Terrane Processes at the Margins of Gondwana (Geological Society Special Publication No. 246)

Terrane Processes at the Margins of Gondwana

Geological Society Special Publications Society Book Editors R. J. PANKHURST (CHIEF EDITOR) P. DOYLE F. J. GREGORY J. S. GRIFFITHS A. J. HARTLEY R. E. HOLDSWORTH

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It is recommended that reference to all or part of this book should be made in one of the following ways: VAUGHAN, A. P. M, LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246. TESSENSOHN, F. & HENJES-KUNST, F. 2005. Northern Victoria Land terranes, Antarctica: fartravelled or local products? In: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 275-291.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 246

Terrane Processes at the Margins of Gondwana

EDITED BY

A. P. M. VAUGHAN, P. T. LEAT British Antarctic Survey, UK

and

R. J. PANKHURST British Geological Survey, UK

2005 Published by The Geological Society London

THE GEOLOGICAL SOCIETY

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Contents Preface 1. VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. Terrane processes at the margins of Gondwana: introduction

vii 1

Regional syntheses 2. GLEN, R. A. The Tasmanides of eastern Australia

23

3. HIBBARD, J. P., MILLER, B. V., TRACY, R. J. & CARTER, B. T. The Appalachian peri-Gondwanan realm: a palaeogeographical perspective from the south

97

4. ADAMS, C. J., PANKHURST, R. J., MAAS R. & MILLAR, I. L. Nd and Sr isotopic signatures of metasedimentary rocks around the South Pacific margin and implications for their provenance

113

5. VAUGHAN, A. P. M. & LIVERMORE, R. A. Episodicity of Mesozoic terrane accretion along the Pacific margin of Gondwana: implications for superplume-plate interactions

143

6. WANDRES, A. M. & BRADSHAW, J. D. New Zealand tectonostratigraphy and implications from conglomeratic rocks for the configuration of the SW Pacific margin of Gondwana

179

7. RAPELA, C. W., PANKHURST, R. J., FANNING, C. M. & HERVE, F. Pacific subduction coeval with the Karoo mantle plume: the Early Jurassic Subcordilleran belt of northwestern Patagonia

217

8. MILLER, H. & SOLLNER, F. The Famatina complex (NW Argentina): back-docking of an island arc or terrane accretion? Early Palaeozoic geodynamics at the western Gondwana margin

241

9. LUCASSEN, F. & FRANZ, G. The early Palaeozoic Orogen in the Central Andes: a non-collisional orogen comparable to the Cenozoic high plateau?

257

10. TESSENSOHN, F. & HENJES-KUNST, F. Northern Victoria Land terranes, Antarctica: far-travelled or local products?

275

Topics and methodologies 11. READING, A. M. Investigating the deep structure of terranes and terrane boundaries: insights from earthquake seismic data

293

12. RAPALINI, A. E. The accretionary history of southern South America from the latest Proterozoic to the Late Palaeozoic: some palaeomagnetic constraints

305

13. CORDANI, U. G., CARDONA, A., JIMENEZ, D. M., Liu, D. & NUTMAN, A. P. Geochronology of Proterozoic basement inliers in the Colombian Andes: tectonic history of remnants of a fragmented Grenville belt

329

14. STONE, P. & THOMSON, M. R. A. Archaeocyathan limestone blocks of likely Antarctic origin in Gondwanan tillite from the Falkland Islands

347

15. LEAT, P.T., DEAN, A.A., MILLAR, I. L., KELLEY, S. P., VAUGHAN, A. P. M. & RILEY, T.R. Lithospheric mantle domains beneath Antarctica

359

16. ZIMMERMANN, U. Provenance studies of very low- to low-grade metasedimentary rocks of the Puncoviscana complex, northwest Argentina

381

17. SIDDOWAY, C. S., SASS, L. C. Ill & ESSER, R. P. Kinematic history of western Marie Byrd Land, West Antarctica: direct evidence from Cretaceous mafic dykes

417

Index

439

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Preface The concept of tectonostratigraphic terranes emerged from the study of Northern Hemisphere active continental margins and erogenic belts; however, it is largely in the Southern Hemisphere, in particular in relation to the Gondwana supercontinent, where it has seen its fullest flowering. The entire Pacific margin of the supercontinent Gondwana is regarded as a collage of accreted terranes and the resulting orogen is one of the world's largest. This book brings together a series of reviews and multidisciplinary research papers that comprehensively cover the terranes from the Tasman Orogen of eastern Australia to the Neoproterozoic and Palaeozoic orogens of South America, taking in New Zealand and Antarctica along the way. The book is a collection of original papers and reviews based on work presented at the Terrane Processes at the Pacific Margin of Gondwana' (TAPMOG) meeting held in Cambridge, UK in September 2003 and at the 10th Chilean Geological Congress held in Conception in October 2003, as well as invited contributions. This volume is a contribution to the British Antarctic Survey SPARC project of the programme Antarctica in the Dynamic

Global Plate System and a posthumous contribution to International Geological Correlation Project 436 'Pacific Gondwana Margin'. The editors thank the sponsors of TAPMOG and the IGCP 436 session at the 10th Chilean Geological Congress for financial support towards the scientific meetings and this special publication. Thanks, in particular, go to the referees for their important contribution: F. G. Acenolaza, N. W. Archbold, C. Augustsson, D. Avigad, K. N. Bassett, T. S. Brewer, P. A. Cawood, U. G. Cordani, A. J. Crawford, I.W. D. Dalziel, V. J. DiVenere, R. W. England, T. H. Fleming, P. G. Flood, G. Franz, G. Gibson, S. Hada, R. Hall, F. Herve, T. R. Ireland, J. L. Isbell, E. C. King, R. L. Larson, R. Maas, W. J. McCourt, C. Mac Niocaill, J. G. Meert, M. A. Menzies, N. Mortimer, J.B. Murphy, S.R. Noble, D. G. Pearson, C. A. Ricci, P. Schmidt, B. C. Storey, F. Tessensohn, R. J. Thomas, K. Ueno, C. R. van Staal, H. von Eynatten, S. J. Whitmeyer and two anonymous reviewers. Alan Vaughan, Phil Leat & Bob Pankhurst Cambridge and Keyworth December 2004

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Terrane processes at the margins of Gondwana: introduction ALAN P. M. VAUGHAN1, PHILIP T. LEAT1 & ROBERT J. PANKHURST2 British Antarctic Survey, High Cross, Madingley Road, Cambridge CB3 OET, UK (e-mail: a. [email protected]) ^British Geological Survey, Keyworth, Nottingham NG12 5GG, UK

l

Abstract: The process of terrane accretion is vital to the understanding of the formation of continental crust. Accretionary orogens affect over half of the globe and have a distinctively different evolution to Wilson-type orogens. It is increasingly evident that accretionary orogenesis has played a significant role in the formation of the continents. The Pacific-margin of Gondwana preserves a major erogenic belt, termed here the 'Australides', which was an active site of terrane accretion from Neoproterozoic to Late Mesozoic times, and comparable in scale to the Rockies from Mexico to Alaska, or the Variscan-Appalachian orogeny. The New Zealand sector of this erogenic belt was one of the birthplaces of terrane theory and the Australide orogeny overall continues to be an important testing ground for terrane studies. This volume summarizes the history and principles of terrane theory and presents 16 new works that review and synthesize the current state of knowledge for the Gondwana margin, from Australia through New Zealand and Antarctica to South America, examining the evolution of the whole Gondwana margin through time.

Why this book? Two main types of orogenic belt have been identified on the Earth: orogens of the Wilson type (e.g. Wilson 1966; Murphy & Nance 2003), involving collision between continents, and orogens of the accretionary, Cordilleran type (e.g. Sengor & Natalin 1996; Tagami & Hasebe 1999; Scarrow et al 2002), where a more steady-state addition of smaller crustal fragments occurs. In simplest terms, collisional orogens are assumed to be the endpoint of cycles of ocean formation and destruction during continental break-up and re-amalgamation (the so-called 'Wilson Cycle'); accretionary orogens are the product of more continuous processes of addition of oceanic, island-arc and ocean-captured continental material to oceanic margins during long-term subduction, often without oceanic closure. Accretionary orogens are often characterized by being much wider across-strike than collisional orogens (Sengor & Okurogullari 1991). Overall, it would seem that collisional and accretionary orogens form end-members of a spectrum (Murphy & Keppie 2003; Murphy & Nance 2003). The Wilson-type end-member is the 'aulacogen' (e.g. Zolnai 1986; PedrosaSoares et al 2001), where there is little or no displacement of continental margins and the ocean basin, which is often narrow, mostly closes up with jigsaw precision. The simplest accretionary end-member consists of a complex or prism (e.g. Leggett 1987; Doubleday et al 1994; Kamp 2000), created through scraping-off

of the upper parts of oceanic lithosphere as it is subducted. This may form at a continental margin or adjacent to an intra-oceanic arc and, ultimately, may be displaced large distances, either across ocean basins or along continental margins, during, or subsequent to, formation. Real situations are a complex mix of Wilsonand accretionary types (e.g. Betts et al 2002), where full-scale oceans that form with or without subsequent closure may have accretionary complexes on their margins, and experience subsidiary terrane and arccollisional orogens that themselves incorporate accretionary complexes. Even 'pure' accretionary orogens, such as the Uralide-Altaid orogen that forms much of Asia (Sengor & Natalin 1996), where there is no evidence of continent-continent collision, consist of many minor terrane and arc-collisional orogens that occurred on the margins of a long-lived ocean. The understanding of the relative significance of Wilson-type and accretionary orogens has changed with time. Historically, much early work focused on Wilson-type orogenesis, particularly in the context of the circumAtlantic orogens affecting NW Europe and eastern North America from Proterozoic through to Late Palaeozoic times (e.g. Phillips et al 1976; Williams & Hatcher 1982; Keppie 19850; Ryan & Dewey 1997; Matte 2001; Young et al 2001; Bandres et al 2002; Gower & Krogh 2002), which may have inflated its significance in global terms. Williams & Hatcher (1982)

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246,1-21. 0305-8719/$15.00 © The Geological Society of London 2005.

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were the first to show that a Wilson-type model may not be appropriate for the case examples associated with the evolution of the lapetus Ocean and this is supported by more recent lithostratigraphic, faunal and palaeomagnetic data (O'Brien etal 1983; Cocks & Torsvik 2002; Hartz & Torsvik 2002). Doubts about its applicability to the Pacific Cordillera of western North America came even earlier (Banner 1970). Although major, Wilson-type continental collision can form long-lived continents such as Gondwana (e.g. Unrug 1992; Boger et al. 2001), global syntheses (e.g. Sengor & Natalin 1996) have emphasized the importance of accretionary orogens, arguing that these were responsible for growth and stabilization of millions of square kilometres of the continental lithosphere from Archaean times onwards (Sengor et al. 1993; Foster & Gray 2000; Polat & Kerrich 2001; Xiao et al 2004). Accretionary orogens are directly and indirectly host to globally important mineral deposits (e.g. Richards & Kerrich 1993; Sherlock et al. 1999; Kerrich et al 2000; Goldfarb et al 2001). For example, volcanic arc and back-arc terranes form an important part of accretionary orogens and it is recognized increasingly that active, submerged arcs and back-arcs are sites of significant metallogenesis (e.g. Fouquet et al 1991; Ishibashi & Urabe 1995; lizasa et al 1999; Fiske etal 2001). Submarine arc-hosted mineral deposits are not easily accessible in their sites of formation because seafloor mining is challenging technically (e.g. Scott 2001); however, terrane accretionary orogenesis has an additional importance in incorporating arc and seafloor mineral deposits in continental lithosphere, making them accessible to simpler extraction techniques. Finally, in this volume, Vaughan & Livermore (2005) present evidence that accretionary orogenesis is not uniformly distributed in time, with implications for our understanding of Earth evolution. Vaughan & Livermore (2005) show that two major pulses of terrane accretion occurred in the Mesozoic, not just affecting the 'Australides' on the Gondwana margin but with global extent, possibly associated with major episodes of flood magmatism. The accretionary orogenic belt that formed on the palaeo-Pacific and Pacific margin of Gondwana in Neoproterozoic-Mesozoic times (Fig. 1) (Ireland et al 1998) is one of the largest known orogenic belts in Earth history. The orogen (the Neoproterozoic to Palaeozoic part of this orogeny has been called the Terra Australis orogen by Cawood & Leitch (2002)) now occupies the eastern third of Australia,

New Zealand, West Antarctica, the Transantarctic Mountains and large parts of southern South America (Fig. 1) (e.g. Bradshaw 1994; Cawood & Leitch 2002). Several factors have hampered reconstruction of this Neoproterozoic to Mesozoic orogenic belt - a timeextended Terra Australis orogen, finishing in the mid-Cretaceous, which will be referred to here informally as the 'Australides'; Figure 1. These include dismemberment and dispersal of its components during Mesozoic break-up of the supercontinent Gondwana, burial of large parts of it beneath ice (in Antarctica) and later sedimentary basins (in all other parts of the belt), local submergence of continental margins (notably the continental margins of New Zealand) and partial obliteration and overprinting by continued subduction-related magmatism and deformation (as in many parts of the Andes and Antarctic Peninsula). The size of the reconstructed orogenic belt - over 7000 km long by over 1500 km wide - is larger than the Mesozoic-Cenozoic orogenic belt that extends from Alaska to Mexico along the Pacific margin of the North American continent and is comparable in scale to the Variscan orogenic belt of Europe and eastern North America (Matte 2001). The better-understood western North American orogenic belt or 'Cordillera' (Fig. 1) was the first to be interpreted as a collage of 'suspect' terranes (as summarized by Coney et al 1980) - terranes being fault-bounded blocks of the Earth's crust characterized by a geological history distinct from that of adjacent terranes. This model has been applied widely and successfully to many ancient orogenic belts. One important aspect of the model is that some of the terranes in a collage may have travelled great distances from their places of origin to their final location adjacent to other terranes or the continental margin. This may have occurred either across oceans or along a continental margin by transcurrent faulting (e.g. Keppie & Dallmeyer 1987; Mankinen et al 1996; Cowan et al 1997; Takemura et al 2002). The terrane model was applied early in its inception to parts of the Pacific margin of Gondwana (Coombs et al 1976; Bradshaw et al 1981; Weaver et al 1984; Murray et al 1987). Now it is accepted almost universally that terrane-style tectonics are of major importance in the development of orogenic belts. Understanding of the 'Australide' orogen (Fig. 1) and the role of terrane processes in its development has progressed rapidly during the last decade. There are several reasons for this.

TERRANE PROCESSES

3

Fig. 1. Time-extended Terra Australis' orogen (cf. Cawood & Leitch 2002) or 'Australides', including Permo-Triassic orogenesis of the Gondwanian and Hunter-Bowen events (e.g. Collins 1991; O'Sullivan et al. 1996; Curtis 2001) and Triassic-Jurassic and mid-Cretaceous deformation events (Vaughan 1995; Vaughan & Livermore 2005) depicted on 200 Ma Pangaea reconstruction of Vaughan & Livermore (2005).

1.

2.

There has been increasing realization that problems of tectonic correlation within the orogen are solved best by comparisons between the now-dispersed parts of the belt. Increasing knowledge of formerly less well known parts of the orogen and greater international co-operation in sharing such knowledge (notably through UNESCOfunded International Geological Correlation Programmes, such as IGCP 436 'Pacific Gondwana Margin') have been important. The formerly adjacent parts of the orogen may have been either sediment sources or terrane sources (e.g. Adams et al 1998; Cawood et al 2002). The advent of routine, accurate and precise U-Pb dating of zircons has led to more refined correlation of events and provided information on the provenance of the huge piles of quartz-bearing sediments that characterize much of the orogen (e.g.

3.

4.

Ireland et al 1998; Fergusson & Fanning 2002; Herve et al 2003; Schwartz & Gromet 2004; Wandres et al 2004). Advances in the routine application of 40Ar/39Ar dating have also been beneficial to provenance studies as well as the dating of deformation events (Adams & Kelley 1998; Vaughan et al 2002). Improved geochemical analytical methods for analysing trace and rare earth elements in volcanic rocks (especially ICP-MS) and greater understanding of their compositional variations has led to growing confidence in assigning tectonic settings to the volcanic arcs that are key elements in the orogenic belt for palaeotectonic reconstructions (e.g. Glen et al 1998; Spandler et al 2004; Wang et al 2004). Increased use of remote sensing techniques and improved regional compilations of data, especially magnetic potential field,

4

5.

A. P. M. VAUGHAN ET AL.

has been highly effective in mapping and characterizing terrane boundaries, especially when submerged or covered by surficial deposits or ice (e.g. Ferraccioli & Bozzo 1999; Sutherland 1999; Direen & Crawford 2003). Palaeontological discoveries and better biostratigraphical correlation have been important in the recognition that terrane activity continued into the Mesozoic and have provided qualitative estimates of terrane transport directions and distances (Benedetto 1998; Fang et al 1998; Kelly et al 2001; Cawood et al 2002).

This book is the first to provide an overview of understanding of the terrane model as it applies to the Australide accretionary orogeny on the Pacific margin of Gondwana. It reviews the work of research groups from North and South America, Europe and Oceania who are engaged in active research on the nature of the Gondwana margin and the accretionary orogeny (regions covered by chapters of this book are indicated on Fig. 2). This volume offers a snapshot of current thinking and current research directions and is a guide for any researcher currently active or about to embark on studies in this dynamic area of investigation. Now is judged the correct time to summarize recent progress and highlight the scientific questions that are currently engaging those active in the field and that will drive future research. Nomenclature According to Coney et al. (1980), terranes 'are characterized by internal homogeneity and continuity of stratigraphy, tectonic style and history'. They stated that 'boundaries between terranes are fundamental discontinuities in stratigraphy that cannot be explained easily by conventional facies changes or unconformity'. The fundamental features of terranes are, therefore, that (a) their boundaries are major faults, and (b) they have different geological histories to adjacent terranes. These features are summarized in the editors' preferred definition of terranes as 'a fault-bounded package of rocks of regional extent characterized by a geologic history that differs from that of neighbouring terranes' (Howell et al 1985; Friend etal 1988). Recognition of terranes is not based on any inferences about distance travelled or relative movement between adjacent terranes (Parfenov et al. 2000). Terranes are 'suspect' if there is doubt about their palaeographical setting with

respect to adjacent terranes or continental margins (Coney et al. 1980; Coombs 1997). Terranes may be described as 'exotic', 'far-travelled' or 'allochthonous' (all meaning about the same thing) if there is sufficient evidence that they originated far from their present locations, often assumed to be hundreds or thousands of kilometres away; however, these distances need not be particularly large in areas of complex geology (Coombs 1997). Problems of definition have been discussed in the literature (e.g. Sengor & Dewey 1991), mainly from perspectives of recursion (i.e. is a seamount in an accretionary complex a separate terrane or just part of the complex?) and problems of lateral and/or vertical extent (i.e. how small, or large, can a terrane be?; Sengor 1990). Development of the terrane concept Recognition that fragments of continental margins had moved long distances came, in the late 1940s and early 1950s, from the discovery that transcurrent faults had hundreds of kilometres of offset (Kennedy 1946; Hill & Dibblee 1953; Wellman 1955). Coombs (1997), in a brief review of the terrane concept, pointed out that the term 'terrane' was in use as early as the 1920s and 1930s, but that modern usage stemmed from the work of Irwin (1964; 1972) in the western Cordillera of the USA. Following further conceptual development in the 1970s in western North America and New Zealand (Berg et al. 1972; Monger et al. 1972; Blake et al. 1974; Coney 1978; Howell 1980), the concept of terranes, or terrane collages, as possibly fartravelled, fault-bounded blocks with geological histories different from that of adjacent blocks, was crystallized by Coney et al. (1980). The model was quickly tested in other orogens (e.g. Bradshaw et al. 1981; Williams & Hatcher 1982; Ziegler 1982; Pigram & Davies 1987) and large numbers of 'suspect' terranes were identified in most. In the case of the lower Palaeozoic Caledonian-Appalachian orogen in Scandinavia, the British Isles and eastern USA and Canada, terranes were sandwiched between continents on opposing sides of the closing lapetus Ocean (Williams & Hatcher 1982; Hutton 1987; Pickering et al. 1988; Rankin et al. 1988; Hibbard 2000; Roberts 2003). This orogen, therefore, had a phase of accretionary tectonics prior to continent-continent collision. As outlined above, studies of the margin of the Pacific basin were instrumental in the creation of the terrane concept and have provided the impetus for its continued development. In the past ten years, the fundamental importance of

Fig. 2. Geographical areas of terrane studies covered by contributions to this volume (Vaughan et al 2005).

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terrane processes in generating and stabilizing continental lithosphere has become apparent from studies of the Altaid belts of Asia (Sengor et al 1993; Sengor & Natalin 1996) and the orogens that comprise eastern Australia (Foster & Gray 2000) - a system comprehensively reviewed by Glen (2005) in this volume. Application of these ideas to older rocks indicates that terrane amalgamation and accretionary orogenesis may be the most important processes in formation of the continental lithosphere through time (Polat & Kerrich 2001).

Terrane processes The key processes of terrane formation are accretion and dispersal. Accretion (or 'docking' (Twiss & Moores 1992)) is the process by which material incorporated in, or transported by, oceanic plates is added to subducting margins, usually separated from the adjacent, overriding oceanic or continental plate by a narrow zone called a suture (Howell 1989). Sutures may be marked by belts of ophiolitic (e.g. Johnson et al 2003) or high pressure rocks, such as blueschists (e.g. Kapp et al 2003), but survival of these rocks is not essential to the definition and sutures may also be represented by strike-slip faults, thrusts or zones of melange (e.g. Abdelsalam et al 2003; Pavlis et al 2004). Suture zones are not exclusive to accretionary orogens (e.g. Vaughan & Johnston 1992). Dispersal is the process by which fragments are detached or redistributed from the overriding plate at active margins during subduction or ridge crest-trench collision (e.g. Nelson et al 1994; Keppie et al 2003) by rifting (e.g. Umhoefer & Dorsey 1997), strike-slip faulting (e.g. Cawood et al 2002) or thrusting (e.g. Fritz 1996). Both accretion and dispersal result in new terranes, either by adding previously separate geological entities such as oceanic plateaux or seamounts to oceanic margins, or by removing pieces of existing margins and transporting them elsewhere. The third main process is amalgamation (e.g. Bluck 1990), by which existing terranes are combined into larger, composite terrane collages or superterranes, ultimately forming stable parts of the continental lithosphere. Examination of Cenozoic to Recent active accretionary orogens in SE Asia suggests that the interaction between accretion, dispersal and amalgamation can be extremely complex, with geologically very rapid changes that may not be recognized in older orogens without high resolution dating (Hall 2002). Other important terrane processes happen after accretion, dispersal or amalgamation. These are the formation of sedimentary or

volcanic overlap sequences and emplacement of igneous complexes that 'stitch' terrane sutures and place time limits on terrane motion (e.g. Gardner et al 1988; Raeside & Barr 1990; Herzig & Sharp 1992).

Types of terrane The common terrane rock types tend to be similar from orogen to orogen and have been grouped into several main associations. The most common types in Phanerozoic orogens around the world are marginal to the ocean basins and can encompass any type of continental or oceanic lithosphere, either with or without a mantle root. There are five most common, non-genetic (i.e. what a geologist would see at outcrop in the field) where possible, terrane rock-type associations, based on the Australides (this book (Vaughan et al 2005) and references therein), western American Cordillera, Caledonian-Appalachian, central and eastern Asia orogens (Coney et al 1980; Williams & Hatcher 1982; Hutton 1987; Parfenov et al 2000; Badarch et al 2002; Xiao et al 2004). 1.

Turbidite terranes. These are volumetrically very significant, forming large parts of the accretionary orogens in New Zealand (e.g. Leverenz & Ballance 2001; Mortimer 2004), Australia (e.g. Foster & Gray 2000), eastern Asia (e.g. Sengor & Okurogullari 1991), the western North American Cordillera (e.g. Rubin & Saleeby 1991; McClelland et al 1992) and in Palaeozoic orogens of NW Europe and eastern North America (e.g. Keppie 19856; Leggett 1987; Lehmann et al 1995; Ryan & Smith 1998). They comprise thick piles of deepmarine sediments, probably representing submarine fans, and are often imbricated by thrusting (e.g. Kusky & Bradley 1999). They are commonly siliciclastic, particularly in the Southern Hemisphere (e.g. Adams et al 1998; Ireland et al 1998), but substantial calcareous complexes also exist (e.g. Robertson & Ustaomer 2004; Wilson et al 2004). Parfenov et al (2000) subdivided these terranes into three types: two accretionary complex types, with greater or lesser proportions of basaltic rocks; and a non-accretionary type where the evolutionary history of the turbidite succession is less certain (e.g. possibly dispersed from a passive continental margin but with no subsequent incorporation in a subduction complex). Turbidite terranes are commonly metamorphosed

TERRANE PROCESSES

2.

3.

(e.g. Herve & Fanning 2001), from anchimetamorphic up to blueschist and amphibolite grade, and associated brittle-ductile and ductile deformation is common (e.g. Wang & Lu 1997; Willner et al 2004). Tectonic and sedimentary melange tenanes. These are associated commonly with turbidite terranes, particularly those generated in a subduction environment (e.g. Ernst 1993; Kusky & Bradley 1999), and often occur along terrane sutures (e.g. Aitchison et al 2002) or at major boundaries within accretionary complex terranes (e.g. Silberling et al 1988). They consist commonly of altered basalt and serpentinite, chert, limestone, greywacke, shale and metamorphic rock fragments (including blueschist) in a fine-grained sheared and cleaved mudstone matrix (e.g. Aalto 1981; Cloos 1983; Carayon et al 1984; Maekawa et al 2004). Magmatic terranes. These can be predominantly mafic or predominantly felsic, reflecting the geological environment in which they formed. Mafic magmatic terranes are dominated by volcanic and plutonic rocks, usually pillow basalts associated with volcanogenic and pelagic sediments (e.g. Takemura et al. 2002), subaerial flood basalts (e.g. Richards et al. 1991), sheeted dyke complexes (e.g. Lapierre et al 2003) and mid-lower crustal lithologies dominated by mafic and ultramafic plutonic complexes (e.g. DeBari & Sleep 1991; Shervais et al 2004). In some cases, related ultramafic rocks in mafic magmatic terranes may be of mantle origin (Fitzherbert et al 2004). Most mafic magmatic terranes are interpreted to have been generated by either seafloor spreading, oceanic intraplate magmatism or in volcanic arc environments, although terranes derived from dispersal of continental flood basalt magmatic rocks are also known (e.g. Song et al 2004). The products of seafloor spreading include ophiolites and other fragments of oceanic basement that were commonly produced in back-arc settings (e.g. Bluck et al 1980; Cawood & Suhr 1992; Bedard 1999; Yumul 2003; Piercey et al 2004). Terranes derived from seafloor volcanic eruptions include oceanic plateaux formed by oceanic flood eruptions (e.g. Wrangellia terrane - Richards et al 1991; Hikurangi Plateau - Mortimer & Parkinson 1996), as well as seamounts and ocean islands (e.g. Jacobi & Wasowski

4.

