic and led initially to the formation of ocean floor in the Riiser–Larsen Sea located between Dronning Maud Land and southeastern Africa and in the Somali Basin between Madagascar and Kenya (Fig. 14). These two areas were linked by a transform known as the Davey Fracture Zone which accommodated Africa's near orthogonal motion relative to Antarctica, and its strike-slip motion relative to India and Madagascar (Reeves and de Wit, 2000). The timing of continental separation is dated via the oldest seafloor anomalies in the Somali Basin and via a break-up unconformity in the Karoo equivalents exposed in Madagascar to approximately 165 Ma (Rabinowitz et al., 1983; Geiger et al., 2004). An equivalent age is obtained from the seafloor anomaly record observed off the coast of Dronning Maud Land (Roeser et al., 1996; Jokat et al., 2003). The initial break-up of Gondwana thus involved only two plates with Africa and South America moving northwards and away from Antarctica–India–Madagascar–Australia (Fig. 14). This two-plate system was not long-lived and by the early Cretaceous India and Australia were both in motion relative to Antarctica. This was driven by the counter-clockwise rotation of India and the opposite clockwise rotation of Australia away from Antarctica (Powell et al., 1988). The result was the formation of ocean floor between India and the western margin of Australia as well as between India and Antarctica by approximately 132 Ma (Fullerton et al., 1989; Rotstein et al., 2001; Robb et al., 2005; Gaina et al., 2007). Within Antarctica, break-up related deformation is described from the Lambert basin (Fig. 14), a 361 modern drainage system that has its origins in the late Carboniferous (Lisker et al., 2003). Within this basin the Permo–Triassic strata of the Amery Group are faulted internally and down thrown against the Proterozoic basement as a result of transtensional intra-plate stresses (Boger and Wilson, 2003; Phillips and Läufer, 2009). The separation of India and Antarctica has widely been ascribed to the interaction of the Kerguelen plume with the Gondwana lithosphere (Kent, 1991; Storey, 1995). However, the most recent data show that the formation of ocean floor between India and Antarctica (~ 130 Ma) preceded the main phase of Kerguelen magmatism (~120 Ma). It is thus not clear if the insipient Kerguelen plume was a driver of rifting, or if rifting and the attendant thinning of the crust was a preferential locus for the plume nucleation (Gaina et al., 2007). The Kerguelen plume did nevertheless have a significant impact on the newly formed ocean floor between India and Antarctica and on the nearby continental crust on either side of the ancestral Indian Ocean. Magma output from the Kerguelen plume commenced between 132 Ma and 123 Ma with the emplacement of the low volume Casuarina-type and Gosselin-type basalts observed in southwestern Australia (Coffin et al., 2002). The peak of magma production was somewhat later. On the ocean floor, Kerguelen magmatism between 119 and 112 Ma extruded 8.5 × 106 km3 of basaltic rock that now defines the southern Kerguelen Plateau (Coffin et al., 2002; Duncan, 2002). Indian onshore equivalents of these rocks include the basaltic rocks that define Rajmahal Traps as well as the mafic and felsic extrusives that underlie the Began basin. These rocks collectively cover an area estimated at 2 × 105 km2 and are dated to approximately 118 Ma (Baksi, 1995; Kent et al., 2002). Basalts of similar age also intrude the Permian to Cretaceous strata within the Mahandi graben (~117 Ma) while low volumes of identically aged (~118–119 Ma) mafic rocks also intrude the upslope equivalents of these strata in Antarctica (Amery Group, Arne, 1994; Lisker and Fachmann, 2001). Having begun synchronously with rifting between India and Antarctica, the rifting between Australia and Antarctica differs in the respect that relative motion effectively stalled just after initiation and extension remained continental for much of the early Cretaceous (Powell et al., 1988; Veevers et al., 1991). Relative motion nevertheless continued extremely slowly and by 96 Ma the first seafloor developed between the Australian and Antarctic plates (Veevers et al., 1991). The