7

1985; Barker et al 1988; Doubleday et al 1994). Mafic magmatic terranes are commonly intra-oceanic and formed on oceanic rather than continental crust (e.g. Weaver et al 1984; DeBari & Sleep 1991; Rubin & Saleeby 1991; Miller & Christensen 1994). Dominantly felsic magmatic terranes consist mostly of broadly calcalkaline, plutonic rocks that represent the interiors of volcanic arcs, although some may represent dispersed fragments of older felsic continental crust, possibly cratonderived, reincorporated in later orogens (e.g. Boger et al 2001). In addition to rocks of the calc-alkaline suite, a common association in felsic magmatic terranes is the tonalite-trondhjemite-granodiorite suite (e.g. Smithies 2000). The Phanerozoic equivalents of these rocks are called adakites and their origin is controversial (e.g. Defant et al 2002; Kay & Kay 2002). They are thought to be generated by high heat flow in several subduction-related settings where partial melting took place in the garnet stability zone, > c. 40 km depth, with end-member models implicating either young subducting slab or partial melting of mafic lower arc crust (e.g. Defant et al 2002; Kay & Kay 2002). Terranes with rocks of this suite are typified by the 'Median batholith' and central magmatic arcs of South Island, New Zealand (Muir et al 1998; Mortimer et al 1999). Similar plutonic rocks are interpreted, from seismic evidence, to characterize some modern oceanic arcs (e.g. Suyehiro et al 1996). Some felsic magmatic terranes, or at least some sequences within them, are dominated by felsic volcanic rocks at exposure level (e.g. Clift & Ryan 1994; MacDonald et al 1996; Bryan et al 2001), which are interpreted to be the erupted equivalents of arc and back-arc plutons, although dispersed terranes derived from continental rhyolite large igneous province magmatism are known (e.g. Heatherington & Muller 2003). Non-turbidite clastic, carbonate or evaporite sedimentary terranes. These terranes fall into two categories: well-bedded, shallowmarine, fluvial or terrestrial sequences, probably representing platform, rift margin or shallow basin deposition, and a category consisting of massive limestones. Wellbedded terrane sequences often represent dispersed fragments of continental margin rocks, including clastic and volcaniclastic sediments (e.g. Campbell et al 2001; Noda

8

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A. P. M. VAUGHAN ET AL.

et al 2004), carbonates (e.g. Gaetani et al 2004) and evaporites (e.g. Thomas et al 2001), but can also be deposits from arcrelated basins (e.g. forearc Nichols & Cantrill 2002) or aulacogens (failed rifts) (e.g. Zolnai 1986). Massive limestones are often masses scraped off seamount summits in accretionary complexes (e.g. Kimura et al 1994; Stevens et al 1997; Cawood et al 2002). Terrane collages. These consist of composite terranes formed by amalgamation of some or all of the above terrane types (e.g. the Argentine Precordillera (Thomas et al 2002) and Avalonia (Nance et al 2004)), with the added complication of internal sutures as well as internal overlap and stitch assemblages.

Size of terranes This is a difficult subject. As with many natural objects, it is easier to know what something is than it is to define it. Parfenov et al (2000) placed a lower size limit on terranes by defining them as units that can be mapped at the 1:5 000 000 scale, although they admitted that size limits are largely arbitrary. Sengor (1990) argued that nappes and blocks in melange units should not be considered terranes, but more recent work would suggest that there is no effective lower size limit (melange zones, for example, can be argued to consist of an arbitrarily large number of faults (e.g. Chang et al 2001) - any exotic block in a melange zone is, therefore, fault bounded and could be considered a terrane, although this is an extreme case). Composite terranes can be very large (e.g. modern New Zealand could be considered a composite terrane) and, although composite terranes should be smaller than continents, there is no arbitrary upper limit to terrane size. Current research themes The study of tectonic plates in terms of terranes is called 'terrane analysis' (e.g. Howell & Howell 1995). Once a terrane has been recognized, by identification of its bounding faults, the next component of terrane analysis is characterization (e.g. Samson et al 1990; Lapierre et al 1992), which uses standard geological techniques such as mapping, geophysics (e.g. Brown 1991; Ferraccioli et al 2002; Armadillo et al 2004), sample collection and follow-up laboratory work etc. It is necessary then to determine the relationship between

the terrane and the adjacent continental margin and other neighbouring terranes (e.g. Samson et al 1991); this is often in combination with efforts to characterize a terrane. In most cases, especially in Palaeozoic orogens, it is easy to designate a terrane as 'suspect', but difficult to prove that it is exotic to the continental margin and its immediately adjacent marginal seas. Several techniques exist for testing the origin of a terrane (Howell & Howell 1995). Where a terrane is suspect, techniques to determine qualitative or semi-quantitative estimates of absolute movement include palaeontology (e.g. Smith et al 2001; Belasky et al 2002; Cawood et al 2002; Kottachchi et al 2002), palaeomagnetism (e.g. Johnston 2001; Keppie & Dostal 2001) and palaeoenvironmental studies (e.g. Condie & Chomiak 1996; Monger 1997; Trop et al 2002). Some techniques do not give movement information directly, but can determine what relationship a terrane has to adjacent terranes and/or determine its ultimate origin. These include petrology (e.g. Barr 1990; Restrepopace 1992), isotope geochemistry (e.g. Samson et al 1990; Leat et al 2005), geochronology (e.g. Herrmann et al 1994; Weber & Kohler 1999), provenance studies of sandstones (e.g. Ireland et al 1998; Friedl et al 2000; Adams et al 2002; Adams et al. 2005) and conglomerates (Wandres & Bradshaw 2005) and sediment geochemistry (e.g. Willan 2003). Of the provenance techniques, U-Pb and Hf isotope dating and fingerprinting of zircon are particularly important because, in addition to age, they provide information about evolution of the crustal sources (Bodet & Scharer 2000; Friedl et al 2000; Knudsen et al 2001; Griffin et al 2004). A second approach to characterizing terranes and identifying their origins comes from comparison with modern analogues. For example, the Japan-Izu Bonin arc collision (e.g. Kawate & Arima 1998; Soh et al 1998) is a modern active example of accretion of a primitive magmatic arc to a composite microcontinental arc terrane. Taiwan preserves an active arc-continent collision zone between the Eurasian plate and the Philippine Sea plate (e.g. Fuh et al 1997; Chang et al 2001). The situation in SE Asia is complex and shows evidence for very rapid changes in plate boundaries on geological time-scales, commonly coeval compressional and extensional regimes and abundant strike-slip (Hall 2002). Pigram & Davies (1987) identified as many as 48 Cenozoic terranes in Papua New Guinea/Irian Jaya and the region shows a long history of terrane processes (Metcalfe 1994). The arc-continent collision between Australia and Indonesia/

TERRANE PROCESSES Papua New Guinea (e.g. Abbott et al. 1994), which is complicated by Pacific plate interactions (e.g. Hall 2002), shows features of terrane dispersal even as terrane accretion is underway (e.g. Milsom et al. 1999) and includes an active collision zone between a submarine plateau (Ontong-Java) and the Melanesian arc (e.g. Hall 2002). Many of the features that developed during the Cenozoic development of the region of SE Asia and the southwest Pacific are at odds with interpretations of older accretionary orogens (R. Hall pers. comm. 2004) and the reasons for this mismatch are unexplained so far. The Lesser Antilles, by showing a system where sediments derived from a primitive arc mix with sediments that are derived cratonically, provide a modern analogue that illustrates the potential complexities of provenance analysis based on sediments (e.g. Marsaglia & Ingersoll 1992; Faugeres et al. 1993; Leitch et al. 2003). Aerogeophysical techniques provide powerful tools for identifying terrane extents and boundaries in areas of ice (e.g. Ferraccioli et al 2002) or thick sediment cover (e.g. Chernicoff & Zappettini 2003). Another approach in characterizing and sourcing terranes is to determine the composition and gross structure of the terrane deep-lithosphere. This can be done by examining the compositions of deeply sourced magmas such as primitive mafic dykes and the compositions of any xenoliths they may have carried from depth (e.g. Yu et al. 2003; Leat et al. 2005), or by quantifying the structure of the lithosphere using energy from distant seismic sources such as earthquakes (e.g. Reading et al. 2003; Reading 2005) or magnetotellurics (e.g. Ledo et al. 2004).

Terrane studies on the margin of Gondwana New Zealand New Zealand was one of the places where the terrane concept was developed (Blake et al. 1974; Coombs et al. 1976; Howell 1980; Coombs 1997) and one of the first parts of the Gondwana margin where the terrane concept of Coney et al. (1980) was applied (Bradshaw et al. 1981). The terrane model has proved highly successful in understanding the pre-Late Cretaceous evolution of the region. It has been well tested (Bradshaw 1989; Adams & Kelley 1998; Cawood et al. 1999; Sivell & McCulloch 2000; Mortimer & Cooper 2004) and there have been no competing models for the last twenty years. Early Palaeozoic terranes form a Western Province of Gondwana affinity, which is

9

separated from late Palaeozoic to Mesozoic accreted terranes of an Eastern Province by a Median Tectonic Zone or batholith (Mortimer et al. 1999) that consists of late Palaeozoic to Mesozoic igneous rocks. The first-accreted Eastern Province terranes include ultramafic rocks (such as the type dunites of Dun Mountain: Coombs et al. 1976). Wandres & Bradshaw (2005) review New Zealand's terranes and present new data on their origins using provenance of clasts in conglomerate deposits, arguing that the Antarctic Sector of the Gondwana margin is a major source of detritus. Similarly, Adams et al. (2005) use Sr and Nd isotopes of metasedimentary sequences in the Australides' from Australia to southern South America to characterize Gondwana margin accretionary complexes and the nature of their sources. A simple conclusion of this work is that at any one time material of different origins was being deposited and accreted in different parts of the orogen and that the accretion history of West Antarctica and southern South America is distinct from that of New Zealand.

Australia The Tasman orogenic system Tasmanides' of Australia occupies the eastern third of the Australian continent. It consists of several orogenic belts whose age of deformation and accretion decreases from west to east (Murray et al. 1987; Coney et al 1990; Flottmann et al 1993; Glen et al. 1998; Ireland et al 1998; Fergusson 2003; McElhinny et al 2003). The Early Palaeozoic Delamerian orogeny formed as Neoproterozoic and Cambrian sedimentary and volcanic arc terranes were accreted along the formerly passive margin of the western Australian Precambrian continental core. This orogen is an along-strike correlative of the Ross orogeny in Antarctica (Stump et al 1986; Flottmann et al 1993). To the east, the Early Palaeozoic Lachlan and Thomson fold belts represent accretion of Cambrian to Silurian volcanic arcs and dominantly siliciclastic sediments to the margin (e.g. McElhinny et al 2003). The eastern New England orogeny represents accretion of terranes during late Palaeozoic to early Mesozoic times. In this volume, Glen (2005) comprehensively reviews current models for the development of the Tasman orogenic system and identifies three supercycles of sedimentation and deformation. His proposed model is one of essentially continuous accretionary orogeny at the Pacific margin of eastern Australia since Neoproterozoic times. The deep

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structure of terranes is also important. Reading (2005) presents a new technique for imaging the deep roots of southwestern and southeastern Australian terranes and terrane boundaries using earthquake seismic data.

South America The first ideas that the South American margin might consist of accreted terranes was presented in relation to a Permian carbonate fragment in southern Chile known as the Madre de Dios terrane (Mpodozis & Forsythe 1983). Subsequent work on most of the metasedimentary rocks of the southernmost Pacific Andean margin has shown that they do not represent Palaeozoic Gondwana basement as once supposed, but are best interpreted as Mesozoic accreted material (Herve 1988; Fang et al 1998; Herve & Fanning 2003). East of the Andes, the Argentine Precordillera is regarded widely as a large-scale exotic terrane derived from Laurentia and accreted during the Early Palaeozoic (e.g. Ramos et al 1986; Moore 1994; Astini et al 1995; Thomas & Astini 2003). Much research has been focused on refining models for the history of this Precordillera or Cuyania terrane. Nevertheless, others have disputed its Laurentian derivation, preferring an autochthonous origin within Gondwana (Acenolaza et al 2002). It is notable that geochronology and detrital zircon analysis have been central to the development of both sides of this controversy (Casquet et al 2001; Thomas et al 2004; Finney et al 2005). In any event, it is becoming increasingly obvious that western South America retains a fragmentary record of high-grade Proterozoic metamorphic rocks coeval with the Grenville belt of North America (e.g. Thomas et al 2004). Cordani et al. (2005) present new evidence for Proterozoic, Grenvillian fragments in the Columbian Andes, and Casquet et al (2005) have identified Grenville-age massif anorthosites in western Argentina, comparable to those of the Grenville province. This 'southern Grenville belt' may well represent a common orogeny linking Laurentia and 'Western Gondwana' within Rodinia. Another aspect of importance in terrane accretion is the tectonic history of the collision zone. Miller & Sollner (2005) present evidence that the Famatina Complex represents autochthonous arc-continent collision on the Gondwana margin in Late Proterozoic to Ordovician times, which could be related to accretion of the Precordillera terrane. Zimmermann (2005) uses new provenance data from Late Proterozoic to Cambrian sediments of the

Puncoviscana basin to show that the rocks represent a peripheral foreland basin succession to the Pampean orogeny. As a counterpoint to terrane interpretations of southern South America, Lucassen & Franz (2005) present an alternative history for the early Palaeozoic of the Central Andes, proposing a non-terrane, tectonic situation similar to the present day. Attempts to view the rest of South America in terrane terms were advanced by Bernasconi (1987) for the Precambrian and Ramos (1988) for the Phanerozoic. The current stage is one where the identification and characterization of terranes in poorly exposed or poorly studied areas, which potentially include Patagonia, has not been demonstrated convincingly. Geophysical evidence is crucial in such circumstances (e.g. Chernicoff & Zappettini 2003). Rapalini (2005) reviews southern South American terranes from east to west and provides new insights into key events during Gondwana assembly in the Neoproterozoic to Late Palaeozoic from the perspective of palaeomagnetic data. Rapela et al. (2005) identify a previously unknown Early Jurassic magmatic arc and show that magmatism in the Triassic-Jurassic interval reveals a rotational tectonic regime which should be a major constraint on the plate configuration of Patagonia and the relationship between southern South America and the Antarctic Peninsula in pre-break-up Gondwana reconstructions.

Antarctica The ice cover of most of both East and West Antarctica has hampered regional correlations. Nevertheless, it is becoming increasingly clear that most of East Antarctica consists of a collage of Archaean blocks and Proterozoic belts that were finally stabilized in their current configuration during the Pan-African orogeny (c. 700-500 Ma) when the Mozambique Ocean separating East and West Gondwana closed to form the Gondwana continent (Fitzsimons 20000, b\ Boger et al 2002; Jacobs et al 2003). Closure of this ocean around the end of Precambrian times formed the continuous Pacific margin of the Gondwana continent, along which the Palaeozoic to Mesozoic orogenic belts developed (Fig. 1). During the Mesozoic break-up of Gondwana, West Antarctica behaved as several crustal blocks separated by rift and strike-slip deformation zones (Dalziel & Elliot 1982). Stone & Thomson (2005) present fossil evidence that supports rotation of the Falklands microplate

TERRANE PROCESSES during Gondwana break-up and has implications for the extent of the Gondwanide ice sheet. Siddoway et al. (2005) show new evidence for strike-slip movements affecting West Antarctica during the Cretaceous just prior to the rifting-off from Gondwana of New Zealand (Laird & Bradshaw 2004). West Antarctica appears to consist mostly of crust accreted to the Antarctic margin during Cambrian to Cretaceous times. Terranes were first identified in the Ross orogeny of the Transantarctic Mountains where Cambrian sedimentary and volcanic arc terranes were accreted to the margin in Cambrian times (Weaver et al 1984; Stump 1995). The extent to which these terranes are exotic to the Gondwana margin is a matter of current debate and Tessensohn & Henjes-Kunst (2005) present a review of the most recent results and models. West Antarctica appears to consist of Early Palaeozoic to Mesozoic provinces, at least some of which are 'suspect' (e.g. Pankhurst etal 1998; Vaughan & Storey 2000; Millar et al 2002). Rocks of Proterozoic age (1176 ± 76 Ma, Millar & Pankhurst 1987) crop out in just one location in West Antarctica, at Haag Nunataks. There is continuing uncertainty as to whether this is an isolated far-travelled terrane derived from a continental margin, a fragment of the Gondwana core or whether it represents more extensive Proterozoic basement to West Antarctica, as indicated by isotope studies of granites and xenoliths (Millar et al 2001; Handler et al 2003). Leaf et al (2005) shed some light on this by using lithosphere-derived mafic magmas to determine differences in lithospheric mantle composition beneath Antarctica.

Other parts of the Gondwana margin Studies of the interaction between Gondwana and Laurentia have been an important driver of orogenic theory. When Wilson (1966) asked if the Atlantic had closed and then reopened, the closure he referred to was between Gondwana and Laurentia. Williams & Hatcher (1982; 1983) demonstrated that this closure had incorporated many exotic terranes in one of the first demonstrations of the utility of the terrane collage model of Coney et al (1980). Hibbard et al (2005) present a re-examination of the Gondwana-Laurentia terrane-collision orogeny in the Carolina Zone of the Appalachian belt and present a new model for middle Palaeozoic interactions of the Appalachian periGondwanan realm with Laurentia.

11

Some suggestions for the future Geological mapping, augmented by geochronology, geochemistry, palaeontology and aerogeophysical methods, continues to be the foundation stone of terrane analysis. New techniques, such as ID seismic analysis, Hf-isotope investigation of zircon and xenolith studies, promise to provide further new insights into terrane deep structure and provenance. Geophysical studies, integrated with geological field data, are allowing better prediction of what lies beneath the ice of Antarctica. Despite the many recent advances, there are still some significant gaps in knowledge. One-dimensional seismic studies would benefit from increased Antarctic coverage of permanent seismic data recorders, which is low relative to other continents. Hf-isotope studies are currently hampered by incomplete sets of representative Hf data from potential source rocks for Gondwana terranes. In a sense, there is a need to know the 'Hf of the world' to provide more confidence in the interpretation of Gondwana provenance data and, to this end, it is recommended that all zircon mounts that have been dated by the U-Pb SHRIMP method are analysed in situ for Hf. In regional terms, linking the Australia-New Zealand sector of the Australides with the South American sector is made more difficult by a gap in geological and high-resolution aerogeophysical data in the Pine Island Bay area of Ellsworth Land. This area is currently a target of glaciological research, but it needs to be made a key target for geology and aerogeophysics. It is clear that the long period of Phanerozoic subduction beneath this margin had a large impact on mantle evolution of the Southern Hemisphere. However, there is a need for more robust regional models - building on excellent local datasets - for the origin and relationship of the diverse mantle reservoirs that have sourced magmatism in the Australides. A deeper understanding of terrane processes is likely to result from closer comparisons between the Australides and the Cenozoic accretionary orogens of SE Asia and the southwest Pacific, by re-assessing data and interpretations of older orogens in the context of the wellconstrained processes and events described from Cenozoic margins. In simple terms, one should look for analogues of Mesozoic, and older, processes in younger, better-constrained Cenozoic orogens.

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Conclusions Accretionary orogenesis is a key process of stabilization and formation of the continental lithosphere. Terrane theory and terrane analysis represent the framework for understanding the processes of crustal accretion. The 'Australides' are one of the largest and longest-lived orogens on Earth and have been a key testing ground for the origination and development of terrane theory. Terrane studies continue to be a vibrant and active area of research in the 'Australides', with new techniques and insights emerging on a regular basis. Many research groups from North and South America, Europe and Oceania are active in the region and their work has provided deep insights into the Proterozoic and Phanerozoic evolution of the orogen and the fundamentals of accretionary orogenesis. The 'Australides' are a key area for terrane research and this contribution has attempted to capture the current state of ideas and provide an introduction and benchmark for future research. The papers in this volume stemmed from work presented at the Terrane Processes at the Pacific Margin of Gondwana' meeting held in Cambridge, UK in September 2003 and at the 10th Chilean Geological Congress held in Conception in October 2003, as well as invited contributions. This volume is a contribution to the British Antarctic Survey SPARC project of the programme Antarctica in the Dynamic Global Plate System and a posthumous contribution to International Geological Correlation Project 436 'Pacific Gondwana margin'. The Editors thank Robert Hall and Brendan Murphy for thoughtful reviews that improved the manuscript substantially.

References AALTO, K.R. 1981. Multistage melange formation in the Franciscan complex, northernmost California. Geology, 9, 602-607. ABBOTT, L.D., SILVER, E.A., THOMPSON, PR., FILEWICZ, M.V., SCHNEIDER, C. & ABDOERRIAS, R. 1994. Stratigraphic constraints on the development and timing of arc-continent collision in northern Papua New Guinea. Journal of Sedimentary Research, 64, 169-183. ABDELSALAM, M.G., ABDEEN, M.M., DOWAIDAR, H.M., STERN, RJ. & ABDELGHAFFAR, A.A. 2003. Structural evolution of the Neoproterozoic western Allaqi-Heiani suture, southeastern Egypt. Precambrian Research, 124, 87-104. ACENOLAZA, F.G., MILLER, H. & TOSELLI, A.J. 2002. Proterozoic-Early Paleozoic evolution in western South America: a discussion. Tectonophysics, 354,121-137. ADAMS, C.J. & KELLEY, S. 1998. Provenance of Permian-Triassic and Ordovician metagraywacke terranes in New Zealand: Evidence

from 40Ar/39Ar dating of detrital micas. Geological Society of America Bulletin, 110, 422-432. ADAMS, C.J., BARLEY, M.E., FLETCHER, I.R. & PICKARD, A.L. 1998. Evidence from U-Pb zircon and 40Ar/39Ar muscovite detrital mineral ages in metasandstones for the movement of the Torlesse suspect terrane around the eastern margin of Gondwanaland. Terra Nova, 10, 183-189. ADAMS, C.J., BARLEY, M.E., MAAS, R. & DOYLE, M.G. 2002. Provenance of Permian-Triassic volcaniclastic sedimentary terranes in New Zealand: evidence from their radiogenic isotope characteristics and detrital mineral age patterns. New Zealand Journal of Geology and Geophysics, 45, 221-242. ADAMS, C.J., PANKHURST, R.J., MAAS, R. & MILLAR, I.L. 2005. Nd and Sr isotopic signatures of metasedimentary rocks around the South Pacific margin, and implications for their provenance. In: VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, RJ. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246,113-142. AITCHISON, J.C., ABRAJEVITCH, A., ALI, J.R. ET AL. 2002. New insights into the evolution of the Yarlung Tsangpo suture zone, Xizang (Tibet), China. Episodes, 25, 90-94. ARMADILLO, E., FERRACCIOLI, F, TABELLARIO, G. & Bozzo, E. 2004. Electrical structure across a major ice-covered fault belt in Northern Victoria Land (East Antarctica). Geophysical Research Letters, 31 (10), 1-4, L10615, 10.1029/ 2004GL019903. ASTINI, R.A., BENEDETTO, XL. & VACCARI, N.E. 1995. The Early Paleozoic evolution of the Argentine Precordillera as a Laurentian rifted, drifted, and collided terrane: a geodynamic model. Geological Society of America Bulletin, 107, 253-273. BADARCH, G, CUNNINGHAM, W.D. & WINDLEY, B.F. 2002. A new terrane subdivision for Mongolia: implications for the Phanerozoic crustal growth of Central Asia. Journal of Asian Earth Sciences, 21, 87-110. BANDRES, A., EGUILUZ, L., IBARGUCHI, J.I.G. & PALACIOS, T. 2002. Geodynamic evolution of a Cadomian arc region: the northern Ossa-Morena zone, Iberian massif. Tectonophysics, 352, 105-120. BARKER, F., JONES, D.L., BUDAHN, J.R. & CONEY, P.J. 1988. Ocean plateau-seamount origin of basaltic rocks, Angayucham terrane, Central Alaska. Journal of Geology, 96, 368-374. BARR, S.M. 1990. Granitoid rocks and terrane characterization: an example from the northern Appalachian orogen. Geological Journal, 25, 295-304. BEDARD, J.H. 1999. Petrogenesis of boninites from the Betts Cove Ophiolite, Newfoundland, Canada: Identification of subducted source components. Journal of Petrology, 40,1853-1889. BELASKY, P., STEVENS, C.H. & HANGER, R.A. 2002. Early Permian location of western North American terranes based on brachiopod, fusulinid, and

TERRANE PROCESSES coral biogeography. Palaeogeography, Palaeoclimatology, Palaeoecology, 179, 245-266. BENEDETTO, J.L. 1998. Caradocian brachiopods from the glacigenic diamictite of the Don Braulio Formation (Late Ashgillian), Sierra de Villicum, Argentine Precordillera. Ameghiniana, 35, 243-254. BERG, H.C., JONES, D.L. & RICHTER, D.H. 1972. Gravina-Nutzotin belt-tectonic significance of an upper Mesozoic sedimentary and volcanic sequence in southern and southeastern Alaska. United States Geological Survey Professional Papers, 800-D, D1-D24. BERNASCONI, A. 1987. The major Precambrian terranes of eastern South America: a study of their regional and chronological evolution. Precambrian Research, 37, 107-124. BETTS, P.O., GILES, D., LISTER, G.S. & FRICK, L.R. 2002. Evolution of the Australian lithosphere. Australian Journal of Earth Sciences, 49,661-695. BLAKE, M.C., JR, JONES, D.L. & LANDIS, C.A. 1974. Active continental margins: contrasts between California and New Zealand. In: BURK, C.A. & DRAKE, C.L. (eds) The Geology of Continental Margins. Springer-Verlag, New York, 853-871. BLUCK, B.J. 1990. Terrane provenance and amalgamation: examples from the Caledonides. Philosophical Transactions of the Royal Society of London Series A, 331, 599-609. BLUCK, B.J., HALLIDAY, A.N., AFTALION, M. & MACINTYRE, R.M. 1980. Age and origin of the Ballantrae ophiolite and its significance to the Caledonian orogeny and the Ordovician timescale. Geology, 8, 492-495. BODET, F. & SCHARER, U. 2000. Evolution of the SEAsian continent from U-Pb and Hf isotopes in single grains of zircon and baddeleyite from large rivers. Geochimica et Cosmochimica Ada, 64, 2067-2091. BOGER, S.D., CARSON, C.J., FANNING, C.M., HERGT, J.M., WILSON, C.J.L. & WOODHEAD, J.D. 2002. Pan-African intraplate deformation in the northern Prince Charles Mountains, east Antarctica. Earth and Planetary Science Letters, 195, 195-210. BOGER, S.D., WILSON, C.J.L. & FANNING, C.M. 2001. Early Paleozoic tectonism within the East Antarctic craton: The final suture between east and west Gondwana? Geology, 29, 463-466. BRADSHAW, J.D. 1989. Cretaceous geotectonic patterns in the New Zealand region. Tectonics, 8, 803-820. BRADSHAW, J.D. 1994. Brook Street and Murihiku terranes of New Zealand in the context of a mobile South Pacific Gondwana margin. Journal of South American Earth Sciences, 7, 325-332. BRADSHAW, J.D., ANDREWS, P.B. & ADAMS, C.J. 1981. Carboniferous to Cretaceous on the Pacific margin of Gondwana: the Rangitata Phase of New Zealand. In: CRESWELL, M.M. & VELLA, P. (eds) Gondwana Five. Balkema, Rotterdam, 217-221. BROWN, L.D. 1991. A new map of crustal terranes in the United States from COCORP deep seismic

13

reflection profiling. Geophysical Journal International, 105, 3-13. BRYAN, S.E., HOLCOMBE, R.J. & FIELDING, C.R. 2001. Yarrol terrane of the northern New England Fold Belt: forearc or backarc? Australian Journal of Earth Sciences, 48, 293-316. CAMPBELL, H.J., MORTIMER, N. & RAINE, J.I. 2001. Geology of the Permian Kuriwao Group, Murihiku Terrane, Southland, New Zealand. New Zealand Journal of Geology and Geophysics, 44, 485-498. CARAYON, V., DEWEVER, P. & RAOULT, J.F. 1984. Franciscan limestone knockers from southwestern Oregon: a new interpretation on melange genesis. Comptes Rendus De L'Academie Des Sciences Serie II, 298, 709-714. CASQUET, C, BALDO, E., PANKHURST, R.J., RAPELA, C.W., GALINDO, C, FANNING, C.M. & SAAVEDRA, J. 2001. Involvement of the Argentine Precordillera terrane in the Famatinian mobile belt: U-PbSHRIMP and metamorphic evidence from the Sierra de Pie de Palo. Geology, 29, 703-706. CASQUET, C, PANKHURST, R.J., RAPELA, C.W. ET AL. in press. Grenvillian massif anorthosites in the Sierras Pampeanas. Journal of the Geological Society, London, 162, 9-12. CAWOOD, PA. & LEITCH, E.C. 2002. Terra Australis orogen: Rodinian breakup and subduction initiation in the Pacific Ocean. Geological Society of Australia, Abstracts, 65, 13-15. CAWOOD, PA. & SUHR, G. 1992. Generation and obduction of ophiolites: constraints from the Bay of Islands complex, western Newfoundland. Tectonics, 11, 884-897. CAWOOD, P.A., NEMCHIN, A.A., LEVERENZ, A., SAEED, A. & BALLANCE, P. 1999. U-Pb dating of detrital zircons: implications for the provenance record of Gondwana margin terranes. Geological Society of America Bulletin, 111, 1107-1119. CAWOOD, P. A., LANDIS, C.A., NEMCHIN, A. A. & HADA, S. 2002. Permian fragmentation, accretion and subsequent translation of a low-latitude Tethyan seamount to the high-latitude east Gondwana margin: evidence from detrital zircon age data. Geological Magazine, 139, 131-144. CHANG, C.P., ANGELIER, J., HUANG, C.Y. & Liu, C.S. 2001. Structural evolution and significance of a melange in a collision belt: the Lichi Melange and the Taiwan arc-continent collision. Geological Magazine, 138, 633-651. CHERNICOFF, C.J. & ZAPPETTINI, E.O. 2003. Delimitation of tectonostratigraphic terranes of the southern-central region of Argentina: aeromagnetic evidences. Revista Geologica de Chile, 30, 299-316. CLIFT, P.D. & RYAN, P.D. 1994. Geochemical evolution of an Ordovician island arc, south Mayo, Ireland. Journal of the Geological Society, London, 151, 329-342. CLOOS, M. 1983. Comparative study of melange matrix and metashales from the Franciscan subduction complex with the basal Great Valley sequence, California. Journal of Geology, 91, 291-306.

14

A. P. M. VAUGHAN ET AL.

COCKS, L.R.M. & TORSVIK, T.H. 2002. Earth geography from 500 to 400 million years ago: a faunal and palaeomagnetic review. Journal of the Geological Society, London, 159, 631-644. COLLINS, WJ. 1991. A reassessment of the Hunter-Bowen orogeny: tectonic implications for the southern New England fold belt. Australian Journal of Earth Sciences, 38,409-423. CONDIE,K.C. & CHOMIAK,B. 1996. Continental accretion: contrasting Mesozoic and Early Proterozoic tectonic regimes in North America. Tectonophysics, 265, 101-126. CONEY, PJ. 1978. Mesozoic-Cenozoic Cordilleran plate tectonics. In: SMITH, R.B. & EATON, G.P. (eds) Cenozoic Tectonics and Regional Geophysics of the Western Cordillera. Geological Society of America Memoirs, 152, 33-50. CONEY, P.J., JONES, D.L. & MONGER, J.W.H. 1980. Cordilleran suspect terranes. Nature, 288, 329-333. CONEY, P.J., EDWARDS, A., HINE, R., MORRISON, F. & WINDRIM, D. 1990. The regional tectonics of the Tasman erogenic system, eastern Australia. Journal of Structural Geology, 12, 519-543. COOMBS, D.S. 1997. A note on the terrane concept, based on a introduction to the Terrane '97 conference, Christchurch, New Zealand, February 1997. American Journal of Science, 297, 762-764. COOMBS, D.S., LANDIS, C.A., NORRIS, R.J., SINTON, J.M., BORNS, D.J. & CRAW, D. 1976. The Dun Mountain ophiolite belt, New Zealand, its tectonic setting, constitution, and origin, with special reference to the southern portion. American Journal of Science, 276, 561-603. CORDANI, U.G., CARDONA, A., JIMENEZ, D.M., Liu, D. & NUTMAN, A.P. 2005. Geochronology of Proterozoic basement inliers in the Colombian Andes: tectonic history of remnants of a fragmented Grenville belt. In: VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, R.J. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 329-346. COWAN, D.S., BRANDON, M.T. & GARVER, J.I. 1997. Geologic tests of hypotheses for large coastwise displacements: a critique illustrated by the Baja British Columbia controversy. American Journal of Science, 297,117-173. CURTIS, M.L. 2001. Tectonic history of the Ellsworth Mountains, West Antarctica: reconciling a Gondwana enigma. Geological Society of America Bulletin, 113, 939-958. DALZIEL, I.W.D. & ELLIOT, D.H. 1982. West Antarctica: problem child of Gondwanaland. Tectonics, 1, 3-19. BANNER, W.R. 1970. Paleontologic and stratigraphic evidence for and against sea floor spreading and opening and closing oceans in the Pacific Northwest. Geological Society of America Abstracts with Programs, 2, 84-85. DEBARI, S.M. & SLEEP, N.H. 1991. High-Mg, low-Al bulk composition of the Talkeetna island arc, Alaska: implications for primary magmas and the

nature of arc crust. Geological Society of America Bulletin, 103, 37-47. DEFANT, M.J., KEPEZHINSKAS, P., Xu, J.F., WANG, Q., ZHANG, Q. & XIAO, L. 2002. Adakites: some variations on a theme. Ada Petrologica Sinica, 18, 129-142. DIREEN, N.G. & CRAWFORD, A.J. 2003. Fossil seawarddipping reflector sequences preserved in southeastern Australia: a 600 Ma volcanic passive margin in eastern Gondwanaland. Journal of the Geological Society, London, 160, 985-990. DOUBLEDAY, PA., LEAT, P.T., ALABASTER, T, NELL,

P.A.R. & TRANTER, T.H. 1994. Allochthonous oceanic basalts within the Mesozoic accretionary complex of Alexander Island, Antarctica: remnants of proto-Pacific oceanic crust. Journal of the Geological Society, London, 151, 65-78. ERNST, WG. 1993. Metamorphism of Franciscan tectonostratigraphic assemblage, Pacheco Pass area, east-central Diablo Range, California Coast Ranges. Geological Society of America Bulletin, 105, 618-636. FANG, Z.-J., BOUCOT, A., COVACEVICH, V. & HERVE, F. 1998. Discovery of Late Triassic fossils in the Chonos Metamorphic Complex, Southern Chile. Revista Geologica de Chile, 25,165-173. FAUGERES, J.C., GONTHIER, E., GRIBOULARD, R. & MASSE, L. 1993. Quaternary sandy deposits and canyons on the Venezuelan margin and south Barbados accretionary prism. Marine Geology, 110,115-142. FERGUSSON, C.L. 2003. Ordovician-Silurian accretion tectonics of the Lachlan Fold Belt, southeastern Australia. Australian Journal of Earth Sciences, 50, 475-490. FERGUSSON, C.L. & FANNING, C.M. 2002. Late Ordovician stratigraphy, zircon provenance and tectonics, Lachlan Fold Belt, southeastern Australia. Australian Journal of Earth Sciences, 49, 423-436. FERRACCIOLI, F. & Bozzo, E. 1999. Inherited crustal features and tectonic blocks of the Transantarctic Mountains: An aeromagnetic perspective (Victoria Land, Antarctica). Journal of Geophysical Research, 104, 25297-25319. FERRACCIOLI, F, Bozzo, E. & CAPPONI, G. 2002. Aeromagnetic and gravity anomaly constraints for an early Paleozoic subduction system of Victoria Land, Antarctica. Geophysical Research Letters, 29,1-4,1406,10.1029/2001GL014138. FINNEY, S.C., PERALTA, S.H., GEHRELS, GE. & MARSAGLIA, K.M. 2005. The Early Paleozoic history of the Cuyania (greater Precordillera) terrane of western Argentina: evidence from geochronology of detrital zircons from Middle Cambrian sandstones. Geologica Acta, in press. FISKE, R.S., NAKA, J., IIZASA, K., YUASA, M. & KLAUS, A. 2001. Submarine silicic caldera at the front of the Izu-Bonin arc, Japan: voluminous seafloor eruptions of rhyolite pumice. Geological Society of America Bulletin, 113, 813-824. FlTZHERBERT, J.A., CLARKE, G.L.,

MARMO, B. &

POWELL, R. 2004. The origin and P-T evolution of peridotites and serpentinites of NE New

TERRANE PROCESSES

15

Caledonia: prograde interaction between conti- GOWER, C.F & KROGH, T.E. 2002. A U-Pb nental margin and the mantle wedge. Journal of geochronological review of the Proterozoic Metamorphic Geology, 22, 327-344. history of the eastern Grenville Province. FITZSIMONS, I.C.W. 20000. Grenville-aged basement Canadian Journal of Earth Sciences, 39, 795-829. provinces in East Antarctica: evidence for at least GRIFFIN, W.L., BELOUSOVA, E.A., SHEE, S.R., PEARSON, NJ. & O'REILLY, S.Y. 2004. Archean three separate collisional orogens. Geology, 28, 879-882. crustal evolution in the northern Yilgam Craton: FITZSIMONS, I.C.W. 2000b. A review of tectonic events U-Pb and Hf-isotope evidence from detrital in the East Antarctic Shield and their implizircons. Precambrian Research, 131, 231-282. cations for Gondwana and earlier superconti- HALL, R. 2002. Cenozoic geological and plate tectonic nents. Journal of African Earth Sciences, 31, 3-23. evolution of SE Asia and the SW Pacific: FLOTTMANN, T., GIBSON, G.M. & KLEINSCHMIDT, G. computer-based reconstructions, model and 1993. Structural continuity of the Ross and animations. Journal of Asian Earth Sciences, 20, Delamerian orogens of Antarctica and Australia 353-431. along the margin of the proto-Pacific. Geology, HANDLER, M.R., WYSOCZANSKI, RJ. & GAMBLE, J.A. 21, 319-322. 2003. Proterozoic lithosphere in Marie Byrd FOSTER, D.A. & GRAY, D.R. 2000. Evolution and Land, West Antarctica: Re-Os systematics of structure of the Lachlan fold belt (orogen) of spinel peridotite xenoliths. Chemical Geology, eastern Australia. Annual Review of Earth and 196,131-145. Planetary Sciences, 28, 47-80. HARTZ, E.H. & TORSVIK, T.H. 2002. Baltica upside FOUQUET, Y., VONSTACKELBERG, U, CHARLOU, J.L. down: A new plate tectonic model for Rodinia ET AL. 1991. Hydrothermal activity and metalloand the lapetus Ocean. Geology, 30, 255-258. genesis in the Lau back-arc basin. Nature, 349, HEATHERINGTON, A.L. & MULLER, PA. 2003. Mesozoic 778-781. igneous activity in the Suwannee terrane, southFRIEDL, G, FINGER, F, MCNAUGHTON, N.J. & eastern USA: petrogenesis and Gondwanan affiniFLETCHER, I.R. 2000. Deducing the ancestry of ties. Gondwana Research, 6, 296-311. terranes: SHRIMP evidence for South America- HERRMANN, U.R., NELSON, B.K. & RATSCHBACHER, L. derived Gondwana fragments in central Europe. 1994. The origin of a terrane: U-Pb zircon Geology, 28,1035-1038. geochronology and tectonic evolution of the FRIEND, C.R.L., NUTMAN, A.P. & MCGREGOR, V.R. Xolapa complex (southern Mexico). Tectonics, 1988. Late Archean terrane accretion in the 13, 455-474. Godthab region, southern West Greenland. HERVE, F. 1988. Late Paleozoic subduction and accreNature, 335, 535-538. tion in Southern Chile. Episodes, 11,183-188. FRITZ, H. 1996. Geodynamic and tectonic evolution of HERVE, F. & FANNING, C.M. 2001. Late Triassic detrital the southeastern Bohemian Massif: The Thaya zircons in meta-turbidites of the Chonos Metasection (Austria). Mineralogy and Petrology, 58, morphic Complex, southern Chile. Revista 253-278. Geologica de Chile, 28, 91-104. FUH, S.-C, LlU, C.-S., LUNDBERG, N. & REED, D.L. HERVE, F. & FANNING, C.M. 2003. Early Cretaceous 1997. Strike-slip faults offshore southern Taiwan: subduction of continental crust at the Diego de implications for the oblique arc-continent Almagro archipelago, southern Chile. Episodes, collision process. Tectonophysics, 274, 25-39. 26, 285-288. GAETANI, M., ZANCHI, A., ANGIOLINI, L., OLIVINI, G, HERVE, F, FANNING, C.M. & PANKHURST, RJ. 2003. SCIUNNACH, D., BRUNTON, H., NICORA, A. & Detrital zircon age patterns and provenance of MAWSON, R. 2004. The Carboniferous of the the metamorphic complexes of southern Chile. Western Karakoram (Pakistan). Journal of Asian Journal of South American Earth Sciences, 16, Earth Sciences, 23, 275-305. 107-123. GARDNER, M.C., BERGMAN, S.C., GUSHING, G.W HERZIG, C.T. & SHARP, W.D. 1992. The Sullivan Creek ET AL. 1988. Pennsylvanian pluton stitching of terrane: a composite Jurassic arc assemblage, Wrangellia and the Alexander terrane, Wrangell western Sierra Nevada metamorphic belt, CaliMountains, Alaska. Geology, 16, 967-971. fornia. Geological Society of America Bulletin, GLEN, R.A. 2005. The Tasmanides of Eastern 104, 1292-1300. Australia: 600 million years of interaction between HIBBARD, J. 2000. Docking Carolina: mid-Paleozoic the proto-Pacific plate and the Australian sector accretion in the southern Appalachians. Geology, of Gondwana. In: VAUGHAN, A.P.M., LEAT, P.T. & 28, 127-130. PANKHURST, RJ. (eds) Terrane Processes at the HIBBARD, J., MILLER, B., TRACY, R. & CARTER, B. Margins of Gondwana. Geological Society, 2005. The Appalachian peri-Gondwanan realm: a London, Special Publications, 246, 23-96. palaeogeographic perspective from the south. In: GLEN, R.A., WALSHE, J.L., BARRON, L.M. & WATKINS, VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, RJ. JJ. 1998. Ordovician convergent-margin volcan(eds) Terrane Processes at the Margins of ism and tectonism in the Lachlan sector of east Gondwana. Geological Society, London, Special Gondwana. Geology, 26, 751-754. Publications, 246, 97-112. GOLDFARB, R.J., GROVES, D.L & GARDOLL, S. 2001. HILL, M.L. & DIBBLEE, T.W, JR. 1953. San Andreas, Orogenic gold and geologic time: a global syntheGarlock, and Big Pine faults, California: a study sis. Ore Geology Reviews, 18, 1-75. of the character, history, and tectonic significance

16

A. P. M. VAUGHAN ET AL.

of their displacements. Geological Society of America Bulletin, 64, 443-458. HOWELL, D.G. 1980. Mesozoic accretion of exotic terranes along the New Zealand segment of Gondwanaland. Geology, 8, 487-491. HOWELL, D.G. 1989. Tectonics of Suspect Terranes: Mountain Building and Continental Growth. Chapman & Hall, London. HOWELL, K. & HOWELL, D.G. 1995. Principles of Terrane Analysis: New Applications for Global Tectonics. Chapman & Hall, New York. HOWELL, D.G, JONES, D.L. & SCHERMER, E.R. 1985. Tectonostratigraphic terranes of the CircumPacific region. In: HOWELL, D.G. (ed.) Tectonostratigraphic Terranes of the Circum-Pacific Region. Council for Energy and Mineral Resources, Earth Science Series, 1, 3-30. HUTTON, D.H.W. 1987. Strike-slip terranes and a model for the evolution of the British and Irish Caledonides. Geological Magazine, 124,405-425. IlZASA, K., FlSKE, R.S., ISHIZUKA, O. ET AL. 1999. A

Kuroko-type polymetallic sulfide deposit in a submarine silicic caldera. Science, 283, 975-977. IRELAND, T.R., FLOTTMANN, T., FANNING, C.M., GIBSON, G.M. & PREISS, W.V. 1998. Development of the early Paleozoic Pacific margin of Gondwana from detrital-zircon ages across the Delamerian orogen. Geology, 26, 243-246. IRWIN, W.P. 1964. Late Mesozoic orogenies in the ultramafic belts of northwestern California and southwestern Oregon United States Geological Survey Professional Papers, 501-C, C1-C9. IRWIN, W.P. 1972. Terranes of the western Paleozoic and Triassic belt in the southern Klamath Mountains, California United States Geological Survey Professional Papers, 800-C, C103-C111. ISHIBASHI, J. & URABE,T. 1995. Hydrothermal activity related to arc-backarc magmatism in the Western Pacific. In: TAYLOR, B. (ed.) Backarc Basins: Tectonics and Magmatism. Plenum Press, New York, 451-495. JACOBI, R.D. & WASOWSKI, JJ. 1985. Geochemistry and plate tectonic significance of the volcanic rocks of the Summerford Group, north-central Newfoundland. Geology, 13,126-130. JACOBS, I, FANNING, C.M. & BAUER, W 2003. Timing of Grenville-age vs Pan-African medium- to high grade metamorphism in western Dronning Maud Land (East Antarctica) and significance for correlations in Rodinia and Gondwana. Precambrian Research, 125,1-20. JOHNSON, P.R., ABDELSALAM, M.G. & STERN, RJ. 2003. The Bi'r Umq-Nakasib suture zone in the Arabian-Nubian shield: a key to understanding crustal growth in the East African orogen. Gondwana Research, 6, 523-530. JOHNSTON, S.T. 2001. The great Alaskan terrane wreck: reconciliation of paleomagnetic and geological data in the northern Cordillera. Earth and Planetary Science Letters, 193, 259-272. KAMP, P.J.J. 2000. Thermochronology of the Torlesse accretionary complex, Wellington region, New Zealand. Journal of Geophysical Research, 105, 19253-19272.

KAPP, P., YIN, A., MANNING, C.E., HARRISON, T.M., TAYLOR, M.H. & DING, L. 2003. Tectonic evolution of the early Mesozoic blueschistbearing Qiangtang metamorphic belt, central Tibet. Tectonics, 22 (4), 1-25, 1043, 10.1029/ 2002TC001383. KAWATE, S. & ARIMA, M. 1998. Petrogenesis of the Tanzawa plutonic complex, central Japan: exposed felsic middle crust of the Izu-Bonin-Mariana arc. Island Arc, 7, 342-358. KAY, R.W. & KAY, S.M. 2002. Andean adakites: three ways to make them. Acta Petrologica Sinica, 18, 303-311. KELLY, S.R.A., DOUBLEDAY, PA., BRUNTON, C.H.C., DICKINS, J.M., SEVASTOPULO, G.D. & TAYLOR, P.D. 2001. First Carboniferous and ?Permian marine macrofaunas from Antarctica and their tectonic implications. Journal of the Geological Society, London, 158, 219-232. KENNEDY, W.Q. 1946. The Great Glen Fault. Quarterly Journal of the Geological Society of London, 102, 41-76. KEPPIE, J.D. 19850. The Appalachian Collage. In: GEE, D.G & STURT, B.A. (eds) The Caledonide Orogen - Scandinavia and Related Areas. John Wiley, Chichester, 1217-1226. KEPPIE, J.D. 1985ft. Geology and tectonics of Nova Scotia. In: KEPPIE, J.D., CURRIE, K., MURPHY, J.B., PICKERILL, R.K., FYFFE, L. & ST. JULIAN, P. (eds) Excursion 1: Appalachian Geotraverse. Geological Association of Canada/Mineralogical Association of Canada, Fredericton, 1-181. KEPPIE, J.D. & DALLMEYER, R.D. 1987. Dating transcurrent terrane accretion: an example from the Meguma and Avalon composite terranes in the northern Appalachians. Tectonics, 6, 831-847. KEPPIE, J.D. & DOSTAL, J. 2001. Evaluation of the Baja controversy using paleomagnetic and faunal data, plume magmatism, and piercing points. Tectonophysics, 339, 427-442. KEPPIE, J.D., NANCE, R.D., MURPHY, IB. & DOSTAL, J. 2003. Tethyan, Mediterranean, and Pacific analogues for the Neoproterozoic-Paleozoic birth and development of peri- Gondwanan terranes and their transfer to Laurentia and Laurussia. Tectonophysics, 365, 195-219. KERRICH, R., GOLDFARB, R., GROVES, D., GARWIN, S. & JIA, Y. 2000. The characteristics, origins, and geodynamic settings of supergiant gold metallogenic provinces. Science in China Series D-Earth Sciences, 43,1-68. KIMURA, G, SAKAKIBARA, M. & OKAMURA, M. 1994. Plumes in central Panthalassa: deductions from accreted oceanic fragments in Japan. Tectonics, 13,905-916. KNUDSEN, T.L., GRIFFIN, W.L., HARTZ, E.H., ANDRESEN, A. & JACKSON, S.E. 2001. In-situ hafnium and lead isotope analyses of detrital zircons from the Devonian sedimentary basin of NE Greenland: a record of repeated crustal reworking. Contributions to Mineralogy and Petrology, 141, 83-94. KOTTACHCHI, N, SCHRODER-ADAMS, C.J., HAGGART, J.W. & TIPPER, H.W. 2002. Jurassic foraminifera

TERRANE PROCESSES from the Queen Charlotte Islands, British Columbia, Canada: biostratigraphy, paleoenvironments and paleogeographic implications. Palaeogeography, Palaeodimatology, Palaeoecology, 180, 93-127. KUSKY, T.M. & BRADLEY, D.C. 1999. Kinematic analysis of melange fabrics: examples and applications from the McHugh Complex, Kenai Peninsula, Alaska. Journal of Structural Geology, 21, 1773-1796. LAIRD, M.G. & BRADSHAW, J.D. 2004. The break-up of a long-term relationship: the Cretaceous separation of New Zealand from Gondwana. Gondwana Research, 7, 273-286. LAPIERRE, H., TARDY, M., COULON, C, HERNANDEZ, E.O., BOURDIER, J.L., REYES, J.M. & FREYDIER, C. 1992. Characterization, genesis and geodynamic evolution of Guerrero terrane (western Mexico). Canadian Journal of Earth Sciences, 29, 2478-2489. LAPIERRE, H., BOSCH, D., TARDY, M. & STRUIK, L.C. 2003. Late Paleozoic and Triassic plume-derived magmas in the Canadian Cordillera played a key role in continental crust growth. Chemical Geology, 201 (1-2), 55-89. LEAT, P.T., DEAN, A.A., MILLAR, I.L., KELLEY, S.P., VAUGHAN, A.P.M. & RILEY, T.R. 2005. Lithospheric mantle domains beneath Antarctica. In: VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, RJ. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 359-380. LEDO, I, JONES, A.G, FERGUSON, I.J. & WOLYNEC, L. 2004. Lithospheric structure of the Yukon, northern Canadian Cordillera, obtained from magnetotelluric data. Journal of Geophysical Research, 109 (B4), 1-19, B04410, 10.1029/ 2003JB002516. LEGGETT, J.K. 1987. The Southern Uplands as an accretionary prism: the importance of analogues in reconstructing palaeogeography. Journal of the Geological Society, London, 144, 737-752. LEHMANN, D., BRETT, C.E., COLE, R. & BAIRD, G. 1995. Distal sedimentation in a peripheral foreland basin: Ordovician black shales and associated flysch of the western Taconic foreland, New York State and Ontario. Geological Society of America Bulletin, 107, 708-724. LEITCH, E.C., FERGUSSON, C.L. & HENDERSON, R.A. 2003. Arc to craton provenance switching in a Late Palaeozoic subduction complex, Wandilla and Shoalwater terranes, New England Fold Belt, eastern Australia. Australian Journal of Earth Sciences, 50, 919-929. LEVERENZ, A. & BALLANCE, P.F. 2001. Terrane affiliation and terrane boundaries of Mesozoic accretionary complexes, northeastern North Island, New Zealand: some implications from recycled elastics. New Zealand Journal of Geology and Geophysics, 44, 589-599. LUCASSEN, F. & FRANZ, G. 2005. The early Palaeozoic orogen in the Central Andes: a non-collisional orogen comparable to the Cenozoic high plateau? In: VAUGHAN, A.P.M., LEAT, P.T. &