rate of drifting remained subdued into the middle Eocene (~ 46 Ma), after which time the rates of divergence increased dramatically. Thereafter Australia moved rapidly north relative to Antarctica. Slow crustal extension between Australia and Antarctica overlapped temporally with the final accretionary events recorded along the Gondwana margin (Fig. 14). Collision sutured the central and western provinces of the Antarctic Peninsula to eastern Province by 107 Ma, which coincided broadly with the collision of the Amundsen and Ross provinces of Marie Byrd Land (Di Venere et al., 1995; Muksaka and Dalziel, 2000; Vaughan et al., 2002). The timing of collision between the components of Marie Byrd Land is further correlated with collision inferred between the western and eastern domains of New Zealand. The western domain defines the Gondwana margin and consists of Gondwana derived Ordovician sediments that are widely intruded by late Devonian to early Carboniferous magmatic arc related intrusions, rocks which correlate directly with the Ross Province of Marie Byrd Land (Ireland and Gibson, 1998; Tulloch et al., 2009). The opposing eastern domain is defined by the Median Tectonic zone, a mostly magmatic region dominated by Jurassic to early Cretaceous plutonic rocks with ages between 250–220 Ma and 170–130 Ma, and the eastern province of New Zealand defined predominantly by the Permian to early Cretaceous turbiditic sediments (Landis and Coombs, 1967; Kimbrough et al., 1994; Muir et al., 1998; Mortimer, 2004). These latter rocks, in addition to being exposed in New Zealand, define much 362 S.D. Boger / Gondwana Research 19 (2011) 335–371 of the eastern Campbell Plateau (Sutherland, 1999). It remains debated whether the magmatic arc represented by the Median Tectonic zone formed along or proximal to the eastern edge of Gondwana or was somewhat removed from this margin (Muir et al., 1995, 1998; Waight et al., 1998; Mortimer et al., 1999). Assembly of these terranes was nevertheless achieved by the early Cretaceous, as these regions are stitched by plutons of the 124–105 Ma Separation Point suite (Waight et al., 1998). In the West Antarctic–New Zealand sector of Gondwana collision was followed by a rapid switch to crustal extension. In Marie Byrd Land, extension was accompanied by widespread A-type granite and mafic dyke intrusion, together with diapiric doming and rapid unroofing of mid-crustal rocks (Muksaka and Dalziel, 2000; Siddoway et al., 2004b, 2005). This occurred between approximately 105 Ma and 95 Ma and resulted in approximately 350 km of crustal extension between Marie Byrd Land and East Antarctica and led to the formation of the thinned continental crust that today underlies the Ross Sea (Storey et al., 1999; Muksaka and Dalziel, 2000; Siddoway et al., 2004a,b). Contemporaneous and along strike extension was also manifested between the eastern coast of Australia and the Lord Howe Rise (Schellart et al., 2006). Similar to the Ross Sea, extension in this region produced distended continental crust, but by 90 Ma, no ocean floor. Continental separation between the Lord Howe Rise and the eastern margin of Australia was achieved at around 83 Ma (Gaina et al., 1998). An identical age for the first appearance of ocean crust is described from between West Antarctica and Campbell Plateau (Larter et al., 2002). Rifting was thus achieved via a two-stage process—extension and crustal thinning between Australia–East Antarctica and New Zealand–West Antarctica between approximately 105 Ma and 95 Ma, followed by extension and then separation along a slightly oblique axis between West Antarctica–Australia and New Zealand (+related blocks) by 83 Ma (Kula et al., 2007). The separation of New Zealand from the margin of West Antarctica effectively formed the modern Antarctic continent as we recognise it today. Since rifting, the Antarctic plate has remained effectively stationary and geological activity has been limited to ongoing extension and volcanism within the West Antarctic rift system. The majority of deformation occurred between 43 Ma and 26 Ma during which time another 180 km of extension occurred (Cande et al., 2000). Deformation at this time is commonly linked to the uplift