17

PANKHURST, RJ. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 257-274. MACDONALD, A.I, LEWIS, P.D., THOMPSON, J.F.H. ET AL. 1996. Metallogeny of an early to middle Jurassic arc, Iskut river area, northwestern British Columbia. Economic Geology, 91, 1098-1114. MAEKAWA, H., YAMAMOTO, K., UENO,T, OSADA, Y. & NOGAMI, N. 2004. Significance of serpentinites and related rocks in the high-pressure metamorphic terranes, circum-Pacific regions. International Geology Review, 46, 426-444. MANKINEN, E.A., IRWIN, W.P & BLOME, CD. 1996. Far-travelled Permian chert of the North Fork terrane, Klamath Mountains, California. Tectonics, 15, 314-328. MARSAGLIA, K.M. & INGERSOLL, R.V. 1992. Compositional trends in arc-related, deep-marine sand and sandstone: a reassessment of magmatic arc provenance. Geological Society of America Bulletin, 104, 1637-1649. MATTE, P. 2001. The Variscan collage and orogeny (480-290 Ma) and the tectonic definition of the Armorica microplate: a review. Terra Nova, 13, 122-128. MCCLELLAND, W.C., GEHRELS, G.E. & SALEEBY, J.B. 1992. Upper Jurassic-Lower Cretaceous basinal strata along the Cordilleran margin: implications for the accretionary history of the AlexanderWrangellia-Peninsular terrane. Tectonics, 11, 823-835. MCELHINNY, M.W., POWELL, C.M. & PISAREVSKY, S.A. 2003. Paleozoic terranes of eastern Australia and the drift history of Gondwana. Tectonophysics, 362, 41-65. METCALFE, 1.1994. Gondwanaland origin, dispersion, and accretion of east and Southeast Asian continental terranes. Journal of South American Earth Sciences, 7, 333-347. MILLAR, I.L. & PANKHURST, RJ. 1987. Rb-Sr geochronology of the region between the Antarctic Peninsula and the Transantarctic Mountains: Haag Nunataks and Mesozoic granitoids. In: McKENZiE, G.D. (ed.) Gondwana Six: Structure, Tectonics, and Geophysics. Geophysical Monographs, 40, 151-160. MILLAR, I.L., WILLAN, R.C.R., WAREHAM, C.D. & BOYCE, A.J. 2001. The role of crustal and mantle sources in the genesis of granitoids of the Antarctic Peninsula and adjacent crustal blocks. Journal of the Geological Society, London, 158, 855-867. MILLAR, I.L., PANKHURST, RJ. & FANNING, C.M. 2002. Basement chronology of the Antarctic Peninsula: recurrent magmatism and anatexis in the Palaeozoic Gondwana margin. Journal of the Geological Society, London, 159, 145-157. MILLER, D J. & CHRISTENSEN, N.I. 1994. Seismic signature and geochemistry of an island arc: a multidisciplinary study of the Kohistan accreted terrane, northern Pakistan. Journal of Geophysical Research-Solid Earth, 99 (B6), 11623-11642. MILLER, H. & SOLLNER, F. 2005. The Famatina complex (NW Argentina): back-docking of an

18

A. P. M. VAUGHAN ET AL.

island arc or terrane accretion? Early Palaeozoic geodynamics at the western Gondwana margin. In: VAUGHAN, A.P.M., LEAT, PT. & PANKHURST, R.J. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 241-256. MILSOM, I, An, J. & SUDARWONO 1999. Structure and collision history of the Buton continental fragment, eastern Indonesia. AAPG Bulletin, 83, 1666-1689. MONGER, J.W.H. 1997. Plate tectonics and northern cordilleran geology: an unfinished revolution. Geoscience Canada, 24, 189-198. MONGER, J.W.H., SOUTHER, J.G. & GABRIELSE, H. 1972. Evolution of the Canadian Cordillera: a plate tectonic model. American Journal of Science, 272, 577-602. MOORE, G.W. 1994. Terranes and the tectonic assembly of South America: the fifth circumPacific terrane conference. Journal of South American Earth Sciences, 1, 235-239. MORTIMER, N. 2004. New Zealand's geological foundations. Gondwana Research, 7, 261-272. MORTIMER, N. & COOPER, A.F. 2004. U-Pb and Sm-Nd ages from the Alpine Schist, New Zealand. New Zealand Journal of Geology and Geophysics, 47, 21-28. MORTIMER, N. & PARKINSON, D. 1996. Hikurangi Plateau: a Cretaceous large igneous province in the southwest Pacific Ocean. Journal of Geophysical Research, 101 (Bl), 687-696. MORTIMER, N., TULLOCH, A.J., SPARK, R.N., WALKER, N.W., LADLEY, E., ALLIBONE, A. & KIMBROUGH, D.L. 1999. Overview of the Median batholith, New Zealand: a new interpretation of the geology of the Median Tectonic Zone and adjacent rocks. Journal of African Earth Sciences, 29, 257-268. MPODOZIS, C. & FORSYTHE, R. 1983. Stratigraphy and geochemistry of accreted fragments of the ancestral Pacific floor in southern South America. Palaeogeography, Palaeoclimatology, Palaeoecology, 41, 103-124. MUIR, R.J., IRELAND, T.R., WEAVER, S.D., BRADSHAW, J.D., EVANS, J.A., EBY, G.N. & SHELLEY, D. 1998. Geochronology and geochemistry of a Mesozoic magmatic arc system, Fiordland, New Zealand. Journal of the Geological Society, London, 155, 1037-1052. MURPHY, J.B. & KEPPIE, J.D. 2003. Collisional orogenesis in the geological record and modern analogues. Tectonophysics, 365,1-5. MURPHY, J.B. & NANCE, R.D. 2003. Do supercontinents introvert or extrovert?: Sm-Nd isotope evidence. Geology, 31, 873-876. MURRAY, C.G., FERGUSSON, C.L., FLOOD, P.G., WHITAKER, WG. & KORSCH, R.J. 1987. Plate tectonic model for the Carboniferous evolution of the New England fold belt. Australian Journal of Earth Sciences, 34, 213-236. NANCE, R.D., MURPHY, J.B. & KEPPIE, J.D. 2004. Nick Rast and the recognition of the Avalonian Arc. Journal of Geodynamics, 37, 437-455. NELSON, E., FORSYTHE, R. & ARIT, I. 1994. Ridge

collision tectonics in terrane development. Journal of South American Earth Sciences, 7, 271-278. NICHOLS, G.J. & CANTRILL, D.J. 2002. Tectonic and climatic controls on a Mesozoic forearc basin succession, Alexander Island, Antarctica. Geological Magazine, 139, 313-330. NODA, A., TAKEUCHI, M. & ADACHI, M. 2004. Provenance of the Murihiku Terrane, New Zealand: evidence from the Jurassic conglomerates and sandstones in Southland. Sedimentary Geology, 164, 203-222. O'BRIEN, S.J., WARDLE, R.J. & KING, A.F. 1983. The Avalon zone: a Pan-African terrane in the Appalachian orogen of Canada. Geological Journal, 18, 195-222. O'SULLIVAN, P.B., FOSTER, DA., KOHN, B.P. & GLEADOW, A.J.W. 1996. Multiple postorogenic denudation events: an example from the eastern Lachlan fold belt, Australia. Geology, 24,563-566. PANKHURST, R.J., WEAVER, S.D., BRADSHAW, J.D., STOREY, B.C. & IRELAND, T.R. 1998. Geochronology and geochemistry of pre-Jurassic superterranes in Marie Byrd Land, Antarctica. Journal of Geophysical Research, 103 (B2), 2529-2547. PARFENOV, L.M., NOKLEBERG, WJ. & KHANCHUK, A.I. 2000. Compilation principles and the main units of the legend for the geodynamics map of north and central Asia, Russian southern Far East, Korea and Japan. Geology of the Pacific Ocean, 15, 463-482. PAVLIS,T.L., PICORNELL, C, SERPA, L., BRUHN, R.L. & PLAFKER, G. 2004. Tectonic processes during oblique collision: insights from the St. Elias orogen, northern North American Cordillera. Tectonics, 23, 1-14, TC3001, 10.10297 2003TC001557. PEDROSA-SOARES, A.C., NOCE, C.M., WlEDEMANN,

C.M. & PINTO, C.P. 2001. The Aracuai-West Congo Orogen in Brazil: an overview of a confined orogen formed during Gondwanaland assembly. Precambrian Research, 110, 307-323. PHILLIPS, W.E.A., STILLMAN, C.J. & MURPHY, T. 1976. A Caledonian plate tectonic model. Journal of the Geological Society, London, 132, 579-609. PICKERING, K.T., BASSETT, M.G. & SIVETER, D.J. 1988. Late Ordovician-early Silurian destruction of the lapetus Ocean: Newfoundland, British Isles and Scandinavia: a discussion. Transactions of the Royal Society of Edinburgh-Earth Sciences, 79, 361-382. PIERCEY, S.J., MURPHY, D.C., MORTENSEN, J.K. & CREASER, R.A. 2004. Mid-Paleozoic initiation of the northern Cordilleran marginal back-arc basin: geologic, geochemical, and neodymium isotope evidence from the oldest mafic magmatic rocks in the Yukon-Tanana terrane, Finlayson Lake district, southeast Yukon, Canada. Geological Society of America Bulletin, 116, 1087-1106. PIGRAM, CJ. & DAVIES, H.L. 1987. Terranes and the accretion history of the New Guinea orogen. Bureau of Mineral Resources Journal of Australian Geology and Geophysics, 10,193-211.

TERRANE PROCESSES POLAT, A. & KERRICH, R. 2001. Geodynamic processes, continental growth, and mantle evolution recorded in late Archean greenstone belts of the southern Superior Province, Canada. Precambrian Research, 112, 5-25. RAESIDE, R.P. & BARR, S.M. 1990. Geology and tectonic development of the Bras D'Or suspect terrane, Cape Breton Island, Nova Scotia. Canadian Journal of Earth Sciences, 27, 1371-1381. RAMOS, V.A. 1988. Late Proterozoic-Early Paleozoic of South America: a collisional history. Episodes, 11,168-174. RAMOS, V.A., JORDAN, T.E., ALLMENDINGER, R.W., MPODOZIS, G, KAY, S.M., CORTES, J.M. & PALMA, M. 1986. Paleozoic terranes of the central Argentine-Chilean Andes. Tectonics, 5, 855-880. RANKIN, D.W., FURNES, H., BISHOP, A.C., CABANIS, B., MILTON, D.J., O'BRIEN, SJ. & THORPE, R.S. 1988. Plutonism and volcanism related to the preArenig evolution of the CaledonideAppalachian orogen. In: HARRIS, A.L. & FETTES, DJ. (eds) The Caledonian-Appalachian Orogen. Geological Society, London, Special Publications, 38, 149-183. RAPALINI, A.E. 2005. The accretionary history of southern South America from the latest Proterozoic to the Late Palaeozoic: some palaeomagnetic constraints. In: VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, R.J. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 305-328. RAPELA, C.W., PANKHURST, R.J., FANNING, C.M. & HERVE, F. 2005. Pacific subduction coeval with the Karoo mantle plume: the Early Jurassic Subcordilleran belt of northwestern Patagonia. In: VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, R.J. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 217-240. READING, A.M. 2005. Investigating the deep structure of terranes and terrane boundaries: insights from earthquake seismic data. In: VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, R.J. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 293-304. READING, A.M., KENNETT, B.L.N. & DENTITH, M.C. 2003. Seismic structure of the Yilgarn Craton, Western Australia. Australian Journal of Earth Sciences, 50, 427-438. RESTREPOPACE, P.A. 1992. Petrotectonic characterization of the central Andean terrane, Colombia. Journal of South American Earth Sciences, 5, 97-116. RICHARDS, J.P. & KERRICH, R. 1993. The Porgera gold mine, Papua New Guinea: magmatic hydrothermal to epithermal evolution of an alkalic-type precious-metal deposit. Economic Geology, 88, 1017-1052. RICHARDS, M.A., JONES, D.L., DUNCAN, R.A. & DEPAOLO, D.J. 1991. A mantle plume initiation model for the Wrangellia flood basalt and other oceanic plateaus. Science, 254, 263-267.

19

ROBERTS, D. 2003. The Scandinavian Caledonides: event chronology, palaeogeographic settings and likely, modern analogues. Tectonophysics, 365, 283-299. ROBERTSON, A.H.F. & USTAOMER, T 2004. Tectonic evolution of the intra-pontide suture zone in the Armutlu Peninsula, NW Turkey. Tectonophysics, 381, 175-209. RUBIN, C.M. & SALEEBY, J.B. 1991. The Gravina sequence: remnants of a Mid-Mesozoic oceanic arc in southern Southeast Alaska. Journal of Geophysical Research, 96 (B9), 14551-14568. RYAN, P.D. & DEWEY, J.F. 1997. Continental eclogites and the Wilson Cycle. Journal of the Geological Society, London, 154, 437-442. RYAN, R.J. & SMITH, P.K. 1998. A review of the mesothermal gold deposits of the Meguma Group, Nova Scotia, Canada. Ore Geology Reviews, 13, 153-183. SAMSON, S.D., PATCHETT, P.J., GEHRELS, G.E. & ANDERSON, R.G 1990. Nd and Sr isotopic characterization of the Wrangellia terrane and implications for crustal growth of the Canadian Cordillera. Journal of Geology, 98, 749-762. SAMSON, S.D., PATCHETT, P.J., MCCLELLAND, WC. & GEHRELS, G.E. 1991. Nd isotopic characterization of metamorphic rocks in the Coast Mountains, Alaskan and Canadian Cordillera: ancient crust bounded by juvenile terranes. Tectonics, 10, 770-780. SCARROW, J.H., AYALA, C. & KIMBELL, G.S. 2002. Insights into orogenesis: getting to the root of a continent-ocean-continent collision, Southern Urals, Russia. Journal of the Geological Society, London, 159, 659-671. SCHWARTZ, J. J. & GROMET, L.P. 2004. Provenance of a late Proterozoic-early Cambrian basin, Sierras de Cordoba, Argentina. Precambrian Research, 129, 1-21. SCOTT, S.D. 2001. Deep ocean mining. Geoscience Canada, 28, 87-96. SENGOR, A.M.C. 1990. Lithotectonic terranes and the plate tectonic theory of orogeny: A critique of the principles of terrane analysis. In: WILEY, T.J., HOWELL, D.G. & WONG, F.L. (eds) Terrane Analysis of China and the Pacific Rim. CircumPacific Council for Energy and Mineral Resources Earth Science Series, 13, 9-44. SENGOR, A.M.C. & DEWEY, J.F. 1991. Terranology: vice or virtue? In: DEWEY, J.F, GASS, I.G., CURRY, G.B., HARRIS, N.B.W & SENGOR, A.M.C. (eds) Allochthonous Terranes. Cambridge University Press, Cambridge, 1-21. SENGOR, A.M.C. & NATALIN, B.A. 1996. Turkic-type orogeny and its role in the making of the continental crust. Annual Review of Earth and Planetary Sciences, 24, 263-337. SENGOR, A.M.C. & OKUROGULLARI, A.H. 1991. The role of accretionary wedges in the growth of continents: Asiatic examples from Argand to plate tectonics. Eclogae Geologicae Helvetiae, 84, 535-597. SENGOR, A.M.C., NATALIN, B.A. & BURTMAN, V.S. 1993. Evolution of the Altaid tectonic collage and

20

A. P. M. VAUGHAN ETAL.

Paleozoic crustal growth in Eurasia. Nature, 364, 299-307. SHERLOCK, R.L., ROTH, T., SPOONER, E.T.C. & BRAY, CJ. 1999. Origin of the Eskay Creek precious metal-rich volcanogenic massive sulfide deposit: fluid inclusion and stable isotope evidence. Economic Geology, 94, 803-824. SHERVAIS, J.W., KIMBROUGH, D.L., RENNE, P. ET AL. 2004. Multi-stage origin of the Coast Range ophiolite, California: implications for the life cycle of supra-subduction zone ophiolites. International Geology Review, 46, 289-315. SIDDOWAY, C.S., SASS III, L.C. & ESSER, R.P. 2005. Kinematic history of the Marie Byrd Land terrane, West Antarctica: direct evidence from Cretaceous mafic dykes. In: VAUGHAN, A.P.M., LEAT, PT. & PANKHURST, RJ. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 417-438. SlLBERLING, N.J., NlCHOLS, K.M., BRADSHAW, J.D. &

BLOME, CD. 1988. Limestone and chert in tectonic blocks from the Esk Head subterrane, South Island, New Zealand. Geological Society of America Bulletin, 100,1213-1223. SIVELL, WJ. & McCuLLOCH, M.T. 2000. Reassessment of the origin of the Dun Mountain Ophiolite, New Zealand: Nd-isotopic and geochemical evolution of magma suites. New Zealand Journal of Geology and Geophysics, 43, 133-146. SMITH, PL., TIPPER, H.W. & HAM, D.M. 2001. Lower Jurassic Amaltheidae (Ammonitina) in North America: paleobiogeography and tectonic implications. Canadian Journal of Earth Sciences, 38, 1439-1449. SMITHIES, R.H. 2000. The Archaean tonalite-trondhjemite-granodiorite (TTG) series is not an analogue of Cenozoic adakite. Earth and Planetary Science Letters, 182, 115-125. Son, W., NAKAYAMA, K. & KIMURA, T. 1998. Arc-arc collision in the Izu collision zone, central Japan, deduced from the Ashigara Basin and adjacent Tanzawa Mountains. Island Arc, 7, 330-341. SONG, X.Y., ZHOU, M.F., CAO, Z.M. & ROBINSON, PT. 2004. Late Permian rifting of the South China Craton caused by the Emeishan mantle plume? Journal of the Geological Society, London, 161, 773-781. SPANDLER, G, HERMANN, J., ARCULUS, R. & MAVROGENES, J. 2004. Geochemical heterogeneity and element mobility in deeply subducted oceanic crust: insights from high-pressure mafic rocks from New Caledonia. Chemical Geology, 206, 21-42. STEVENS, C.H., DAVYDOV, V.I. & BRADLEY, D. 1997. Permian Tethyan Fusulinina from the Kenai Peninsula, Alaska. Journal of Paleontology, 71, 985-994. STONE, P. & THOMSON, M.R.A. 2005. Archaeocyathan limestone blocks of likely Antarctic origin in Gondwanan tillite from the Falkland Islands. In: VAUGHAN, A.P.M., LEAT, PT. & PANKHURST, RJ. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 347-358. STUMP, E. 1995. The Ross Orogen of the Transantarctic

Mountains. Cambridge University Press, Cambridge. STUMP, E., WHITE, A.J.R. & BORG, S.G. 1986. Reconstruction of Australia and Antarctica: evidence from granites and recent mapping. Earth and Planetary Science Letters, 79, 348-360. SUTHERLAND, R. 1999. Basement geology and tectonic development of the greater New Zealand region: an interpretation from regional magnetic data. Tectonophysics, 308, 341-362. SUYEHIRO, K., TAKAHASHI, N., ARIIE, Y. ET AL. 1996. Continental crust, crustal underplating, and lowQ upper mantle beneath an oceanic island arc. Science, 272, 390-392. TAGAMI, T. & HASEBE, N. 1999. Cordilleran-type orogeny and episodic growth of continents: insights from the circum-Pacific continental margins. Island Arc, 8, 206-217. TAKEMURA, A., AITA, Y, HORI, R.S. ET AL. 2002. Triassic radiolarians from the ocean-floor sequence of the Waipapa terrane at Arrow Rocks, Northland, New Zealand. New Zealand Journal of Geology and Geophysics, 45, 289-296. TESSENSOHN, F. & HENJES-KUNST, F 2005. Northern Victoria Land terranes, Antarctica: far-travelled or local products? In: VAUGHAN, A.P.M., LEAT, PT. & PANKHURST, RJ. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 275-292. THOMAS,W.A. & ASTINI, R.A. 2003. Ordovician accretion of the Argentine Precordillera terrane to Gondwana: a review. Journal of South American Earth Sciences, 16, 67-79. THOMAS, W.A., ASTINI, R.A. & DENISON, R.E. 2001. Strontium isotopes, age, and tectonic setting of Cambrian salinas along the rift and transform margins of the Argentine Precordillera and southern Laurentia. Journal of Geology, 109, 231-246. THOMAS, W.A., ASTINI, R.A. & BAYONA, G 2002. Ordovician collision of the Argentine Precordillera with Gondwana, independent of Laurentian Taconic orogeny. Tectonophysics, 345, 131-152. THOMAS, W.A., ASTINI, R.A., MULLER, PA., GEHRELS, G.E. & WOODEN, J.L. 2004. Transfer of the Argentine Precordillera terrane from Laurentia: constraints from detrital-zircon geochronology. Geology, 32, 965-968. TROP, J.M., RIDGWAY, K.D., MANUSZAK, J.D. & LAYER, P. 2002. Mesozoic sedimentary basin development on the allochthonous Wrangellia composite terrane, Wrangell Mountains basin, Alaska: a long-term record of terrane migration and arc construction. Geological Society of America Bulletin, 114, 693-717. Twiss, RJ. & MOORES, E.M. 1992. Structural Geology. Freeman & Co, New York. UMHOEFER, PJ. & DORSEY, RJ. 1997. Translation of terranes: lessons from central Baja California, Mexico. Geology, 25, 1007-1010. UNRUG, R. 1992. The supercontinent cycle and Gondwanaland assembly: component cratons and the timing of suturing events. Journal of Geodynamics, 16, 215-240.

TERRANE PROCESSES VAUGHAN, A.RM. 1995. Circum-Pacific midCretaceous deformation and uplift: a superplume-related event? Geology, 23, 491-494. VAUGHAN, A.P.M. & JOHNSTON, ID. 1992. Structural constraints on closure geometry across the lapetus Suture in eastern Ireland. Journal of the Geological Society, London, 149, 65-74. VAUGHAN, A.P.M. & LIVERMORE, R.A. 2005. Episodicity of Mesozoic terrane accretion along the Pacific margin of Gondwana: implications for superplume-plate interactions. In: VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, RJ. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 143-178. VAUGHAN, A.P.M. & STOREY, B.C. 2000. The eastern Palmer Land shear zone: a new terrane accretion model for the Mesozoic development of the Antarctic Peninsula. Journal of the Geological Society, London, 157, 1243-1256. VAUGHAN, A.P.M., KELLEY, S.P. & STOREY, B.C. 2002. Mid-Cretaceous ductile deformation on the Eastern Palmer Land Shear Zone, Antarctica, and implications for timing of Mesozoic terrane collision. Geological Magazine, 139, 465-471. VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, R.J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246. WANDRES, A.M. & BRADSHAW, ID. 2005. New Zealand tectonostratigraphy and implications from conglomeratic rocks for the configuration of the SW Pacific of Gondwana. In: VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, R.J. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 179-218. WANDRES, A.M., BRADSHAW, ID., WEAVER, S., MAAS, R., IRELAND, T. & EBY, N. 2004. Provenance analysis using conglomerate clast lithologies: a case study from the Pahau terrane of New Zealand. Sedimentary Geology, 167, 57-89. WANG, K.L., CHUNG, S.L., O'REILLY, S.Y., SUN, S.S., SHINJO, R. & CHEN, C.H. 2004. Geochemical constraints for the genesis of post-collisional magmatism and the geodynamic evolution of the northern Taiwan region. Journal of Petrology, 45, 975-1011. WANG, Z.H. & Lu, H.F. 1997. 40Ar/39Ar geochronology and exhumation of mylonitized metamorphic complex in Changle-Nanao ductile shear zone. Science in China, Series D-Earth Sciences, 40, 641-647. WEAVER, S.D., BRADSHAW, ID. & LAIRD, M.G. 1984. Geochemistry of Cambrian volcanics of the Bowers Supergroup and implications for the early Paleozoic tectonic evolution of Northern Victoria Land, Antarctica. Earth and Planetary Science Letters, 68, 128-140. WEBER, B. & KOHLER, H. 1999. Sm-Nd, Rb-Sr and U-Pb geochronology of a Grenville terrane in southern Mexico: origin and geologic history of the Guichicovi complex. Precambrian Research, 96, 245-262.

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WELLMAN, H.W 1955. New Zealand Quaternary tectonics. Geologisches Rundschau, 43, 248-257. WILLAN, R.C.R. 2003. Provenance of TriassicCretaceous sandstones in the Antarctic Peninsula: implications for terrane models during Gondwana breakup. Journal of Sedimentary Research, 73, 1062-1077. WILLIAMS, H. & HATCHER, R.D 1982. Suspect terranes and accretionary history of the Appalachian orogen. Geology, 10, 530-536. WILLIAMS, H. & HATCHER, R.D. 1983. Appalachian suspect terranes. In: HATCHER, R.D., WILLIAMS, H. & ZEITZ, I. (eds) Contributions to the Tectonics and Geophysics of Mountain Chains. Geological Society of America Memoirs, 158, 33-53. WILLNER, A.P., HERVE, E, THOMSON, S.N. & MASSONNE, H.I 2004. Converging P-T paths of Mesozoic HP-LT metamorphic units (Diego de Almagro Island, Southern Chile): evidence for juxtaposition during late shortening of an active continental margin. Mineralogy and Petrology, 81, 43-84. WILSON, IT. 1966. Did the Atlantic close and then reopen? Nature, 211, 676-681. WILSON, R.A., BURDEN, E.T., BERTRAND, R., ASSELIN, E. & MCCRACKEN, A.D. 2004. Stratigraphy and tectono-sedimentary evolution of the Late Ordovician to Middle Devonian Gaspe Belt in northern New Brunswick: evidence from the Restigouche area. Canadian Journal of Earth Sciences, 41, 527-551. XIAO, W, WINDLEY, B.F., BADARCH, G, SUN, S., Li, I, QIN, K. & WANG, Z. 2004. Palaeozoic accretionary and convergent tectonics of the southern Altaids: implications for the growth of Central Asia. Journal of the Geological Society, London, 161, 339-342. YOUNG, G.M., LONG, D.G.E, FEDO, CM. & NESBITT, H.W. 2001. Paleoproterozoic Huronian basin: product of a Wilson cycle punctuated by glaciations and a meteorite impact. Sedimentary Geology, 141, 233-254. Yu, J.H., Xu, X.S., O'REILLY, S.Y, GRIFFIN, WL. & ZHANG, M. 2003. Granulite xenoliths from Cenozoic basalts in SE China provide geochemical fingerprints to distinguish lower crust terranes from the North and South China tectonic blocks. Lithos, 67, 77-102. YUMUL, G.P. 2003. The Cretaceous southeast Bohol ophiolite complex, central Philippines: a highly disaggregated supra-subduction zone ophiolite. Journal of Asian Earth Sciences, 21, 957-965. ZIEGLER, PA. 1982. Geological Atlas of Western and Central Europe. Shell International Petroleum Mijnbouw B. V., The Hague. ZIMMERMANN, U. 2005. Provenance studies of very low- to low-grade metasedimentary rocks of the Puncoviscana complex, northwest Argentina. In: VAUGHAN, A.P.M., LEAT, P.T & PANKHURST, R. I (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 381-416. ZOLNAI, G. 1986. North American aulacogens. Bulletin de la Societe Geologique de France, 2, 809-818.