and 4–9 km of denudation along the rift flank now defined by the Transantarctic Mountains (e.g. Fitzgerald and Gleadow, 1988; Fitzgerald and Stump;, 1997; Miller et al., 2010). Neogene to Quaternary alkaline volcanism defined by the McMurdo Volcanic Group also occurred along the margins of the Ross Sea. 13. Some concluding remarks The modern Antarctic continent together with its inferred geological make-up is illustrated in Fig. 14. Given most of Antarctica remains glaciated, the geology shown for most of the continent has been inferred from the limited inland or coastal outcrops and by extrapolating across previously connected rifted margins. Although, clearly this process leaves significant room for error, I nevertheless believe the gross distribution of rocks shown in this figure likely reflects a first order approximation of what underlies the Antarctic ice. One of the first and more important conclusions is that the core of East Antarctica (domains 3 and 4) is probably dominated by Archaean and Palaeoproterozoic crust. For domain 3, this is consistent with the known rock exposures along the Terre Adélie coastline and from the Transantarctic Mountains and Shackleton Range. It is further supported by granitoid boulders of Mesoproterozoic age that are found in moraines in various localities (Peucat et al., 2002; Goodge et al., 2008). These boulders, once of intrusive origin, imply older country rocks. The case for an Archaean or Palaeoproterozoic age of the rocks of domain 4 is given by the antiquity of the exposures in the Obrechev Hills and southern Prince Charles Mountains (e.g. Black et al., 1992b; Boger et al., 2006; Phillips et al., 2006). The implied late Mesoproterozoic collision between the Crohn and Mawson cratons, further implies that both sides of this suture were cratonised by at least the middle Mesoproterozoic. The rocks underlying the remainder of East Antarctica (domains 1 and 2) are more or less exposed along the coast between longitudes 30°W and 100°E. Although commonly overprinted to varying degrees by Ediacaran to Early Cambrian events related to their amalgamation with the rest of East Antarctica, these rocks nevertheless have wellestablished Mesoproterozoic and early Neoproterozoic protolith ages. Interestingly this age pattern is not well replicated in detrital zircon datasets described from Permo–Triassic and younger sediments shed from the interior of Antarctica (Veevers and Saeed, 2007, 2008; van de Flierdt et al., 2008; Veevers et al., 2008). Although commonly containing some older zircons, these sediments are by and large dominated by Ediacaran and Cambrian age zircons with a secondary population of late Mesoproterozoic to Tonian grains. Given that the sampled sediments are mostly from the coast, it is possible that they are tapping relatively local sources, which between Dronning Maud and Wilkes Lands (longitudes 30°W and 100°E) are dominated by rocks with ages similar to those obtained. However, more distal inland sources, such as the Gambertsev Subglacial Mountains are commonly ascribed to these rocks. If these inland sources are correct the predominance of relatively young grains (1200–500 Ma) appears contradictory to the conclusions drawn here—that most of the interior of Antarctica is defined by Mesoproterozoic or older rocks. This contradiction is perhaps explained if one considers that: (1) detrital zircons are not necessarily derived directly from the rocks in which they originally formed, but can undergo a number of phases of erosion and redeposition and, (2) the reconstruction of the crust underlying the Antarctic ice undertaken here cannot account for unexposed sedimentary rocks that may overlie these basement rocks. It is thus possible that the dominance of Edicaran–Cambrian and late Mesoproterozoic–Tonian zircons in the Permo–Triassic and younger coastal strata may reflect the reworking of for example Cambro–Ordovician sedimentary rocks that: (1) define large but unknown sequences that overlie the basement illustrated in Fig. 14 and, (2) are dominated by detrital zircons with late Mesoproterozoic to Cambrian ages. With this in mind, it is interesting to note that much of the vast turbidite fan deposited along Gondwana's