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The Tasmanides of eastern Australia R. A. GLEN Geological Survey of New South Wales, Department of Primary Industries, PO Box 344, Hunter Region Mail Centre, New South Wales, 2310 Australia (e-mail: [email protected]) Abstract: The Tasmanides of eastern Australia record the break-up of Rodinia, followed by the growth of erogenic belts along the eastern margin of Gondwana. Spatially, the Tasmanides comprise five erogenic belts, with an internal Permian-Triassic rift-foreland basin system. Temporally, the Tasmanides comprise three (super)cycles, each encompassing relatively long periods of sedimentation and igneous activity, terminated by short deformational events. The Neoproterozoic-earliest Ordovician Delamerian cycle began by rifting, followed by convergent margin tectonism and accretion of island-arc forearc crust and ?island arcs in the Middle-Late Cambrian. The Ordovician-Carboniferous convergent margin Lachlan supercycle consists of three separate cycles, each ending in major deformation. The Ordovician Benambran cycle includes convergent (island-arc) and transform margin activity terminated by terrane accretion in the latest Ordovician-earliest Silurian. The Silurian-Middle Devonian Tabberabberan cycle reflects development of a large back-arc basin system, marked by rift basins and granite batholiths, behind intra-oceanic arcs and an Ordovician-Early Devonian terrane that were accreted in the Middle Devonian. The Middle Devonian to Carboniferous Kanimblan cycle began by rifting, followed by continental sedimentation inboard of a major convergent margin system that forms the early part of the Late Devonian-Triassic Hunter-Bowen supercycle. This supercycle comprises a Late Devonian-Carboniferous continental arc, forearc basin and outboard accreted terranes and subduction complexes intruded by the roots of a Permian-Triassic continental margin arc. Complex deformation ended with accretion of an intra-oceanic arc in the Early Triassic. Key features of the Tasmanides are: continuity of cycles across and along its length, precluding growth by simple eastwards accretion; development of a segmented plate margin in the Late Cambrian, reflected by major rollback of the proto-Pacific plate opposite the southern part of the Tasmanides; rifting of parts of the Delamerian margin oceanwards, to form substrate to outboard parts of the Tasmanides; the presence of five major Ordovician terranes in the Lachlan Orogen; and the generation of deformations either by the accretion of arcs, the largely orogenparallel 'transpressive' accretion of Ordovician turbidite terranes (in the Lachlan Orogen), or by changes in plate coupling.

The Tasmanides of eastern Australian represent one sector of the Pacific margin of Gondwana that stretched 20 000 km through New Zealand, Antarctica (North Victoria Land, Transantarctic Mountains, Antarctic Peninsula) and into South America. This paper reviews that history and

discusses tectonic interpretations of the

Tasmanides in three parts: Part 1 introduces the concept of the Tasmanides; Part 2 presents an up-to-date synthesis of its development; Part 3 focuses on some key issues and processes distilled from that synthesis. This paper draws on published, or in press, papers on specific aspects of the Tasmanides, as well as building on previous syntheses in Coney (1992), Scheibner & Veevers (2000), and Betts etal (2002), and of Scheibner & Basden (1998) for New South Wales, VandenBerg et al (2000) for Victoria, Gray & Foster (1997) largely for Victoria,

Seymour & Calver (1995) for Tasmania, Bain & Draper (1997) for North Queensland and Crawford et al (20030, b) for the Cambrian of Tasmania and Victoria.

Part

1: The Tasmanides - definitions

Tasmanides is the name given to a collection of orogenic belts that records the break-up of a Mesoproterozoic supercontinent, the formation of a passive margin in the Late Neoproterozoic and the establishment of a series of convergent margin orogenic belts along part of east Gondwana from the Middle Cambrian, through collision of Gondwana with Laurussia to form Pangaea c. 320-330 Ma ago (Veevers 20006), until the beginning of Gondwana-Pangaea break-up, around 227 Ma.

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 23-96. 0305-8719/$15.00 © The Geological Society of London 2005.

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THE TASMANIDES OF EASTERN AUSTRALIA

Extent of the Tasmanides Where and what are the boundaries of the Tasmanides? The eastern boundary was rifted off by opening of the Tasman Sea, beginning at c. 90 Ma, and of the Coral Sea to the north at c. 61 Ma (Sdrolias et al 2003). As a result, fragments of the Tasmanides are preserved in parts of New Zealand, New Caledonia and, presumably, the Lord Howe Rise (Sutherland 1999). The northern boundary lies in eastern and northern Papua New Guinea (Hill & Hall 2003). The western boundary is more difficult to define and, here, it is suggested that it is a structural boundary that represents both Proterozoic rifting and, thus, major thinning of Precambrian continental crust from a normal thickness of 38-41 km (Collins et al 2003), as well as later extensional/reverse/strike-slip reactivation. The Torrens Hinge Zone along the western edge of the Adelaide Rift Complex represents a part of this margin that has undergone only minor contractional reactivation. In contrast, the Palmerville Fault System in northern Queensland represents part of the margin that has undergone major contractional reactivation in the Devonian-Carboniferous. Using this definition, it is suggested that the western edge of the Tasmanides is a zig-zag line that crosses the Australian continent from north to south (Fig. 1). In northern Queensland, the western margin of the Tasmanides coincides with the N-trending Palmerville Fault System (Fig. 1). Shaw etal (1987) showed that this fault system is a major west-dipping DevonianCarboniferous thrust system with unknown amounts of displacement. In central Queensland, the western margin is drawn along the NEtrending Diamantina River Lineament. Although this lineament truncates the southern part of the Palaeoproterozoic to Mesoproterozoic Mt Isa succession, it has no surface expression. Further south, it is argued that the western line of the Tasmanides runs along the N-S Torrens Hinge Zone that separates the Gawler craton - with thin Neoproterozoic cover - in the west, from the 4 km thick Adelaide Geosyncline (e.g. Preiss 2000, herein called the Adelaide Rift Complex) to the east. This line runs N-S, west of Mesoproterozoic rocks of the Curnamona craton, which is bounded to the east and west by the Adelaide Rift Complex. The western margin of the Tasmanides disappears offshore south of the Australian mainland, with

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Tasmania lying east of that boundary. As thus defined, the western margin of the Tasmanides separates continental crust of 'average thickness' in the west from crust in the east that underwent thinning during supercontinent break-up, mainly in the Neoproterozoic. Except in Queensland, this boundary of the Tasmanides does not coincide with the Tasman Line, described originally by Hill (1951) as the eastern boundary of Precambrian 'basement' rocks in eastern Australia. Subsequently, the concept of the Tasman Line has been broadened to represent the western boundary of the Tasmanides, with the line considered to mark the place of break-up of a Mesoproterozoic supercontinent of which Australia was part (e.g. Veevers & Powell 1984; Powell et al. 1994; Scheibner & Basden 1998; Scheibner & Veevers 2000; but not Direen & Crawford 2003ft). This paper does not subscribe to this view, as discussed in Part 3.

'Continent-ocean boundary' within the Tasmanides Outboard of the western edge of the Tasmanides, it is possible to recognize a second line that marks the approximate boundary between extended continental crust and oceanic igneous crust (Fig. 1). This line runs along the northern and western edge of the Kanmantoo Trough and under the Bancannia Trough. Extension further north is uncertain. In far north Queensland, it lies outboard of the present-day coastline, since continental crust underlies the North Queensland Orogen (Bain & Draper 1997; Hutton et al 1998). Offshore plateaux adjacent to the Queensland coast are composed of granitic intrusions into Devonian metasediments similar to that of the North Queensland Orogen (Feary et al 1993). Between the Bancannia Trough and far north Queensland, the location of the 'continent-ocean boundary' is based on the inference that probable thinned continental crust underlies most of the Thomson Orogen (Finlayson 1990). There is general agreement that the accretionary parts of the New England Orogen developed on oceanic crust. The forearc basin of that orogen formed on arcs of the underlying Lachlan Orogen that had become accreted by the Middle Devonian. In this interpretation, the Mesoproterozoic rocks of the Curnamona craton (encompassing

Fig. 1. Map of eastern Australia showing the Tasmanides, different locations of the Tasman Line (based on Direen & Crawford 2003&) and the 'continent-ocean boundary'. Geophysical image from Geoscience Australia (Milligan & Franklin 2004).

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THE TASMANIDES OF EASTERN AUSTRALIA

the Mt Painter, Olary and Broken Hill areas) are surrounded by the Adelaide Rift Complex. Rutland (1976) suggested that the Adelaide Rift Complex was a multi-branched rift, with a central unextended basement horst, the Curnamona craton. If not, then one is forced to look at core-complex type extensional models to exhume the craton in the middle of a single rift system.

Subdivision of the Tasmanides Rutland (1976) suggested that the Tasmanides constituted one of the major erogenic provinces of Australia (his Tasman Province), which was divided into a number of subprovinces: Delamerian, Lachlan and New England. Subsequent usage has been to elevate these subprovinces to province status, calling them orogenic belts (Fig. 2). Subdivisions of these orogenic belts are called subprovinces and, in some cases, these have been further divided into structural zones and blocks (e.g. in New South Wales by Owen & Wyborn (1979), Glen (1992) and Scheibner & Basden (1996) and in Victoria by Gray (1988) and VandenBerg et al (2000)). Orogenic belts that constitute the Tasmanides are now described from west to east.

The Delamerian Orogen The extent of the Delamerian Orogen is defined by the distribution of rocks that have undergone a multistage Mid-Late Cambrian to earliest Ordovician Delamerian deformation. The Delamerian Orogen thus encompasses western Tasmania, western NSW and Victoria, and eastern South Australia (Fig. 2). In western Tasmania, the Delamerian (or Tyennan) Orogen extends west from the West Tamar Fault zone of Reed et al (2002) (Fig. 2) and contains rocks that range in age from Neoproterozoic to Early Ordovician, the latter being post-tectonic to the Tyennan Orogeny, which is equivalent in time and dynamics to the Delamerian Orogeny on the mainland. On the mainland, the Delamerian Orogen extends westwards from western New South

27

Wales and Victoria into South Australia, with the northern and eastern parts obscured largely by Mesozoic and Cenozoic cover. In the Koonenberry area, the eastern boundary is taken to be the Olepoloko Fault System (Stevens 1991) (Fig. 2). In western Victoria, the boundary has been controversial and has been moved over 200 km east and west over the last ten years (see VandenBerg et al 2000). Although most authors had settled on an eastdipping thrust - the Moyston Fault (Fig. 2) (VandenBerg et al 2000) - as the Delamerian/ Lachlan boundary, 500 Ma (Delamerian) mica cooling ages in the Stawell Zone (Miller et al 2003; 2004) suggest that the boundary may be the west-dipping Avoca Fault (Fig. 2). The whole of the Stawell Zone is either in the Delamerian Orogen [assumed in this paper, as suggested by Glen (1992) and Glen et al (1992)], or in a transitional intermediate zone (Miller et al 2004). The Stawell Zone was reactivated in the Ordovician-Devonian during the Lachlan supercycle. The eastern edge of the Delamerian Orogen in NSW is uncertain, being obscured by Mesozoic and Cenozoic cover. A recent interpretation (Hallett et al 2005), suggests that the Delamerian Orogen extends northeastwards from Victoria with a cusp shape into western New South Wales and, in the north, lies partly as basement to the western part of the Lachlan Orogen (Figs 2, 6). On the mainland, the Delamerian Orogen is divided into two parts (Fig. 2). In the west is an external fold-thrust belt, called the Adelaide Fold-Thrust Belt by some, that developed from inversion of the Adelaide Rift Complex and overlying shallow-water Cambrian sediments. Further east, the internal part of the Delamerian Orogen, called the Kanmantoo Fold Belt by some workers (e.g. Scheibner 1987), consists of multiply-deformed and metamorphosed sediments (intruded by granitoids) that developed from the inversion of the deep-water Cambrian Kanmantoo Trough. These metamorphic rocks and granitic rocks extend eastwards under Cenozoic cover into western New South Wales and Victoria. They pass eastwards into two elongate (350 X 50 km) belts of volcanic rocks

Fig. 2. Subdivisions of the Tasmanides. Also shown is the I-S line separating I-Type granitoids on the east from S±I-type granitoids on the west, subprovinces (subp) within the Lachlan Orogen (eastern, central, western and southwestern), subdivision of the Delamerian and New England orogens into internal (INT) and external (EXT) parts, and subdivision of the North Queensland Orogen into subprovinces (H, Hodgkinson; B, Broken River; L, Lolworth-Ravenswood: Ba, Barnard Metamorphics). DRL, Diamantina River Lineament; AF, Avoca Fault; MF, Moyston Fault; PFS, Palmerville Fault System; OF, Olepoloko Fault; WTF, West Tamar Fault. The Lolworth-Ravenswood block also shows basement inliers and Cambro-Ordovician volcanic rocks ( A ) of the Seventy Mile Group. Polygon west of the I-S line is the Australian Capital Territory.

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concealed beneath Mesozoic and younger cover, as revealed by aeromagnetic data (Figs 1, 5a). Basement to the Delamerian Orogen on the mainland is Palaeoproterozoic and Mesoproterozoic crust brought to the surface in the hangingwalls of some thrusts south of Adelaide. These rocks are inferred to be part of the Curnamona craton, exposed in the Broken Hill-Olary region (located in Fig. 1). Granite data (Foden et al 2002) suggest that the Kanmantoo Trough (Fig. 5a) was developed directly on oceanic crust. Similarly, the presence of Proterozoic E-MORB basalts east of the Devonian Bancannia Trough in the Koonenberry area in the northeast of the orogen (Mills 2003) suggests that Cambrian turbidites there were developed on pre-break-up extensional volcanic rocks (A. Crawford pers. comm. 2004). Os isotopic data suggest that Proterozoic lithospheric mantle extends east into the Stawell Zone (Handler et al 1997). The Delamerian Orogen in Tasmania and on the mainland has undergone Carboniferous deformations, as well as Early and Middle Devonian heating, that reflect events in the neighbouring Lachlan Orogen to the east.

Lachlan Orogen The Lachlan Orogen contains mainly rocks that range in age from basal Ordovician through to Carboniferous. It does, however, contain several belts of Cambrian greenstone in central Victoria that overlap in age with the Delamerian Orogen, as well as Cambrian rocks on the south coast of New South Wales. Deformation events occur around the Ordovician-Silurian boundary (Benambran Orogeny), locally around the Silurian-Devonian boundary (Bindi deformation), late Early to Middle Devonian (Tabberabberan Orogeny) and Early Carboniferous (Kanimblan Orogeny). The orogen is divided structurally into four subprovinces (Figs 2, 6), separated by major faults. The boundaries between the Eastern and Central, and between the Central and Southwestern subprovinces are major oblique overthrusts and are called sutures because of disappearance of small amounts (hundreds of kilometres) of oceanic crust (e.g. Scheibner 1987). Although Gray and coworkers (e.g. Gray et al 1997) do not distinguish between the Southwestern and Western subprovinces, interpretation of geophysical data indicates north-to-south differences in geology (Fig. 6). A variety of possibilities has been put forward for the substrate of the Lachlan Orogen:

continental (Rutland 1976; Chappell & White 1974), oceanic (Crook (1980), mixed oceanic and continental (Scheibner 1973), and oceanic in central Victoria with subsequent major underthrusting of continental material (Crawford et al 1984). Traditional models of granite genesis suggested that west of the I-S line, the Lachlan Orogen was underlain by Precambrian continental crust (Chappell & White 1974), but results of detrital zircon dating indicate that this crust is no older than Early Ordovician (Williams & Chappell 1998). I-type granites east of the I-S line were thought to be sourced from 500-600 Ma tonalites (Chappell & Stephens 1988; Williams & Chappell 1998). Alternative models of granite genesis suggest that a mafic igneous substrate underlies Ordovician turbidites (Keay et al 1997; Collins 1998) and this is consistent with high positive epsilon Nd data from the Ordovician Macquarie Arc (Glen et al 1998; Crawford et al 2005). In the Southwestern and Central subprovinces, conformable relations between Cambrian volcanic and sedimentary rocks and overlying Ordovician turbidites are used to infer that the turbidites were deposited on oceanic igneous substrate (Crawford et al 1984; Gray & Willman 19910). Some authors have suggested that rifted-off fragments of Precambrian craton form the substrate to parts of the Lachlan Orogen. Scheibner (1989) suggested that there was a 'Molong microcontinent' under the Ordovician Macquarie Arc because of the shoshonitic affinities of some of the volcanism (based on the interpretation of Wyborn 1992). However, shoshonitic volcanism, now known to be restricted largely to the Late Ordovician in that arc (Glen et al 1998; Crawford et al 2005), is also present in modern intra-oceanic volcanic arcs such as Fiji. A continental substrate is precluded effectively by the high positive epsilon Nd values (Glen et al 1998; Crawford et al 2005) and primitive Pb isotope ratios (Carr et al 1995). Packham (1973) and Scheibner (1989) also suggested that there was Precambrian continental crust under central Victoria: the latter basing his interpretation on the view that the thin-skinned thrusts described from the western, Bendigo Zone of the Southwestern subprovince (Figs 2,6) by Cox et al (1991) were foreland-type thrusts developed on rigid basement. Although Gray & Willman (19910, b) and Gray (1995) showed that these thrusts were not of a foreland type (see also Glen & VandenBerg 1996) and could develop on oceanic igneous crust, Cayley et al (2002) resurrected the idea of a continental basement (Selywn

THE TASMANIDES OF EASTERN AUSTRALIA

Block), based on apparently continuous magnetic zones from western Tasmania towards central Victoria.

The Thomson Orogen The Thomson Orogen underlies much of central and western Queensland where it is concealed by Mesozoic cover. The western margin against obscured rocks of the Delamerian Orogen is uncertain (Murray 1994): both contain lowgrade rocks. As suggested from geophysical trends by Murray & Kirkegaard (1978) and Wellman (1995), the southern boundary of the Thomson Orogen against the Lachlan Orogen is a curvilinear east-west fault zone in northwestern New South Wales (Fig. 2), (Olepoloko Fault in the west, Louth-Eumarra Shear Zone in the east: Stevens 1985 and Glen et al 1996, respectively). This boundary truncates N-S structures of the Lachlan Orogen at a high angle, and curved and concealed magnetic volcanic units in the southern Thomson Orogen at a low angle (Fig. 2). This fault zone is inferred to be a major suture between two orogenic belts (see below). It swings northwards into Queensland where it may correspond to the Foyleview Geosuture of Finlayson et al. (1990). Further north, the eastern boundary of the Thomson Orogen extends east to the Permian-Triassic Bowen Basin on the surface. Interpretation of seismic reflection data (Korsch et al. 1997) suggests that the orogen passes beneath that basin, to underlie structurally the western part of the New England Orogen. To the north, the Thomson Orogen is bounded by the North Queensland Orogen and Meso- and Palaeoproterozoic cratonic Australia. Drill hole data indicate that rocks in the Thomson Orogen range in age from Precambrian through to Late Devonian (Murray 1994; Scheibner & Veevers 2000). However, Cambrian or older rocks deformed in the 500 Ma Delamerian Orogeny occur along the eastern margin, in the Anakie Inlier (Fig. 2). This inlier extends in the subsurface to the south, into the Nebine gravity ridge (Withnall 1995). Cores from this ridge consist of metasediments, phyllites and multiply-deformed schist at lower greenschist to ?amphibolite grade (Murray 1994). However, the ages of these rocks are unknown, the only indicator being a 416 Ma K-Ar date on biotite (Murray 1994). The south end of the Nebine ridge swings to the west and coincides with belts of major aeromagnetic and gravity highs that reflect concealed igneous rocks (proven by drilling, but

29

of unknown age) that lie along the southern boundary of the orogen (Fig. 2, see below). It is suggested here that they may be of Cambrian age, since there is some evidence (below) that the oldest deformation across the boundary between the Lachlan and Thomson orogens is Late Cambrian, probably reflecting accretion of an intra-oceanic arc. When coupled with the uncertain nature of the western boundary of the Thomson Orogen (above), this continuity allows two possible suggestions. 1.

2.

The Nebine ridge and the Anakie Inlier are physically part of an enlarged Delamerian Orogen that extended eastwards for some 600 km. Such a margin is similar to the Nebine arc inferred by Harrington (1974). If validated by future work, such an orientation implies that a major sector of the Australian plate margin was orientated east-west in the Cambrian. Older rocks along the southern and eastern margin of the Thomson Orogen are pieces rifted away from Delamerian Gondwana, with the Thomson Orogen being the site of post-500 Ma rifting. In this model, the Thomson Orogen is floored by extended and thinned continental crust (Harrington 1974; Murray 1994) that might correspond to the geophysically layered crust of Finlayson et al. (1990). Gravity data show that this crust has WNW trends in the southern third of the orogen and NE trends in the north (Milligan et al. 2003).

The New England Orogen The New England Orogen is the most easterly component of the Tasmanides. It occupies much of coastal Queensland and extends south below the Mesozoic cover of the Clarence-Moreton and Surat basins into northeastern New South Wales (southern New England Orogen) (Figs 2, 13a; Leitch 1974; Scheibner & Veevers 2000). The orogen also forms basement to the eastern part of the Sydney Basin and extends offshore as the Currarong Orogen (Jones & McDonnell 1981; Jones et al. 1984), which is represented in seismic data as the 'offshore uplift' (Bradley 1993; Alder et al 1998). The New England Orogen has an inferred thrust contact with the Eastern subprovince of the Lachlan Orogen. The New England Orogen is divided into two structural subprovinces (Leitch 1974): a western, external part that constitutes a fold-thrust belt (with intrusive granites in the north) and an internal part in which accretionary complex rocks are multiply-deformed

R. A. GLEN

30

and metamorphosed, and intruded by granitoids. This subdivision reflects the development of a Late Devonian-Carboniferous classical convergent margin consisting of arc, forearc basin and accreted terranes. The subsequent history involves Permian rifting and Triassic subduction followed by a protracted PermianTriassic (Hunter-Bowen) deformation. Substrate to the New England Orogen is inferred to be oceanic, east of the Peel-Manning Fault System in New South Wales and the Yarrol Fault System in Queensland, and mixed oceanic and continental to the west, becoming continental by the Carboniferous and undergoing rifting in the Early Permian. However, small fault-bounded outcrops of Neoproterozoic-Early Devonian subduction-related and accretionary material overlap with the development of Delamerian and Lachlan orogens and point to the presence of older substrate. This is consistent with several other lines of evidence: 1.

2. 3. 4.

presence of Neoproterozoic Re-Os ages, commonly in the range of 0.6 to. 1.8 Ga (Powell & O'Reilly 2001; Bennett et al 2002); Neoproterozoic zircon model ages for some young granites (Shaw & Flood 2002); seismically fast lithospheric mantle interpreted from SKIPPY data by Van der Hilst et al (1998); and c. 562 Ma Sm-Nd isochron age on gabbro in ophiolitic rocks of a mid-Palaeozoic accretionary complex in central Queensland (Bruce et al. 2000).

Together, they point to the persistence of old lithosphere.

Permian-Triassic Bowen-GunnedahSydney Basin system This system is remarkable for its length of c. 1600 km (Fig. 2). It originated by rifting in the Early Permian and was converted into a basin foreland in the mid-Triassic, with the eastern parts constituting the foreland fold-thrust belt of the New England Orogen (Glen & Beckett 1997). The geology has been summarized by several authors (Harrington et al 1989; Veevers et al 1994; updated in Veevers 2000d, Murray 1990; Scheibner & Basden 1998; Fielding et al 2001). The substrate of the Bowen Basin is the Thomson Orogen in the west and New England Orogen in the east. The Gunnedah Basin was

built over crust of the Lachlan Orogen (Korsch et al 2002). The substrate of the Sydney Basin is inferred to be Lachlan Orogen in the west (felsic volcaniclastic rocks have been brought up in diatremes that intrude the basin, O'Reilly 1990) and the New England Orogen in the east, with a depositional relationship preserved north of Newcastle (Roberts & Engel 1987).

The North Queensland Orogen This term has been introduced in order to group together a series of tectonic provinces (which have been downgraded to subprovince status) in far north Queensland, known as Hodgkinson Province, Broken River Province, LolworthRavenswood Block and Barnard Province (Fig. 2). These were reviewed most recently by Bain & Draper (1997). The broad Ordovician to Carboniferous evolution of the North Queensland Orogen is similar broadly to that of the Lachlan Orogen, with two important exceptions - the presence of scattered inliers of Neoproterozoic continental crust that reflect growth on Precambrian continental crust, and the presence of Cambrian to Ordovician arc-related rocks in the south that are similar to the Tasmanian part of the Delamerian Orogen. To the west, the North Queensland Orogen is fault-bounded against Meso-and Palaeoproterozoic rocks of cratonic Australia along the Palmerville Fault System, an imbricate Devonian-Carboniferous thrust system (Shaw et al 1987). The North Queensland Orogen extends offshore, as indicated by inferred Palaeozoic strata underlying the offshore Marion and Queensland plateaux (Feary et al 1993), although Henderson et al (1998) suggested that these plateaux were underlain by accretionary complex rocks of the New England Orogen. The southern boundary against the Thomson Orogen is marked by truncation of the N-trending pre-500 Ma Anakie Inlier, seen also in gravity images (Murray et al 1989; Milligan et al 2003), although similar-aged Precambrian crust occurs as inliers within the orogen. The North Queensland Orogen underwent deformation in the Ordovician, Middle Devonian and Carboniferous (Bain & Draper 1997).

Tectonic cycles The subdivision into different orogenic belts is based on the age distribution of constituent strata and the ages of 'climactic' deformation. The eastwards-younging of ages in those belts (Fig. 2) has been cited widely as evidence that the Australian craton grew by eastward accretion

THE TASMANIDES OF EASTERN AUSTRALIA from the Neoproterozoic until the plate boundary jumped to the New Caledonia-New Zealand area in the Mesozoic (Sutherland 1999). However, the increasing recognition of older rocks in parts of the so-called younger orogenic belts shows that growth of the Australian sector of Gondwana was not so simple. Cambrian rocks occur in parts of the Lachlan and New England orogens. They also occur, or are inferred, from the eastern and southern parts of the Thomson Orogen. This means that subdivision of the Tasmanides into orogenic belts based solely on age criteria is not as valid as thought previously. While it is still useful to show the different orogenic belts on a map (Fig. 2), time-space plots show old rocks extend right across the Tasmanides (Fig. 4). It is thus proposed here to describe the evolution of the Tasmanides in terms of tectonic cycles, in the South American sense, to encompass depositional/magmatic as well as contractional deformational histories of rock packages developed along plate margins. A complete idealized cycle begins with a rift phase that may pass into drift, then into convergence (if subduction can be recognized), then into deformation/ collision that terminates the cycle although a post-collisional phase may be present. In some cases there is no rift/drift phase. The term 'collision' is applied to the terminal deformation of many cycles in the belief that it reflects the accretion of an arc to the developing Gondwana landmass. It does not represent a continentcontinent collision. Acenolaza & Toselli (1981) and Ramos (1999) have described the cycle concept for the Palaeozoic of the South American margin of Gondwana. Seven out of the eight cycles there close with a deformation event. The exception is the Famatinian cycle that includes two significant deformations in the Ordovician (at 465 Ma and c. 440 Ma). The use of cycles here is similar to the concept of stages used previously in New South Wales (Scheibner 1973; Scheibner & Basden 1998) and in the New England Orogen by Korsch & Harrington (1981).

Stewart (Monash University) and the collection of deep seismic reflection profiles. Evolution of the Tasmanides is now described from the oldest cycle to the youngest, proceeding from south to north and west to east. Each cycle is described first from its 'home' orogen and then from other orogens (such as the Delamerian cycle from the Delamerian Orogen first and then from other orogens).

Delamerian cycle While best developed in the Delamerian Orogen on the mainland and in west Tasmania, the Delamerian cycle is also represented in the early histories of the Lachlan, New England and North Queensland orogens. In west Tasmania, this cycle lasted c. 300 million years, from >780 Ma until 490 Ma, although it may have begun much earlier, depending on the tectonic significance of 1270 Ma and 1100 Ma ages below. On the mainland, it lasted c. 350 million years, from 830 Ma to 480 Ma.

Extension and passive margin phase The oldest rocks in the Tasmanides occur in western Tasmania and on King Island between Tasmania and the mainland (Figs 3, 4a). They consist of: • •

basalts in the west of King Island that underwent amphibolite-grade metamorphism at c. 1270 Ma (Holm et al 2003); shallow-water undeformed sandstone in the NW corner of Tasmania with inferred c. 1100 Ma depositional ages (Turner et al. 1992).

Younger rocks that fit the rift history on the mainland better (see below) comprise: •

Part 2: Synthesis of the Tasmanides This synthesis is based largely on new datasets collected in the last ten years, in the period since the Tectonophysics issue on the Lachlan Fold Belt and related regions (Fergusson & Glen 1992). Key datasets include new aeromagnetic and gravity databases, new geological mapping, new mineral and in situ age-dating techniques, new conodont identification techniques by Ian

31

•

•

>780 Ma mafic volcanic rocks of the allochthonous Bo wry Formation that underwent blueschist metamorphism (Holm et aL 2003). The Cooee Dolerite (with a minimum K-Ar age of 725 Ma, Crook 1979) is a correlative. Together they represent an early phase of rift volcanism (Holm et al 2003); the Wickham deformation, dated by the emplacement of associated granites on King Island, at 760 Ma (Turner et al 1998). This deformation also affected NW Tasmania, where similar granites are dated at 777 Ma (Turner et al 1998); c. 700 Ma glacially-derived Sturtian and c. 635 Ma Marinoan conglomerates in

32

R. A. GLEN

Fig. 3. Precambrian part of the Delamerian cycle. Rift cycle 1, Adelaide Rift Complex evolution, largely after Preiss (2000). Coloured polygons represent major depocentres and arrows represent inferred directions of maximum extension, (a) Willouran; (b) Torrensian (THZ, Torrens Hinge Zone); (c) Sturtian; (d) Marinoan (M, Marlborough; SDRS, seaward-dipping reflector sequences after Direen & Crawford (2003)), with inferred old subcontinental lithosphere (s.c.l.) under the New England Orogen.

younger basins in NW Tasmania (Figs 3c, d, 4) (Calver & Walters 2000); volcanic-rich rift basins, containing tholeiites passing up into second-stage melt lavas (picrites) (Crawford & Berry 1992) (Fig. 3d).

These igneous rocks formed during a major 600-580 Ma rifting event in western Tasmania and were interpreted as seawarddipping reflector sequences by (Direen & Crawford 20030) and (Crawford etal 20030).

THE TASMANIDES OF EASTERN AUSTRALIA

33

Fig. 4. Time-space plots, (a) Southern plot from Delamerian Orogen in South Australia and west Tasmania through to Lachlan Orogen in Victoria. Stratigraphic data mainly from Seymour & Calver (1995), Holm et al. (2003), Preiss (2000), VandenBerg et al. (2000) plus other sources cited in text. Time-scale - Veevers (2000e) is used as the source for most of the time-scale, with the exception of that part from the Ordovician to the Carboniferous which (rounded up or down) comes from Pogson & Percival (2003). For the Ordovician this latter scale is based on Cooper (1999). British stage names are also provided.

34

R. A. GLEN

Fig. 4. (b) Central plot through New South Wales, from Delamerian Orogen (Koonenberry) through Lachlan Orogen into New England Orogen. Stratigraphic data mainly from Mills (2002, 2003), Colquhoun et al (2004), Percival & Glen (2006), Meffre & Glen (unpublished), Thomas et al (2002), Glen et al (20040), Roberts & Geeve (1999), Aitchison et al (19920), and other sources cited in text.

THE TASMANIDES OF EASTERN AUSTRALIA

35

Fig. 4. (c) Northern plot through Queensland. Stratigraphic data mainly from Bain & Draper (1997), Bultitude et al (1993), Withnall et al (1996), Henderson et al. (1998), Leitch et al. (2003), the Yarrol Project Team (1997) and other sources cited in text.

36

R. A. GLEN

On the mainland, there is no record of rifting before 830 Ma. The rift phase of the Delamerian cycle on the mainland is divided into two rift cycles. Rift cycle 1 is represented by growth of the Adelaide Rift Complex. Rift cycle 2 is represented by formation of the Kanmantoo Trough in eastern South Australia and equivalents in western Victoria and in the Koonenberry Belt of western New South Wales (Figs 1, 3,5). The Adelaide Rift Complex lies mainly west, but also east, of the Curnamona craton (Cooper & Tuckwell 1971). Its development records supercontinent break-up in the Neoproterozoic, beginning at 827 Ma and continuing to the base of the Cambrian. Five rifting events within this cycle have been identified by Preiss (2000) (Figs 3, 4a). In contrast to west Tasmania, rifting is essentially non-volcanic, although events 1 and 3 began with volcanism. U-Pb constraints on the ages of rifting are provided in the first event by the Wooltana Volcanics (c. 827 Ma, see Preiss 2000) coeval with the Gairdner Dyke Swarm to the west dated at c. 827 Ma by Wingate et al (1998). Event 2 began with the Rook Tuff dated at c. 802 Ma (Fanning et al 1986). Event 3 began with the rhyolites of the Boucaut Volcanics dated at c. Ill Ma (C. M. Fanning 1994, quoted by Preiss 2000). Event 4 began with Sturtian glacials, dated at c. 700 Ma and event 5 in the middle Marinoan at c. 650 Ma, below the Marinoan glacials dated at 635 Ma (Preiss 2000). Deeper-water equivalents (Kara beds) occur east of the Bancannia Trough in the Koonenberry Belt (Mills 1992, 2003). A key feature of event 5 is the presence of alkaline rift volcanic rocks extruded over a large area from the Koonenberry Belt (where they are dated at c. 586 Ma, Crawford et al 1997), southwards to the Truro alkaline volcanic rocks east of Adelaide, correlated by Crawford et al (1997), and into the alkaline volcanic rocks of King Island and northwestern Tasmania. Also included in this event is the eastern 350 X 50 km belt of volcanic rocks in western New South Wales and Victoria concealed beneath Tertiary cover and revealed by aeromagnetic data (Figs 1, 5a). This belt, the Dimboola Igneous Complex of VandenBerg et al (2000), contains ultramafic and mafic tholeiites, boninites, volcaniclastics and cumulate gabbros. They are interpreted as a set of c. 600 Ma seaward-dipping reflector sequences by Direen & Crawford (20030). VandenBerg et al (2000), however, suggested they represent a Cambrian intra-oceanic arc (see below). Major tectonic changes occurred in the

Middle Cambrian on the mainland (Figs 4a, 5) with the onset of rift cycle 2, which is absent in western Tasmania. On the mainland, a shallowwater Early Cambrian limestone shelf (part of the Stansbury Basin, Ardrossan Shelf in the west and Arrowie Basin in the north, e.g. Belperio et al 1998; Preiss 2000) was developed above a regional hiatus on top of the western and northern parts of the Adelaide Rift Complex (Figs 4a, 5a). A tuff band at the top of the shelf sequence has a U-Pb date of c. 526 Ma (Cooper et al 1992) (Fig. 4a). This shelf was followed by a major rifting event in the east (Waitpingan Subsidence, Thomson 1969; Haines & Flottmann 1998) that led to the formation of a major new deep-water rift basin - the Kanmantoo Trough - that extended eastwards into western Victoria and has analogues in the Ponto beds of the Koonenberry Belt (Mills 2003). The Kanmantoo Trough was filled by rapid deposition of c. 7-8 km thick (Jago et al 2003), high-density turbidites (Haines et al 2001) from 526 Ma till the onset of deformation around 514 Ma (Foden et al 2002) (Fig. 4a). The Kanmantoo Trough formed as a transtensional basin in response to NE-SW extension, with W-E palaeocurrents reflecting deflection by bounding faults (Flottmann et al 1998; Haines & Flottmann 1998). The presence of basalts and gabbros, dated at c. 524 Ma and with within-plate and MORB chemistry (Rankin et al 1991; VandenBerg et al 2000), suggest that rifting had progressed to the formation of Cambrian oceanic crust. The Kanmantoo Group was derived from Antarctica (Flottmann et al 1998) and shares the 600-500 Ma zircon population (the Pacific Gondwana zircons of Ireland et al 1994) with Ordovician turbidites of the Lachlan Orogen, which are derived from the Ross Orogen of Antarctica (but see Williams et al 2002). This source contrasts with derivation of the Adelaide Rift Complex (Turner et al 1996; Ireland et al 1998) and the partly coeval Early Ardrossan Shelf (Ireland et al 1998), largely from the craton to the west (Veevers 20000). In the New England Orogen, the rift phase of the Delamerian cycle is represented by an ophiolite in the Marlborough Block of central Queensland that has a 562 ± 23 Ma Sm-Nd isochron age (Bruce et al 2000) (Figs 3d, 4c). The ophiolite has depleted MORB-like trace element characteristics that suggest formation as oceanic crust at a Neoproterozoic ocean ridge (Bruce et al 2000). These data are consistent with the existence of a proto-Pacific Ocean east

THE TASMANIDES OF EASTERN AUSTRALIA

37

of the Delamerian Orogen after supercontinent break-up. In the western part of the Thomson Orogen, the Delamerian rift cycle is represented by continent-derived sandstones and mudstones with 580-480 Ma K-Ar deformation ages (Murray 1994; Scheibner & Veevers 2000). Correlation with rocks of rift cycles 1 or 2 to the southwest is uncertain. In the eastern part of the orogen, alkaline and tholeiitic volcanic rocks in the Anakie Inlier (Fig. 2) provide further evidence of rifting, while the presence of quartz sandstone suggests an intracontinental setting (Withnall 1995). These rocks may be as old as Neoproterozoic (Withnall et al 1996), or late Neoproterozoic to Middle Cambrian (Fergusson et al 2001). The North Queensland Orogen includes inliers of metamorphic complexes, which appear to represent exhumed 1100-1200 Ma ('Grenvillian') continental crust (Blewett et al 1998) as well as Neoproterozoic rift-related sediments and volcanic rocks indicative of an intraplate extensional origin (Draper et al 1998; Hutton et al 1998) (Figs 2, 4c). These inliers occur in both the Broken River and LolworthAnakie subprovinces, while the metamorphic rocks in the Barnard subprovince predate the intrusion of Ordovician granites, the oldest of

Fig. 5. Delamerian cycle, rift cycle 2, convergent, collisional and post-collisional phases, (a) Blue unit is c. 522-514 Ma Kanmantoo Trough with mafic volcanic rocks ( A ). Pale green units are deformed volcanic packets. Eastern volcanics = Dimboola Igneous Complex - either a Cambrian arc or a 600-580 Ma seaward-dipping reflector sequence (see Fig. 3d). Subduction zones shown in red with red barbs. East-dipping subduction zone based on model of Crawford & Berry (1992). Possible north- and west-dipping subduction zones postulated in this paper based on the presence of inferred Cambrian volcanic rocks. Eclogite at Attunga (A) in northern NSW and 530 Ma ophiolites along the Peel-Manning Fault System (pmfs) are also part of the convergent phase, (b) Collisional phase, showing regional trends in black wrapping around Precambrian continental buttresses shown in yellow. Arrows show directions of maximum shortening. Fleurieu structural arc highlighted, (c) Syn- to post-collisional phase showing syn- and post-orogenic granites and post-collisional volcanic rocks and sedimentary rocks. Dark blue area represents post-collisional turbidites; light blue areas represent post-collisional shallow-water deposits. Abbreviations: A, Attunga; dt, Dundas Trough; g, Gidgealpa; k, Koonenberry; mrv, Mt Read Volcanics; p, Padthaway; pmfs, Peel-Manning Fault System; wb, Waratah Bay.

38

R. A. GLEN

which is 486 Ma (Bultitude & Garrad 1997) (Fig. 2). These inliers may be correlatives of the Anakie Inlier (Withnall 1995), since they predate Delamerian deformation and cooling (Rb-Sr isochron at c. 500 Ma from the Cape River Metamorphics in the Broken River Subprovince, Draper et al 1998) (Fig. 5a). In the Cape River Metamorphics, amphibolite-grade sandstones with c. 1145 Ma dominant detrital zircons are intruded by Late Cambrian-Early Ordovician granite and overlain by Late Cambrian-Early Ordovician volcanic rocks (see below) (Hutton et al. 1998). In the LolworthRavensworth Block, the Charters Towers Metamorphics contain c. 507 Ma mafic volcanic rocks (Bain & Draper 1997). The presence of a Grenvillian basement is supported further by Neoproterozoic detrital zircons in Palaeozoic strata, old zircons in early Palaeozoic granites and, significantly, negative epsilon Nd values (Bain & Draper 1997).

Convergent phase The convergent phase of the Delamerian cycle is represented by development of crust formed in the forearc of an intra-oceanic island arc. Convergence between the extended west Tasmania craton and the proto-Pacific plate is reflected by the development of Cambrian mafic-ultramafic complexes. These constitute relics of forearc igneous crust that developed on the proto-Pacific plate above an eastdipping subduction zone and west of an intraoceanic arc (Crawford & Berry 1992). Reed et al (2002) suggested that the east-dipping seismic reflectors of Barton (1999) in the west Tamar region reflect the location of the Cambrian Delamerian subduction zone (Fig. 5a). A zircon date c. 510 Ma from the Heazlewood Complex (Turner et al. 1998) reflects ongoing subduction and dates convergence from c. 520 to 510 Ma (A. Crawford pers. comm. 2004) (Fig. 4). Candidate arc-rocks on the mainland are concealed beneath Mesozoic and younger cover, but are revealed by geophysical data. Two NSW-NNE-trending belts, an eastern and western, each c. 350 X 50 km, occur in western New South Wales, Victoria and eastern South Australia. The third, a curved east-west belt, lies along the southern margin of the Thomson Orogen (Figs 1, 5a). The eastern belt is the Dimboola Igneous Complex of VandenBerg et al. (2000), interpreted as 580-600 Ma rift volcanic rocks by Direen & Crawford (20030) but as a Cambrian intra-oceanic arc reflecting convergence along

the Gondwana proto-Pacific plate boundary by VandenBerg et al. (2000). Elements of rift, forearc crust and post-collisional volcanic rocks may be present since the small Stavely Volcanic Complex just to the east and south contains slices of serpentinized boninitic ultramafics, interpreted as remnants of the forearc of an intra-oceanic island arc, overlain by postcollisional felsic volcanic rocks (Crawford et al. 1996) (Fig. 4a). The tholeiitic Magdala Volcanics in the Stawell Zone lie immediately east of the Dimboola Igneous Complex and the Moyston Fault. Although equivalent in age, they lack the strong geophysical response on a regional scale. Crawford et al. (2003b) quoted a mixing age of 518 Ma. The western volcanic belt in the Delamerian Orogen (Figs 1, 5 a) wraps around the Precambrian Curnamona craton and contains a mixture of back-arc, arc-like and ?post-collisional volcanic rocks with zircon ages of 521-480 Ma. The Padthaway area (Fig. 1) contains MORB volcanic rocks (Rankin et al. 1991) and felsic volcanic rocks intruded by syn- to postdeformation mafic and granitic intrusive rocks. One felsic volcanic rock has been U-Pb dated at c. 493 Ma, and one granite at c. 480 Ma (Fanning 1996). Gabbro (just east of the P, Fig. 1) has been U-Pb dated at 525 Ma (Maher et al. 1977). Further north, in the ENE part of the belt, Neoproterozoic and Cambrian mafic volcanic and sedimentary rocks are intruded by Cambrian-Ordovician, late- to post-deformation granitoids and mafic to intermediate igneous rocks such as microdiorites (U-Pb dated at c. 482 Ma, Fabris 2003). Just ESE of Broken Hill, K-rich volcanic to subvolcanic rocks are intruded by granite, diorite and monzodiorite U-Pb dated at 521 Ma, 519 Ma and 505 Ma (cited by Mills 2001). The Koonenberry area near the northern part of this western belt (Fig. 5a) contains calc-alkaline andesitic volcanic rocks overlain by late Early Cambrian felsic volcanic rocks U-Pb dated at 525 Ma (J. Claoue-Long 1992, cited in Zhou & Whitford 1994) below Middle Cambrian limestone, shale and volcanolithic conglomerate. These shallowwater rocks pass eastwards into fault-bounded turbidites of the Teltawongee beds and overlying Ponto beds that are equivalent to the Kanmantoo Group (Mills 1992). Concealed Early Cambrian volcanic rocks at Gidgealpa at the northern tip of the western belt consist of a trachytic lower part and an upper part of rhyolitic to dacitic tuff beds with some andesites (Gatehouse 1986; Gravestock & Gatehouse 1995) (Fig. 5a).

THE TASMANIDES OF EASTERN AUSTRALIA The calc-alkaline volcanic rocks at Koonenberry were interpreted as an arc by Scheibner (1987) and Scheibner & Basden (1998), although Zhou & Whitford (1994) and Crawford et al. (1997) suggested that they represented rift volcanism from their geochemistry, their isotope contents and, in part, also from the small amount of volcanic material. A rift setting was also ascribed to the Gidgealpa Volcanics by Gravestock & Gatehouse (1995). However, Sharp & Buckley (2003) suggested that the Koonenberry rocks are subductionrelated and included with them 520 Ma volcanic rocks intersected in drill holes beneath the Bancannia Trough just to the west. In the Lachlan Orogen, elements of Delamerian convergence are represented by narrow zones of Cambrian mafic and ultramafic rocks in the Southwestern subprovince (Figs 5a, 6). These Cambrian igneous rocks are exposed mainly as fault slices in the hangingwalls of major thrust faults (Fig. 6): Heathcote Fault zone (the western edge of the Melbourne Trough), Mt Wellington Fault zone (the eastern edge of the Melbourne Trough) and the Governor Fault zone (the western edge of the Tabberabbera Zone). They also occur as very small bodies within the Bendigo and Melbourne structural zones (Ceres, Phillip Island, Waratah Bay - only the last is shown in Fig. 5). Most of these igneous rocks are boninitic andesitic lavas and ultramafic equivalents of forearc affinity, as well as tholeiitic basalt, the latter with back-arc basin geochemical signatures that may reflect rifting of the forearc (Crawford et al 1984; Crawford et al. 20035). A gabbro (at Dookie) at the northern end of the Governor Fault zone was dated at 502 Ma by Spaggiari etal (20030). In contrast, calc-alkaline andesites occur along, and in windows below, the Mt Wellington Fault zone (renamed the Mt Useful Fault zone by VandenBerg et al 2000). These andesites pass up into volcaniclastic rocks and are overlain by sandstones and cherts. One body (at Licola in the south) has been dated by U-Pb at 500 Ma (Spaggiari et al 20030). The third volcanic belt, largely concealed, lies in the southern margin of the Thomson Orogen and may also contain Cambrian volcanic rocks of the convergent phase of the Delamerian cycle. The presence of a significant gravity high and several magnetic ridges suggests that this margin of the orogen contains curvilinear east-west-trending major igneous bodies (Fig. 5a). Drill hole intersections (in the southern high near longitude 145°30' S (Fig. 6) indicate the presence of gabbro, andesitic volcanic and volcaniclastic rocks, amphibole peridotites and

39

sediments (Savage 2000), but there are few data as to the nature or ages of these hidden bodies. Scheibner & Basden (1998) inferred a Devonian rift, but the presence of Late Ordovician graptolites (summarized in Glen et al 1996) and Cambrian to Early Ordovician zircons in the Easter Monday beds north of Koonenberry (Stevens 1991; Stevens & Fanning unpublished) indicates that they are older. If they predate Delamerian north-south shortening in the Koonenberry Belt (Mills 2003), they could represent a mixture of Cambrian rift and convergent margin igneous rocks and ultramafics along the southern margin of the Thomson Orogen. In the southern New England Orogen, the convergent part of the Delamerian cycle is represented by eclogite blocks, exhumed in a major serpentinite-lubricated fault at Attunga on the Peel-Manning Fault System (Fig. 5a). Originally dated at 571 Ma by Watanabe et al (1998), this date has now been revised to 536 Ma by Fanning et al (2002) and is similar to the 530 Ma zircon ages obtained from plagiogranites and metadiorites in schistose serpentinite along the same fault system (Figs 4b, 5a) (Aitchison et al 19925; Sano et al 2004 respectively). The eclogite indicates Cambrian subduction. The plagiogranite ages suggest that the enclosing ophiolitic low-Ti tholeiitic basalts and boninitic ultramafic rocks are relics of a Cambrian suprasubduction zone forearc and, thus, of a Cambrian convergent plate boundary (Aitchison et al 1994). Chemically, the basalts resemble Cambrian basalts of western Tasmania (Aitchison & Ireland 1995). Convergence is also indicated by Middle to early Late Cambrian volcaniclastic rocks of the Murrawong Creek Formation that occur immediately west of the Peel-Manning Fault System (Figs 4b, 5a), dated from fossils in limestone clasts (Cawood 1976; Cawood & Leitch 1985). Although Aitchison & Flood (1990) suggested that Cambrian fossils came from allochthonous blocks in Devonian matrix, the conformably overlying Pipeclay Argillite contains Middle to early Late Cambrian conodonts (Stewart 1995). Cawood & Leitch (1985) suggested that volcanic clasts in the conglomerate were derived from a low-K intra-oceanic island arc.

Collisional phase Western Tasmania records the collision of the forearc of an intra-oceanic arc with extended East Gondwana crust around 510-505 Ma. This resulted in the southwest transport of allochthonous thrust sheets of forearc mafic and

R. A. GLEN

40

0

THE TASMANIDES OF EASTERN AUSTRALIA

ultramafic rocks over the thinned passive margin sequence and its cover of late Neoproterozoic rift basins (Berry & Crawford 1988; Crawford & Berry 1992; Berry 1994; 1995) (Figs 5a, 8). Ultramafic detritus in the middle Middle Cambrian basal units of the Dundas Trough indicates that obduction (Fig. 5b) had occurred between latest Early Cambrian and middle Middle Cambrian. Obduction was accompanied by the formation of metamorphic complexes beneath the ophiolitic sheets (Meffre et al. 2000), with one eclogite U-Pb dated at 502 Ma (Turner et al 1998). Subsequent deformation between 505 Ma and 495 Ma is subdivided into an early phase of N-S compression in the late Middle-early Late Cambrian and a later phase of E-W compression in the Late Cambrian (Berry 1994; Turner et al 1998). The high-strain Arthur Lineament running NE across western Tasmania is interpreted by Holm & Berry (2002) as a series of thrust sheets emplaced N-S in the first event and refolded and steepened by subsequent E-W folding and faulting. Woodward et al (1993) showed that allochthonous Precambrian massifs were also emplaced in the Late Cambrian. On the mainland, the Delamerian Orogeny produced an orogenic belt between 300 km and 600 km wide (Fig. 5b). The western external part of this orogen is marked by a 100-300 km wide fold-thrust belt in the external part of the orogen in the west, and an extensive, internal high-T low-P zone extending eastwards under the Cenozoic Murray Basin into western Victoria. This zone is marked by multiple deformation and metamorphism and the emplacement of syn- and post-kinematic granites. Low-grade (dominantly volcanic) rocks also occur along the eastern margin of the orogen in western Victoria, suggesting that the Delamerian Orogen may have lower-grade margins flanking a higher-grade core. The western external fold-thrust belt has variable geometry. The central and northern parts were deformed by south-vergent thrusting that is thin skinned, except where evaporites of

41

Fig. 7. Benambran cycle showing craton-derived turbidite, arc and Narooma terranes in Lachlan Orogen, blueschist ages in New England Orogen (from text) and inferred tectonic elements in northern New England Orogen and in North Queensland Orogen. Platform deposits and uplift in Delamerian orogen to west also shown. P, Port Macquarie; a, location of Anakie Inlier.

Fig. 6. Map of Lachlan Orogen showing subdivision into subprovinces, major faults and tectonostratigraphic units. Red numbers are Ar-Ar plateaux ages from Foster et al (1999). Blue numbers are deformation ages from the stratigraphic record (Glen et al 2004b). Highs and lows (L) in the southern Thomson Orogen refer to gravity features. Abbreviations: SZ, Stawell Zone; BZ, Bendigo Zone; MZ, Melbourne Zone; TZ, Tabberabbera Zone; OZ, Omeo Zone: mf, Moyston Fault; af, Avoca fault; hf, Heathcote Fault; gof, Governor Fault Zone; if, Long Plain & Indi faults; gf, Gilmore Fault Zone; tf, Tullamore Fault; ks, Kiandra-Narromine Structure; CT, Cowra Trough; CH, Capertee High; WS, Winduck Shelf; KS, Kopyje Shelf; CMH VB, Canbelego-Mineral Hill Volcanic Belt; MHT, Mt Hope Trough; WPS, Walters Range Shelf; RT, Rast Trough; MT, Melrose Trough; QB, Quidong Basin; C-Y Shelf, Canberra-Yass Shelf. Geology in Victoria based on VandenBerg et al (2000).