Pacific margin is dominated by detrital zircons with similar age distributions to those obtained from the Antarctic Permo–Triassic strata. These sediments reflect a vast outpouring of detritus related to the formation of Gondwana. These sediments originated in part from the Gondwana collision between Africa and Antarctica (Squire et al., 2006) and were probably recycled through the Kuunga orogenic cycle enroute to their final site of deposition along Gondwana'a eastern margin. The extent to which these rocks were deposited onshore in Antarctica remains unknown. Although it remains the author's opinion that the majority of the metamorphic basement rocks of central East Antarctica are at least of Mesoproterozoic age, I would not discount the potential significance of post-Gondwana strata which may overlie these rocks. Acknowledgements The present manuscript was written in response to an invitation from M Santosh to write an overview of Antarctic tectonics for Gondwana Research. I sincerely thank him for firstly suggesting this project, without such a nudge I may not otherwise have undertaken such a review. I must also thank him for his patience as an editor—it has been a long time between the inception of this study and its final publication. Chris Wilson and Chris Carson introduced me to Antarctic geology and both have remained supporters of my endeavours there. To both I am very grateful. I am similarly grateful to my various S.D. 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Journal of Petrology 45, 949–973. Zhang, X., Luttinen, A.V., Elliot, D.H., Larsson, K., Foland, K.A., 2003. Early stages of Gondwana breakup: The Ar40/Ar39 geochronology of Jurassic basaltic rocks from western Dronning Maud Land, Antarctica, and implications for the timing of magmatic and hydrothermal events. Journal of Geophysical Research 108. doi:10.1029/2001JB001070. Zhao, Y., Song, B., Wang, Y., Ren, L., Chen, T., 1992. Geochronology of the late granite in the Larsemann Hills, East Antarctica. In: Yoshida, Y., Kaminuma, K., Shiraishi, K. (Eds.), Recent Progress in Antarctic Earth Science. Terra Scientific Publishing, Tokyo, pp. 155–161. Zhao, Y., Liu, X., Song, B., Zhang, Z., Li, J., Yao, Y., Wang, L., 1995. Constraints on the stratigraphic age of metasedimentary rocks from the Larsemann Hills, East Antarctica: possible implications for Neoproterozoic tectonics. Precambrian Research 75, 175–188. Zhao, J.-X., Ellis, D.J., Kilpatrick, J.A., McCulloch, M.T., 1997a. Geochemical and Sr–Nd isotopic study of charnockites and related rocks in the northern Prince Charles Mountains, East Antarctica: implications for charnockite petrogenesis and Proterozoic crustal evolution. Precambrian Research 81, 37–66. Zhao, Y., Xiaohan, L., Shicheng, W., Song, B., 1997b. Syn- and post-tectonic cooling and exhumation in the Larsemann Hills, East Antarctica. Episodes 20, 122–127. Zhao, Y., Liu, X.H., Liu, X.C., Song, B., 2003. Pan-African events in Prydz Bay, East Antarctica and its inference on East Gondwana tectonics. In: Yoshida, M., Windley, B., Dasgupta, S. (Eds.), Proterozoic East Gondwana: Supercontinent Assembly and Breakup: Geological Society of London, Special Publication, London, 206, pp. 231–245. Steven Boger graduated with a BSc (Hons) from Monash University (1995). A short stint in industry preceded his PhD, which he undertook at the University of Melbourne (2001). This was followed by post-doctoral fellowships at both Monash and Melbourne universities as well as stints as the Prince of Asturias Fellow at the University of Bremen and as a visiting scientist at the universities of Bergen and Mainz. Boger is presently an Honorary Fellow at the University of Melbourne. He additionally works as a freelance geologist mostly on development aid projects funded by the likes of the World Bank. He is a broad based field geologist with particular interests in applying structural, petrologic, and geochronological techniques to solving tectonic problems. He has a particular interest in the evolution of the components of Gondwana and the events that lead to the eventual amalgamation. He has ongoing research interests in Antarctica, Madagascar and Morocco.

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