42

R. A. GLEN

THE TASMANIDES OF EASTERN AUSTRALIA

the Callana Group are absent (Preiss 2000) (Fig. 5b). The southern part of the fold-and-thrust belt verges westwards towards the Gawler craton and outlines the Fleurieu structural arc (Fig. 5b), which reflects variable shortening against the Gawler craton buttress (Flottmann et al 1994; Flottmann & James 1997). The thrust belt involved only very limited translation, since thrusts do not extend west of the Torrens Hinge Zone - the western edge of the Adelaide Rift Complex from Sturtian times onward. This is consistent with formation of a very restricted foreland basin that lies immediately west of that fault zone (Flottmann et al 1997). Haines & Flottmann (1998) also suggested that some Cambrian sediments east of the Torrens Hinge Zone were deposited in a foreland basin that began at c. 523 Ma, but this is much older than the onset of deformation in the internal zone at 514 Ma (Foden et al 2002). It is, however, consistent with a 531 Ma Rb-Sr age of cleavage formation in one part of the thrust belt (Turner et al 1994). In the internal part of the orogen, the Kanmantoo Trough underwent severe inversion during this Delamerian Orogeny, undergoing multiple deformation, metamorphism and granite emplacement (e.g. Jenkins & Turner 1992) (Fig. 4a). Old growth faults were reversereactivated and there was major thrusting (Flottmann & James 1997; Flottmann et al 1998). Syn-kinematic granite ages indicate that deformation commenced at 514 Ma and lasted till 490 Ma (Foden et al 2002). Subsequent intrusive rocks are largely post-kinematic, highlevel A-type granites and gabbros (497-481 Ma, Foden et al 2002) that were emplaced during 15 km of exhumation (S. P. Turner et al 1992). In the outboard part of the orogen, in the Koonenberry area, deformation was somewhat younger, dated as Middle to Late Cambrian (Mills 2003). Neoproterozoic rocks west of the Bancannia Trough were deformed into westverging structures (Cooper et al 1975). East of the trough, turbidites were juxtaposed against shallow-water limestones and volcanic rocks. Increasing deformation in Cambrian rocks to the northeast (Mills 2002) implies a component of N-S shortening (Fig. 5b). Direen (1997) and Direen & Crawford (20035) used geophysical data to infer a stack of west-vergent thrusts, but more recent interpretation of deep seismic data

43

suggests major west-dipping structures and a major crust al antiform, with an amplitude of c. 15 km and cored by Neoproterozoic strata (Mills & David 2003). In the outboard part of the orogen, in western Victoria, volcanic rocks thrust westwards over the Kanmantoo Group equivalents and older packages are interpreted to be either forearc crust (Crawford et al 1996) or an arc (Scheibner 1989; VandenBerg et al 2000) (Fig. 6). Thrusting predated the 500 Ma post-collisional volcanic rocks in the Mount Stavely Volcanic Complex (Crawford et al 1996) (Fig. 4a). Most of the deformation (D1-D5) and high-T low-P heating were taken up by the sediments in the collision zone west of the volcanic rocks, where a 40 km wide high-grade zone containing synand post-tectonic granites and migmatites passes across faults into lower-grade sedimentary and volcanic rocks to the east and west (e.g. Gibson & Nihill 1992; Gray et al 2002; Kemp et al 2002). Peak deformation and metamorphism occurred around 516 Ma (Turner et al 1993), with cooling ages of 500-480 Ma from metamorphic rocks and granites (Richards & Singleton 1981; Turner et al 1993). Amphibolite-grade rocks occur further east in the Stawell Zone, in the hanging wall of the west-vergent Moyston Thrust (Fig. 6), which has undergone vertical displacement of 15-20 km (Phillips etal 2002). Miller etal (2003) reported that mineral growth occurred at c. 500 Ma in the high-grade rocks, thereby indicating that the Moyston Fault was a Delamerian structure (compare Taylor & Cayley 2000 and Gray & Foster 2000). East of the Moyston Thrust, metamorphic grade drops back to low-grade greenschist facies, with most of the Stawell Zone cut by east-vergent thrusts (VandenBerg et al 2000). The Stawell Zone lacks clear-cut evidence of a Delamerian unconformity. In the Lachlan Orogen, Delamerian (or Tyennan) deformation has been inferred in one locality only - at Waratah Bay (Fig. 5b) - where an unconformity separates basalt from chert (Cayley et al 2002). Spaggiari et al (2003c) suggested, however, that this discordance could represent local features on a topographic high rather than part of a major orogenic event. Conformable Cambrian-Ordovician relationships are, thus, the inferred norm, based on relations in:

Fig. 8. Ordovician terranes in the Lachlan Orogen: (a) distribution; (b) time-space plot. In the Bendigo terrane, the thickness of the Bendigonian 4 to Castlemainian 1 part of the package thins eastwards from 550 m in the west through 450 m down to 24 m in the southwestern corner of the Melbourne Zone (central column). In the eastern part of the Melbourne Zone (right-hand column) the Bendigonian to Darriwilian sequence is <10 m thick. Thicknesses from VandenBerg et al. (2000). Other sources cited in text.

44

1. 2.

3.

R. A. GLEN

the hanging wall of the Heathcote Fault zone, at Lancefield (Cas & VandenBerg 1988; VandenBerg et al 1992); the hanging wall of the Governor Fault in the upper Howqua River (Crawford 1988; Fergusson 1998), where Cambrian volcanic and volcaniclastic rocks pass up into clastic rocks and then Late Cambrian chert, below thick Early Ordovician sandstone; in the Narooma Terrane, where Late Cambrian and Early Ordovician cherts appear to be conformable (Glen et al 20040).

These observations imply that the Delamerian/ Tyennan deformation did not extend into the Central or Eastern subprovinces of the Lachlan Orogen (but see Part 3). Nor is there any record in the New England Orogen (Cawood & Leitch 1985). The Delamerian Orogeny affected the eastern part of the Thomson Orogen, the Anakie Inlier. Rocks here underwent a c. 500 Ma K-Ar cooling that postdates formation of a shallow S2 cleavage (Withnall et al 1996) (Fig. 5c). Fergusson et al (2001) suggested that SI occurred at c. 583 Ma and S2 at c. 540 Ma. In the North Queensland Orogen, the Late Cambrian-Early Ordovician Seventy Mile Group (Henderson 1986) described below from the Benambran cycle along the southern margin of the orogen shows no effects of any Delamerian Orogeny. However, the presence of c. 500 Ma cooling ages just to the north in basement inliers (the Cape River Metamorphics and Barnard Metamorphics; e.g. Draper et al (1998)) implies either that deformation predated the Late Cambrian or that some units are allochthonous. Post-collisional phase Collision in west Tasmania was followed by extension and dismemberment of the thickened crust, leading to formation of the Dundas Trough and the Mount Read Volcanics to the east (Figs 4a, 5c) (Crawford & Berry 1992). These volcanic rocks were erupted between 505 Ma and 495 Ma (Perkins & Walshe 1993) and intruded by late dacites and granites. Continued rifting led to exhumation of the eastern basement and opening of rift basins, filled first by the Owen Conglomerate, with local internal unconformities that date deposition from c. 490 to 470 Ma (Seymour & Calver 1995), passing up into sandstone and then a sag phase platform sequence of Ordovician

limestone (Noll & Hall 2003; 2004) (Fig. 4a). Deformation phases of the Delamerian/Tyennan Orogeny continued into the earliest Ordovician (Seymour & Calver 1995; Holm et al 2003). In the outboard part of the orogen, in the Koonenberry area, post-collisional units consist of basal conglomerate passing up into shallowwater to fluviatile Late Cambrian-Early Ordovician sediments (Mootwingee and Kayrunnera groups) (Mills 2002; 2003) (Fig. 5c). To the north at Gidgealpa, equivalent 'molassic' strata form the lower part of the Dullingari Group, which extends into the Late Ordovician (Gravestock & Gatehouse 1995). Further south in the outboard part of the orogen in Victoria, accretion of forearc (and arc) volcanics was followed by rapid extension, reflected by: 1.

2.

3.

the emplacement of calc-alkaline postcollisional volcanic rocks into the accreted arc (upper part of the Mount Stavely Volcanic Complex) (Crawford et al 1996); formation of accommodation space for accumulation of turbidites of the Glenthompson Sandstone and in the Stawell Zone, the 2-2.5 km thick St Arnaud Group of inferred Middle-Late Cambrian age above the Magdala Volcanics (VandenBerg et al 2000; I. Williams cited in VandenBerg et al 2000; Squire et al 2003) (Figs 4a, 5c, 6); and formation of extensional faults (Cayley & Taylor 1996; VandenBerg et al 2000).

The post-collisional turbidites were themselves deformed in the Late Cambrian, in a second phase of the Delamerian Orogeny, which is constrained by post-tectonic granites, the oldest of which has a U-Pb age of 489 Ma (Crawford et al 1996; VandenBerg et al 2000). The vast amounts of detritus shed eastwards from the eastern part of the orogen suggests that deformation here was also east directed, producing a divergent orogen. Lachlan supercycle The Lachlan supercycle characterizes the Ordovician to Carboniferous history of the Lachlan Orogen and lasts for c. 170 million years, from c. 490 Ma to 320 Ma. This supercycle is divided into three cycles, separated by major contractional deformations that appear to be orogen-wide in space and time, although commonly they may be multiphase, diachronous in detail and variable in intensity. These three cycles are the Benambran (c. 50 million

THE TASMANIDES OF EASTERN AUSTRALIA

years, 490-440 Ma), the Tabberabberan (c. 50 million years, 430-380 Ma) and the Kanimblan (c. 60 million years, 380-320 Ma). All three cycles can be recognized to varying degrees in the North Queensland Orogen, with the Benambran and Tabberabberan cycles also seen in the New England Orogen. The early part of the Benambran cycle corresponds in time with the post-collisional phase of the Delamerian cycle.

Lachlan supercycle 1 (Benambran cycle) Convergent phase. There are four key lithotectonic elements in the Benambran cycle (Figs 6, 7,8): (1) widespread quartz-rich, craton-derived turbidites, overlain in the east by a condensed sequence of black shale; (2) intra-oceanic Macquarie Arc; (3) tholeiitic basalt-chert association with MORB-like chemistry; and (4) the ocean floor Narooma Terrane. An additional element is a very small outcrop of Ordovician limestone (Early Ordovician Digger Island Marlstone lying on Cambrian ? chert) on the Victorian south coast (VandenBerg et al 2000). (1) Early and Middle Ordovician turbidites in the Central and Eastern subprovinces contain both thick and thin beds of sandstone grading up into siltstones and interbedded with multiply-cleaved slates. VandenBerg et al. (2000) recognized several facies. Sandstones are generally quartz-rich with up to 10% detrital feldspar as well as detrital white mica, especially in the lower part of the sequence, which also includes lithic grains (M. Scott & O. Thomas, unpublished work). In the Tabberabbera area, basal sandstones are lithic and lie on Cambrian sedimentary rocks above Cambrian tholeiitic basalts (Crawford 1988; Fergusson 1998). In the Central and Eastern subprovinces, thin chert bands or lenses occur in the Early Ordovician in the Bendigonian (Stewart & Glen 1991) and the Chewtonian (Colquhoun et al. 2004). A prominent chert packet up to 100 m thick near the top of the turbidite packet is late Darriwilian, or Darriwilian-Gisbornian in age (Glen 1992; 1994; VandenBerg & Stewart 1992; Colquhoun et al. 20040) (Fig. 8b). Turbidites, contourites or siltstones above the chert (Jones et al. 1993; Glen 1994; Thomas et al 2002) pass up into a condensed sequence (c. 400 m thick) of black shale and siltstone (Bendoc Group, Warbisco Shale), which spans all or almost all of the Late Ordovician, beginning in the Gisbornian (Fig. 8b) (VandenBerg 1981; VandenBerg et al. 1992; VandenBerg & Stewart 1992; Glen 1992). In some localities, black shales pass up into buff-

45

coloured shales and then local quartz sandstone (Glen & VandenBerg 1987; Thomas et al 2002; Colquhoun et al. 2004). Another packet of Early and Middle Ordovician turbidites occurs in the Southwestern subprovince in central Victoria (Figs 4a, 6, 7, 8). Unlike the other turbidite packets, these contain abundant graptolitic shales, thin from west to east (VandenBerg et al 2000) and lack cherts (Fig. 8b). Unlike the subprovinces to the east, there are Late Ordovician turbidites in this packet that pass eastwards into Late Ordovician black shales (VandenBerg & Stewart 1992). Ordovician turbidites in northeastern Tasmania are dated poorly and consist of the sandstone-rich Tippogoree Group overlain by a slate-rich unit, containing a sole Middle Ordovician graptolite (Reed 2001). Reed (2001) suggested that these Ordovician units should be correlated with similar rocks of the Central subprovince (terminology of this paper) in the Tabberabbera area of central Victoria rather than with rocks of Southwestern subprovince as inferred previously (e.g. Powell et al 19930). Comparison of the different turbidite packets is now possible with the introduction of the new conodont identification techniques of I. Stewart and I. Percival (e.g. Stewart 1988; Glen et al 1990; VandenBerg & Stewart 1992; Percival et al 2003). This comparison shows that rather than representing a homogeneous mud pile (Coney 1992), there are differences in facies, lithology and sequence between turbidite packets across the different subprovinces (summarized in Fig. lOb). These differences lead to the concept that the turbidites were deposited from different fan systems subsequently structurally juxtaposed along major faults. They are thus parts of different terranes (Glen et al 1992; VandenBerg & Stewart 1992; Glen 1993; Glen & Percival 2002; 2003), rather than the one Bengal-size megafan (Cas 1983; Powell 1983; Coney 1992; Fergusson & Coney 1992a). This is discussed further in the last section of the paper. A feature of all these turbidites is the pattern of detrital zircon ages with a dominance of 650 (600-500) Ma ages and a lesser population of 900-1200 Ma detrital zircons (Williams et al 2002). Local differences between different terranes are emerging. The younger population constitutes the Pacific Gondwana population of Ireland et al (1994). Veevers (2000c) suggested that the 600-550 Ma zircons came from the Beardmore-Ross Orogen area of Antarctica, with Goodge (2002) showing that the Ross Orogen began to emerge at 580-520 Ma.

46

R. A. GLEN

Williams et al (2002), on the other hand, suggested that the Mozambique Belt was the most likely source of the zircons. (2) The intra-oceanic Macquarie Arc (Figs 4a, 6, 7, 8) existed from the earliest Ordovician to the Llandovery (494-438 Ma) and is now represented by four main structural belts of volcanic and volcaniclastic rocks, separated by younger Silurian-Devonian rifts (Glen et al. 1998). The Macquarie Arc grew in three pulses (Glen et al 2003; Percival & Glen 2006) (Fig. 8b). Pulse 1, 490-475 Ma, is best developed in the western and central belts, where volcanic and volcaniclastic rocks pass up into siltstones. The early arc formed on a rifted fragment of Cambrian forearc crust, intruded by post-collisional volcanic rocks (Glen et al 2003; Crawford et al 2005) 1000 km east of the old Delamerian margin. It was growing actively at the same time as the Delamerian highlands were rising. After a c. 11 million years hiatus, possibly reflecting back-arc spreading, pulse 2 of arc growth occurred at 466-455 Ma. It is represented by volcanic and volcaniclastic rocks in the western and central belts passing into thinner-bedded turbidites to the east. After a partial hiatus in volcanism in the west of c. 4 million years, pulse 3 activity (c. 452-438 Ma) resulted in coherent volcanic rocks in the western belt and southern part of the central belt and c. 438 Ma (Llandovery) porphyries. In contrast to pulses 1 and 2, pulse 3 rocks are shoshonitic in chemistry (Glen et al 2003; Crawford et al 2005). This third pulse is coeval with Late Ordovician Alaskan-type zoned mafic to felsic igneous complexes intruding back-arc basin turbidites to the west (Suppel & Barron 1986; Elliott & Martin 1991). Early workers maintained that there was interfingering between rocks of the Macquarie Arc and Ordovician craton-derived turbidites. However, there is no provenance mixing (Glen & Wyborn 1997; Colquhoun et al 1999; Meffre et al 2005) and there are faults between these coeval packages, which are thus separate terranes (Glen & Wyborn 1997; Glen etal 1998; Glen & Percival 2003; Glen 2004; Meffre et al 2005). (3) An association of basalts, with MORBlike tholeiitic chemistry, and cherts occurs adjacent to major faults in the Lachlan Orogen - the Gilmore and Tullamore fault zones between the Eastern and Central subprovinces and along the Kiandra-Narromine Structure in the Eastern subprovince. This association has proven or inferred Ordovician ages (e.g. Basden 1990; Lyons & Percival 2002) and is inferred to represent igneous crust to the back-arc basin

turbidites that was exhumed and translated along major faults. (4) The Narooma Terrane (Figs 4b, 6, 8) is an Ordovician oceanic terrane that consists almost wholly of chert, but which coarsens at the top into argillite, siltstone, sandstone and conglomerate as it approached the Gondwana margin. This interpretation follows the model of Glen et al (20040) rather than those of Miller & Gray (1996) or Fergusson & Frikken (2003). Elements of the Benambran cycle are also recognized from other orogenic belts. In the Delamerian Orogen, it is represented by platform deposits in west Tasmania (largely Gordon Limestone) and by shallow-water clastic rocks in the Koonenberry Belt (Figs 4,7). In the southern New England Orogen, the Benambran cycle is represented both west and east of the Peel-Manning Fault System (Glen & Scheibner 1993) (Figs 4b, 7). To the west in fault slices are arc-derived sediments, with blocks of Late Ordovician limestone (Cawood 1976), that lie unconformably on Early Ordovician and Cambrian strata (Cawood & Leitch 1985). The Trelawney Beds' of Philip (1966) are probably limestone olistoliths in Early Devonian units (P. A. Cawood cited in Furey-Greig 1999). In Queensland, Ordovician limestone occurs in the Devonian Calliope arc. Benambran elements in accretionary complex rocks east of the Peel-Manning Fault System include: (a) limestone blocks with Eastonian conodonts immediately east of the fault (Furey-Greig 1999). They probably represent seamount cappings (Scheibner 1973; Flood 1999), although Furey-Greig (1999) preferred a sedimentary derivation from volcanic islands and fringing reefs to the west; (b) gabbro in the Peel-Manning Fault System at Attunga (locality of Attunga (at), Fig. 7) with zircons dated at 460 Ma (Watanabe et al 1998) but corrected to 479 Ma by Fanning et al (2002). Further south, phengite in a gabbro gave a K-Ar age of 481 Ma (Fukui et al 1995); (c) Ordovician blueschist and related rocks. At Port Macquarie (p, Fig. 7), white mica has yielded K-Ar ages of 444 Ma (M. A. Lanphere quoted by Scheibner 1985) and ages of 476 and 471 Ma (Fukui et al 1995). White mica K-Ar ages of 482-467 Ma from blueschist knockers and related rocks (originally basalts and gabbros) in schistose serpentinite matrix along the PeelManning Fault System at Glenrock probably represent uplift ages: subduction

THE TASMANIDES OF EASTERN AUSTRALIA

47

the orientation of subduction - was it east-west (Henderson 1986), north-south (e.g. Kay 1998) or was it rotated from a SW-NE trend, parallel There are few data on the Benambran cycle to the Diamantina River Lineament with in the Thomson Orogen. The generally low- subduction to the southwest (Stolz 1995)? grade metasediments encountered in drill holes Palaeomagnetic evidence in support of a 90° in the west are undated and Murray (1994) clockwise rotation between 425 Ma and 400 Ma suggested they are Cambrian to early Late was presented by McElhinny et al (2003). This Ordovician in age. Local metamorphic rocks is consistent with the presence, to the present have Cambrian to Early Ordovician K-Ar mica north, of rapidly deposited, moderately poorly ages, and intrusive granites have Late Ordovi- sorted Ordovician quartz-rich turbidites derived cian to Devonian K-Ar mica ages. On the from the craton to the west (Bain & Draper eastern margin of the orogen (Fig. 7), the Late 1997). These turbidites constitute the Mulgrave Ordovician Fork Lagoons beds (quartz-rich Formation in the Hodgkinson subprovince (but sandstones, conglomerates, limestones and contain more immature sandstones in the east) felsic volcanic detritus and basalts with island- and the Early Ordovician Judea beds in the arc tholeiite affinity; Withnall et al 1996) (Fig. Broken River subprovince (Donchak 1993) 4c) are faulted against the Neoproterozoic- (Figs 4c, 9(1)). Lenses of mafic volcanic rocks Cambrian rocks of the Anakie Inlier. Withnall and jaspers in the turbidites and in slices along et al (1996) suggested correlations with arc- the Palmerville Fault System may be similar to related Late Ordovician units in the North the back-arc basin basalts in the Lachlan Queensland Orogen and it is possible the Fork Orogen, although they also may relate to a Lagoons beds indicate extensions of this orogen suprasubduction zone environment (Bain & to the south. Draper 1997). The Benambran cycle is also represented in Middle to early Late Ordovician tectonism in the North Queensland Orogen, which contains the North Queensland Orogen in the Hodgkinevidence for Early and Late Ordovician subduc- son subprovince is reflected by deposition of tion (Figs 4c, 7). The Lolworth-Ravenswood Richmondian (c. =Bolindian) limestone above Block (Fig. 2) contains the c. 12 km thick Late the Mulgrave Formation (Nicoll cited in Bain & Cambrian to Early Ordovician Seventy Mile Draper 1987) and beneath a conglomerate Group, which shows a geochemical progression containing clasts of veined quartz-sandstone, from intraplate basalt, andesite and alkaline derived from the Mulgrave Formation, and magmas at the base to subduction-related, low- felsic-intermediate volcanics derived from an to medium-K calc-alkaline lavas higher up arc (Donchak 1993). A 455 Ma U-Pb age by (Henderson 1986; Stolz 1995; Berry et al 1992; C. M. Fanning on one dacite clast (cited by Perkins et al 1993; Paulick & McPhie 1999). Donchak 1993; Bain & Draper 1997) suggests Tectonically, this group represents initial rifting that these Late Ordovician sediments accumuof a continental margin, leading to formation of lated in a back-arc basin behind a west-dipping a back-arc basin just inboard of a continental Late Ordovician subduction zone, and that a margin arc no longer preserved (Henderson continental arc was established on a basement 1986; Stolz 1995). Basement inliers were of 500 Ma metamorphic rocks (Fig. 9(3)). deformed strongly, with formation of a flat-lying Initiation of subduction in the Middle Ordoviextensional fabric at this time (Henderson et al cian is approximately coeval with closure of the 2004). Early Ordovician sedimentary basin to the west This suprasubduction zone setting is (Fig. 9(2)). consistent with the subduction signature of the coeval Late Cambrian to Middle Ordovician I- Collisional phase. Evidence for some type of type, medium- to high-K granitoids in the Middle Ordovician tectonism occurs in three Ravenswood Batholith (Hutton et al 1994; Bain orogenic belts. In the Lachlan Orogen, & Draper 1997). Model ages of 1120-1230 Ma, 455-458 Ma Ar-Ar mica ages were obtained detrital zircons of 1100 Ma (Bain & Draper from quartz veins in the Southwestern 1997; Draper et al 1998) and large negative subprovince (Foster et al 1999). In the southern epsilon Nd indicate the involvement of old New England Orogen, there is a mid-Ordovicrust. Bain & Draper (1997) attributed the cian unconformity (Cawood & Leitch 1985). In subduction signature to derivation from melting the North Queensland Orogen, Middle Ordoviof underplated mafic melts. cian deformation is recorded by an inferred The anomalous c. 150 km east-west trend of unconformity above quartz-rich sediments of the Seventy Mile Group raises questions about the Mulgrave Formation and below Late was, thus, earlier (Fukui et al 1995; Offler 1999).

48

R. A. GLEN

THE TASMANIDES OF EASTERN AUSTRALIA

Ordovician arc-derived conglomerate in the Hodgkinson subprovince (Fig. 9(3)). This deformation is attributed to closure of the 'Mulgrave Basin' related to initiation of subduction to the east (Donchak 1993; Bain & Draper 1997). A major end Ordovician-Early Silurian deformation was responsible for thrusting between the craton and the western margin of the orogen (Bain & Draper 1997). In the southern part, it produced multiple deformation in the Mt Windsor Group, especially the east-west strike of the unit that reflects a structural position on the south-dipping limb of a D2 fold (Berry et al 1992). In the Lachlan Orogen, the Benambran cycle was terminated by the Benambran Orogeny that affected all of the orogen, except for the Melbourne Zone in the Southwestern subprovince. Ar-Ar mica ages indicate the Bendigo Zone underwent formation of eastvergent folds, thrusts and related cleavage around 440 Ma (Foster et al 1999). In contrast, in the Central and Eastern subprovinces, the Benambran Orogeny consists of two identifiable phases phases over c. 10 million years (Packham 1969; VandenBerg 1999; Collins & Hobbs 2001; Glen et al 20040) (Fig. 4b). Phase 1 occurred around the end of the Ordovician (c. 443 Ma) and was marked by deformation of Ordovician turbidites and overlying black shales by folding, thrusting, major strike-slip faulting and multiple cleavage formation in response to east-west and north-south shortening. Part of the extinct Macquarie Arc was upthrust as it was accreted into back-arc basin turbidites (Glen et al 20040). Phase 1 was followed by extension leading to formation of deep- and shallow-water Llandovery sedimentary basins and is thus manifest by a major facies change from black shales to Llandovery turbidites (Glen et al 20040). Phase 2 (c. 433-430 Ma) of the Benambran Orogeny was marked by oblique thrusting of Ordovician turbidites in the Central subprovince over the Macquarie Arc, which was being translated to the southwest. Uplift of other parts of the Macquarie Arc, the folding and thrusting of Llandovery turbidites (VandenBerg 1999), and the syn-tectonic emplacement of some granites (VandenBerg 1999) are also parts of phase 2. In eastern Tasmania the Benambran Orogeny is poorly constrained in time, but involved recumbent folding with vergence to the northeast (Reed 2001).

49

In the Koonenberry area of the Delamerian Orogen, the Benambran Orogeny is expressed by ESE-trending cleavage in post-Delamerian strata (Mills 2002). Mills also suggested that an ESE-trending fold belt in the SW corner of the Thomson Orogen formed during the Benambran Orogeny. In the New England Orogen, the Benambran Orogeny coincides with the hiatus between the Middle Ordovician Haedon Formation and Early Devonian strata at the base of the Tamworth Group (Cawood & Leitch 1985).

Lachlan supercycle 2 (Tabberabberan cycle) Lachlan cycle 2 affected the Lachlan Orogen and the North Queensland Orogen, both of which have extensional character, lying behind the Gondwana-proto-Pacific plate margin identified from the New England Orogen by the presence of an intra-oceanic arc-subduction system. Granites of this cycle also intrude basement west and north of the North Queensland Orogen and the Delamerian Orogen west of the Lachlan Orogen back-arc basin elements Convergent margin phase. Elements of a Late Silurian-Middle Devonian arc occur in faultbounded blocks along the western flank of the Peel-Manning Fault System in the southern New England Orogen (Figs 4b, lOa). These volcanic rocks lie below Middle DevonianCarboniferous strata of the Tamworth Trough, although they were placed originally in that forearc basin sequence by Crook (1960). Two volcanic packets are present: one Late Silurian to Middle Devonian and the other Middle to Late Devonian in age. The older packet of volcaniclastic, extrusive and intrusive rocks was interpreted as an intra-oceanic arc with a low-K calc-alkaline signature by Cawood & Flood (1989) and Offler & Gamble (2002). Both papers suggested the arc developed above a west-dipping subduction zone, although Aitchison & Flood (1995) argued for east-dipping subduction. Offler & Gamble (2002) suggested that the Middle-Late Devonian volcanic rocks represented inter-arc rifting (see also Stratford 1993; Aitchison & Flood 1995; Stratford & Aitchison 1996). The significance of a 436 Ma U-Pb zircon age from tonalite in the Pigna Barney Ophiolite Complex is uncertain: it could be related to pieces of an Ordovician arc or to

Fig. 9. Geological history of the central and northern parts of the North Queensland Orogen, Hodgkinson subprovince. From Paul Donchak, Geological Survey of Queensland, with permission.

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R. A. GLEN

THE TASMANIDES OF EASTERN AUSTRALIA

Pb loss from the Cambrian arc (Kimbrough et al 1993). In the northern New England Orogen, Upper Silurian to Middle Devonian rocks within the area of the Yarrol Trough were placed in an island-arc setting by Marsden (1972) and grouped into the Calliope Volcanic Arc by Day et al (1978) (Fig. lOa). These rocks are dominantly shallow-marine volcaniclastic sediments with varying amounts of felsic to mafic volcanic rocks. From geochemical data, Morand (1993) suggested that these rocks formed in a continental margin arc setting, whereas Offler & Gamble (2002) used REE data to argue for an island-arc setting. Messenger (1996) suggested a continental island arc undergoing local rifting. Murray (2003) suggested the 380 Ma Mt Morgan trondhjemite has an arc or rifted-arc geochemistry. In contrast, Bryan et al. (2003Z?) suggested formation in a back-arc basin and argued that zircon and Pb isotopic data support the involvement of continental crust in magma generation. In the southern New England Orogen, the accretionary complex of the Tabberabberan cycle occurs as a narrow 10-20 km wide terrane east of the Peel-Manning Fault System (Fig. lOa). This is the Woolomin Terrane (Cawood & Leitch 1985), broadly correlating with the Djungati Terrane of Aitchison et al. (19920), which comprises a fault-repeated basalt-chertsandstone package (Woolomin Group) (Cawood 19825; Aitchison et al. 19920) (Fig. lla). Radiolaria indicate a mid-Silurian age for basalt, Late Silurian-Frasnian age for chert and a Famennian age for siliceous chert and overlying volcanogenic sandstone (Ishiga et al. 1988; Aitchison et al. 19920). Limestone olistoliths in the sandstone range in age from Late Ordovician through to the Devonian (Cawood 1980; Pickett 1982; Furey-Greig 1996; Flood 1999) and may represent partly accreted seamounts (Scheibner 1973; Aitchison et al. 19920; Flood 1999). In the northern New England Orogen, the presence of a Late Silurian-Middle Devonian conodont fauna (Fergusson et al. 1993) indicates some accretion in the Tabberabberan cycle, but accretion was mainly younger. Back-arc system. In the Lachlan Orogen, the Tabberabberan cycle is characterized by major

51

Fig. 11. Stratigraphy of two New England Orogen terranes: (a) Djungati terrane from the Tabberabberan cycle; (b) Anaiwan terrane in the Hunter-Bowen cycle. After Aitchison et al. (19920) with permission of the journal of Palaeogeography, Palaeoclimatology and Palaeoecology, and the authors.

basin formation and emplacement of granitoids (Figs 4, 6, 10). Basins in the Eastern, Centr and Western subprovinces formed by rifting or transtension. In the absence of published subsidence curves, three criteria suggest that basins in the Eastern, Central and Western subprovinces had rift and/or transtensional origins: 1.

the presence of bounding and internal growth faults (e.g. Stuart-Smith 1990; Glen et al. 1996; Glen 1999; VandenBerg et al. 2000) revealed by thickness variations and/or facies changes (e.g., megabreccias

Fig. 10. Tabberabberan cycle, (a) Convergent margin and back-arc units, (b) Collisional phase showing complex thrust ± strike-slip faulting in the Lachlan Orogen with directions of maximum horizontal shortening indicated by open red arrows, (c) Simplified time-space plot of some Tabberabberan basins in the Lachlan Orogen.

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2.

3.

R. A. GLEN

on the margins of the Early Devonian Buchan Rift (VandenBerg et al 2000); the change in basin fill, from volcanic and/or coarse elastics at the base, passing up into limestone and shale, and/or finegrained elastics, at the top. These changes suggest that basin growth may be divided into syn-rift and post-rift or sag phases (Fig. lOc) and, thus, that these basins formed as rifts, half-graben or transtensional basins (Glen 1992); the rift geochemical signature of volcanic and volcaniclastic rocks, with most volcanic rocks plotting in the fields of an incipient back-arc, probable back-arc extension, continental rift setting, or even intra-arc settings (e.g. Collins 2002). Some volcanic units have subduction zone signatures indicative of melting of crustal or mantle material modified previously by subduction (Basden 1990; Dadd 1998; Watkins 19986) (see also Part 3).

Using these criteria, it appears that most/all of the Middle Silurian-Middle Devonian basins in the Eastern and Central subprovinces formed during rifting that commenced around the lower part of the Wenlock. Multiple rift-sag sequences occur within some basins, indicating renewed extension in the middle to upper Lochkovian (Fig. lOb), leading to the formation of intrabasin unconformities below Lochkovian conglomerates from the Canberra-Yass Shelf and Goulburn Trough. Although attributed previously to the contractional Bowning Orogeny (Brown et al 1968; Crook & Powell 1976; Owen & Wyborn 1979; Felton 1976; Bain et al. 1987), the unconformities are regarded here as extensional in nature. It is suggested that the presence or absence of volcanic rocks in rift basins can be used to divide rifts into hot ones that contain felsic volcanic and volcaniclastic rocks with or without comagmatic granites, and cold ones filled either by epiclastic or siliciclastic detritus (unpublished data). The Hill End Trough and Cobar Basin are the only two major cold rifts in the Central and Eastern subprovinces (Fig. 6) and both lie north of a major cross fault - the Lachlan Transverse Zone (Glen & Walshe 1999). This largescale difference in heating of the extended crust is still not understood, but appears to be largely independent of basement type, with the Cobar Basin developed on a basement of deformed turbidites (Glen et al 1996) and the Hill End Trough on a basement of Ordovician volcanic rocks (Glen et al 2002; David et al 2003) To the south and west, the Southwestern

subprovince was occupied by one large basin, the Melbourne Trough (now Melbourne Zone, Fig. 6), that was open to the east, and filled gradually from the west, south and southwest, right though to the late Early Devonian (Garratt 1983; VandenBerg et al 2000; Powell et al 2003), when detritus began to be supplied from an encroaching eastern source. In contrast to evidence of widespread extension in the three subprovinces above, there is little evidence of extension in the Southwestern subprovince. The Melbourne Trough is best understood as a composite foreland basin, yoked for some of its life to the deforming Bendigo structural zone to the west (western part of the Bendigo terrane) (Gray & Foster 1998; VandenBerg et al 2000; Cayley et al 2002) and containing packets of sandstone derived from the southwest and west as well as unknown areas to the southwest and south. In the Emsian, the Melbourne Trough was also yoked to the uplifted Tabberabbera Zone to the east, receiving volcanic-rich detritus from the east (Powell et al 2003). Despite this foreland basin model, some extension in the eastern sector is reflected in part by the emplacement of the Middle Devonian (375 Ma) Woods Point Dykes with shoshonitic chemistry (Bierlein et al 2001). In eastern Tasmania, the Tabberabberan cycle is represented by turbidites of the Ludlow to Pragian Panama Group, intruded by Early Devonian I- and S-type granitoids (Reed 2001). In the western Tasmanian part of the Delamerian Orogen, the absence of a Benambran Orogeny means the stable Ordovician platform (Gordon Group) continued into the Late Silurian, with shelf deposition of the Eldon Group (Seymour & Calver 1995). Reed (2001) pointed out that the overlying deep-water Early Devonian turbidites were correlatives of the Panama Group in eastern Tasmania, thereby indicating that the two terranes were close to each other by then. In the western Victoria part of the Delamerian Orogen, the Tabberabberan cycle is represented by deposition of fluviatile and marginal marine sandstone-rich sediments of the Grampians Group, which are Late Silurian in the upper part (VandenBerg et al 2000) (Figs 6, 13a). This unit was then thrust imbricated, extended (420-410 Ma) and, subsequently, intruded by I- and A-type Early Devonian 410-400 Ma post-tectonic granites (Cayley & Taylor 1996). These were then overlain by the Rocklands Rhyolite dated at 410 Ma (VandenBerg et al 2000). In the Koonenberry area of the Delamerian Orogen, the Tabberabberan cycle is represented

THE TASMANIDES OF EASTERN AUSTRALIA by deposition of the Late Silurian-Devonian Mt Daubney Formation (Fig. 6), followed by its deformation in the late Early Devonian (Neef et al 1989; Buckley 2003). In the Thomson Orogen, the Tabberabberan cycle is represented by development of the Adavale Basin and adjacent troughs that are concealed beneath Permo-Triassic basins and younger cover (Figs lOa, b). Early Devonian felsic volcanic and volcaniclastic rocks pass up with apparent conformity into Middle Devonian marine sediments and evaporites (Murray 1994). In the North Queensland Orogen, there is no major mid-Devonian deformation: deformation in the Broken River and Hodgkinson subprovinces occurred later, around the Devonian-Carboniferous boundary and, as a result, the Tabberabberan cycle extends into the early Tournaisian. There are, however, lowangle unconformities and disconformities around the Middle-Late Devonian boundary in both the Lolworth-Ravensworth and Broken River subprovinces (Bain & Draper 1997). The Tabberabberan cycle was marked by the formation of rift basins, filled by limestone, turbidites and shallow to continental siliciclastic rocks, and by the emplacement of granites. These rocks are mainly overlain by Visean volcanic rocks deposited in graben (Bain & Draper 1997). In the Hodgkinson subprovince, the Tabberabberan cycle began with extensional faulting that created accommodation for deposition of thin-bedded siliciclastic turbidites in the lower part of the Chillagoe Formation (Fordham 1993) and early to mid-Llandovery submarine volcaniclastic rocks and basaltic lavas (of possible MORE affinity), especially in the north. By the middle to late Llandovery, carbonate platforms developed in shallow subbasins. Coeval deeper sub-basins received turbidites and continent-derived sediments. Carbonate deposition generally ceased in the Early Devonian (around 410 Ma, in the late Lochkovian), with the collapse of the carbonate platform in a renewed phase of basin extension, characterized by deposition of craton-derived breccia fans and conglomerates in submarine channels (Fordham 1993). This event was synchronous with the emplacement of the Cape York Peninsula Batholith in basement rocks to the west (Fig. 2) (Domagala et al 1998). Conglomerates were succeeded by quartzintermediate turbidites of the Hodgkinson Formation that covered a large area (Fig. 9(5)) and which are at least as young as mid-Famennian (Domagala et al. 1998). Most of these

53

turbidites were derived from Proterozoic basement and >420 Ma Silurian-Devonian granites, although some sandstones in the east contain detritus from the old Ordovician arc (Domagala et al 1998). The presence of lenses of submarine basalt, chert and local limestone and broken formation has been used in the past to invoke an accretionary prism setting for the Hodgkinson Formation (Henderson 1987), but the back-arc basin rift model of Fawckner (in Arnold & Fawckner 1980), Donchak (1993) and Domagala et al (1998) is followed in this paper. The Broken River subprovince to the southwest is dominated by siliciclastic rocks, although there is a local regression leading to formation of carbonates in the Emsian-Givetian. These pass up across a low-angle unconformity to disconformity into shallow-water to alluvial strata in the pull-apart Bundock Creek Basin. In the Lolworth-Ravenswood Block, Givetian arkoses, limestone and sandstone are overlain with local unconformity by Frasnian-Famennian continental sediments that pass up into alternating shallow-marine to continental sediments (fill of the Burdekin Basin) below a late Tournaisian unconformity and younger felsic volcanic rocks. Tabberabberan cycle granites. Tabberabberan cycle granites occur in North Queensland and in the Lachlan Orogen. In North Queensland, granites were emplaced into the North Queensland Orogen and also into Precambrian basement to the west (e.g. the c. 430 Ma Nundah Granite, Fig. 9) (Bultitude et al 1993). Basement northwest of the North Queensland Orogen was intruded by the >450 km long Cape York Peninsula Batholith, which consists of a core of I-type granites surrounded by dominant S-type granites that are mainly Early Devonian in age. Both types have Early Proterozoic model ages, suggesting derivation from underlying old continental crust. Within the North Queensland Orogen, Tabberabberan cycle granites in the Ravensworth Batholith in the LolworthRavensworth Block comprise regional aureole I- and S-type granites (the oldest, 435-424 Ma), followed by 425-410 Ma oxidized I-type granites, 410-400 Ma regional aureole granites and 400-380 Ma high-level granites. The oxidized I-type granites have subductionrelated chemistry, but were believed by Draper & Bain (1997) to have been derived from mid-deep crustal sources as a result of underplating by mafic magmas during regional extension. In the Lachlan Orogen, Tabberabberan cycle granites are prevalent in the Eastern and Central subprovinces where they occupy up to

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R. A. GLEN

36% of the surface area (Chappell et al 1988) (Fig. 6). Granites are divided into I-type granites in the east and mixed I- and S-type granites in the west (Chappell & White 1974): very few granites have mixed or zoned I-S character. I-types are subdivided further into high and low temperature types and fractionated or non-fractionated types (Chappell et al. 1998; 2000). Some granites were emplaced into the extending upper crust, where they are comagmatic with, and overlain by, Late Silurian and Early Devonian felsic volcanic rocks erupted in extensional basins (Wyborn & Chappell 1986). Other, foliated, granites were emplaced into the middle crust at depths of c. 10 km and then exhumed by thrusting a further 1-2 km (V. Morand, cited in Glen & Wyborn 1997). Granite emplacement ages migrate eastwards from the Silurian in the Central subprovince into the Silurian to Devonian in the Eastern subprovince (e.g. Powell 1983, 19840; Collins 2002), but there are young Late DevonianCarboniferous granites in central Victoria in the Southwestern subprovince as well in the eastern part of the Eastern subprovince. Some of this 'migration' might actually reflect cooling patterns, since recent zircon dating has shown that granites thought previously to be Late Silurian to Early Devonian from K-Ar and Rb-Sr dating are older and formed during early parts of the Tabberabberan cycle, if not at the end of the Benambran cycle, e.g. Cooma granodiorite (Williams 2001). If the model of Glazner et al. (2004) applies to Lachlan granites, individual plutons may have been emplaced over millions of years, and can no longer be represented by a single zircon date. Collisional phase. In the southern New England Orogen, Flood & Aitchison (1992)

Fig. 12. Kanimblan cycle and Hunter-Bo wen supercycle. (a) Fluviatile Lambie facies represented by mustard-coloured units in the Lachlan, Delamerian and Thomson orogens. Hachured area in the Lachlan Orogen represents area of concealed Lambie facies developed above Tabberabberan cycle strata. Elements of Hunter-Bowen supercycle also shown: back-arc Drummond Basin, continental margin arc, forearc basin (tt, Tamworth Trough; yt, Yarrol Trough) and subduction complex and accreted terranes outboard of major faults (pmfs, Peel-Manning Fault system; yfs, Yarrol Fault System), (b) Collisional and post-collisional phase, Kanimblan cycle.(c) Hunter-Bowen supercycle, deformation phase, cycle 2.

THE TASMANIDES OF EASTERN AUSTRALIA suggested that the intra-oceanic arc of the Tabberabberan cycle became accreted to the Australian plate in the Late Devonian (Fig. lOb). This is controversial, however, and no clear evidence of contractional deformation related to arc accretion has been identified. The candidate Bective unconformity below the Late Devonian Keepit Conglomerate is a semiregional disconformity only, and is not present in the centre of the Tamworth Trough (Russell 1979). In the northern New England Orogen, proponents of the intra-oceanic model for the Calliope arc argue that an unconformity close to the Middle-Late Devonian boundary reflects accretion of the arc to the Gondwana margin (e.g. Murray et al 2003). Others (Leitch et al 1992; Morand 1993) have argued that the unconformity is low angle, deformation was minor and that there was no break at the base of the overlying sequence (see also Bryan et al. 2003); C. Murray (pers. comm. 2004) points to removal of a major stratigraphic unit along the unconformity. In the Lachlan Orogen, major basin inversion occurred in the late Early-Middle Devonian, with earlier Devonian deformations reflecting localized strike-slip tectonics. There are differing views as to whether these deformations are part of a longer-lived deformation prograding from the west (Gray & Foster 1997; Gray et al 1997), or are related to the oblique strike-slip collision of an allochthonous terrane into the Southwestern subprovince (Glen et al. 1992; VandenBerg & Stewart 1992; Willman et al 2002; see Part 3). Middle Devonian orogeny is reflected by basin inversion, with growth faults undergoing varying degrees of reverse/oblique reactivation and development of thin-skinned and basininversion structures (e.g. Glen 1995) (Fig. lOb). In the Eastern subprovince, Middle Devonian basin inversion involved fold and cleavage formation and NE-SW thin-skinned thrusting in the Hill End Trough (Vassallo et al 2003), deformation of the shelf to the west (c. 380 Ma cooling ages, Glen et al. 1999) and in the 'Captains Flat Trough' to the south (370-380 Ma K-Ar ages on sericite, 374 Ma Rb-Sr age on biotite, Abell 1993). The Jemalong Trough (Fig. 6) (Raymond et al. 2001) and Cowra Trough were deformed at c. 393 Ma (Raymond et al 2001), while the overlying Middle Devonian A-type volcanic rocks of the Rocky Ponds Group (Dulladerry Volcanics) were deformed c. 372 Ma (Raymond et al. 2001). The presence of metamorphic biotite in deformed sedimentary rocks of the Hill End

55

Trough (Vernon & Flood 1979) and 'Captains Flat Trough' (Smith 1969; Barron 1999; Abell 1993) and metamorphic white mica in the eastern part of the Cobar Basin (Brill 1989; Glen et al 1992) implies more intense deformation and structural thickening than first apparent (e.g. Vassallo et al. 2003). The N-S elongate shapes of some Silurian-Devonian granites (Fig. 6) also reflect this Middle Devonian deformation, with the formation of solid-state foliations and well developed S-C fabrics and mylonite zones on the (commonly) eastern margins of mid-crustal granites such as the Wyangala Batholith (Morand 1988; Paterson et al 1990; Lennox et al 1998) and Wologorong Batholith (Vernon et al. 1983). Cooling ages of c. 380 Ma (Foster et al 1999; Glen et al 1999) have been obtained from Ar-Ar dating of metamorphic biotite. The eastern, highly strained, margin of the Murrumbidgee Batholith might also have been deformed at this time. No evidence is found for the concept that rising granites were responsible for deformation seen in Silurian-Devonian rifts through downward movement (or sagductional deformation) of upper crustal rocks lying between plutons or batholiths (Warren & Ellis 1996; Blevin 1998). The last stage of this Tabberabberan Orogeny in the Eastern subprovince was marked by brittle conjugate NE- and NW-trending faults offsetting major plutons. These faults produced small amounts of extension parallel to the orogen and were, thus, able to distribute the decreasing amounts of east-west shortening in a crust strengthened by cooling granites (Lambert & White 1965; Glen 1992). An earlier deformation in the Early Devonian is concentrated along the boundary between the Eastern and Central subprovinces and reflects thrusting and translation of the Central subprovince to the SSE (Glen 1991; Morand & Gray 1991). Deformation was concentrated along a linked system of major and splay faults - the Tullamore Fault Zone, Gilmore Fault Zone and Indi-Long Plain Fault - that marks the boundary between these subprovinces (Fig. 6). To the east, the Tumut Trough was inverted between 418 Ma and 411 Ma (Basden 1990) and the largely transtensional rifts of the Cowombat Trough in Victoria were deformed by the Bindian deformation around 418 Ma (VandenBerg et al 2000). This Bindian deformation is also recorded by Ar-Ar dates of 409 Ma on cleavage sericite at Peak Hill (Perkins et al 1995) and 411 Ma on sericite in the Booberoi Fault Zone (dating by Foster et al. 1999, cited by Lyons 2000). To the west, the

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eastern margin of the Cobar Basin was deformed at 400-395 Ma, according to wholerock Ar-Ar dating (Glen et al 19926). Movement on the Indi Fault is dated at 405 Ma (Foster et al 1999) while faults near the leading edge of the south-moving Central subprovince have Ar-Ar mica ages of 413-395 Ma (Foster et al 1999). The latitudinal boundary with the Thomson Orogen underwent major dextral strike-slip (Glen et al 1996) probably after contractional dip-slip faulting. Interpretation of geophysical data, coupled with limited outcrop information, suggests that significant basin inversion also occurred in the Western subprovince in the mid-Devonian (Glen et al 1996). In the Southwestern subprovince, most Middle Devonian structures formed in response to broadly east-west shortening coupled with lesser north-south shortening. Effects of both shortenings are best seen in the Melbourne Zone, where the low-strain, western part contains curvilinear (NE)-NW-trending folds, refolded about east-west folds (see summary in VandenBerg et al 2000). East-west strain becomes higher eastwards, with the eastern Melbourne Zone containing west-dipping thrusts (Gray 1995) and duplexes in Cambrian igneous units (VandenBerg et al 1995). Ar-Ar dating of mica gives ages of 415-390 Ma (Foster et al 1999) (Fig. 6). A two-stage Tabberabberan deformation is present in eastern Tasmania. Early Devonian, pre-granite, NE-vergent structures were followed in the late Early to Middle Devonian by syn-granite structures as this part of the Lachlan Orogen was thrust southwest over western Tasmania (Reed 2001). As a result of this deformation, the Delamerian Orogen in western Tasmania underwent tightening of older Delamerian folds, generation of NWtrending folds, cleavage and faults, and heating that led to the intrusion of syn- to post-tectonic granites (Berry 1994). No clear Tabberabberan deformation has been recognized from the concealed Adavale Basin, although Evans et al (1990) suggested the basin was converted to a foreland basin around this time. In the North Queensland Orogen, the Tabberabberan cycle was brought to a close by an Early Carboniferous deformation (Fig. 9(6)). In the Hodgkinson subprovince, this is best constrained by a syn-tectonic 357 Ma granite that was intruded during formation of a widespread S2 foliation (Davis et al 1998). In the Broken River subprovince, deformation separates little-deformed Tournaisian sedi-

mentary rocks in the Clarke River Basin from underlying strata (Zucchetto et al 1999) (Fig. 4c). This deformation did not apparently affect the Lolworth-Ravenswood subprovince to the southwest, where sedimentation in the Burdekin Basin continued into the Tournaisian.

Lachlan supercycle 3 (Kanimblan cycle) Elements of the Kanimblan cycle occur in the Lachlan and Delamerian orogens (Figs 4,12a). Rifting and loading phases. In the Eastern and Central subprovinces of the Lachlan Orogen, a brief late Early-Middle Devonian rifting event marked the onset of the Kanimblan cycle and generated restricted A-type volcanic rocks and local granites and rift-related sediments (Mcllveen 1974; Dadd 1992; Wyborn & Owen 1986; Raymond 1996; Collins 2002) (Figs 4, 6). In some areas there is a low-angle unconformity beneath 3-4 km of largely fluviatile sandstone, conglomerate and siltstones/mudstones that were deposited in one or more terrestrial basins and which constitute the Lambie (or Lambian) facies of Powell (19846) (Fig. 12a). However, in general, rocks of the Lambie facies lie on a range of older stratigraphic units that were assembled in the preceding Tabberabberan Orogeny (Glen & Watkins 1999). In the Western, Central and Eastern subprovinces, rocks of the Lambie facies either pass up from a basal marine interval into continental fluviatile deposits, or contain a marine interval (that disappears westwards) near their base. Further west, in the Central and Western subprovinces, the Mulga Downs Group (Glen et al 1996) contains, in its lower part, a late Early-Middle Devonian package (Bembrick 1997). Rocks of the Lambie facies are overwhelmingly quartz-rich and possibly derived from major uplift in central Australia (Alice Springs Orogeny), although there are local facies variants, especially near the base and variations in palaeocurrent directions. One prominent interval with volcanic detritus was derived from the south (Powell 19846), another from the arc in the New England Orogen to the north (Powell et al 1984). It is still uncertain whether the Lambie facies formed a continuous 3-4 km thick blanket across the Lachlan Orogen or was deposited in interconnected intermontane basins (Powell 19846). O'Halloran & Cas (1995) favoured intermontane basins in central Victoria and recognized intra-formational unconformities that they attributed to syndepositional contractional deformation, probably reflecting long-lived movement on the

THE TASMANIDES OF EASTERN AUSTRALIA

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Governor Fault Zone between the Eastern and Central subprovinces (Fig. 6). Different relations occur in the Southwestern subprovince. Here, the Kanimblan cycle began with the widespread intrusion of high-level Sand I-type granites and the outpouring of associated felsic volcanic rocks preserved in cauldron complexes of the Central Victorian Magmatic Province (Figs 4a, 6,12a). Local mafic magmas are shoshonitic, characteristic of postorogenic regions (VandenBerg et al. 2000). The succeeding Lambie facies, while broadly similar to units further east, contains a prominent volcanic interval marked by rhyolitic ignimbrites derived from cauldron activity and is intruded by high-level granites (VandenBerg etal 2000). Lambie facies rocks in the Delamerian Orogen are restricted to the Koonenberry area, where a late Early-Middle Devonian unit and a Late Devonian unit are preserved in the largely subsurface Bancannia and Menindee troughs as well in the Koonenberry Belt (Sharp & Buckley 2003; Neef 2004) (Figs 6, 12a). In the Thomson Orogen, Late Devonian redbeds in the upper part of the concealed Adavale Basin correspond to rocks of the Lambie facies (Murray 1994). The concealed E-W- to WSWtrending Paka Tank Trough in the southern part of the orogen is of this age as well (Alder 1999) (Fig. 12a). In the North Queensland Orogen, Lambie facies rocks are restricted to a c. 600 m thick sequence of siltstones and tuffaceous sandstone in the Lolworth-Ravensworth Block and in the Bundock Basin in the Broken River subprovince (Bain & Draper 1997).

regional cleavage-forming event in the NE of the Eastern subprovince (dated at c. 340 Ma) was followed by granite emplacement (Pogson & Watkins 1998). Carboniferous strain decreases to the south and west, where Lambie facies rocks have been deformed into large N- to NW-trending broad synclinoria and anticlinoria that may also show effects of north-south shortening (Figs 6, 12b) (Glen 1992; Glen et al 1996). Effects of the Kanimblan Orogeny extend westward into the Delamerian Orogen: in the Koonenberry area, the Bancannia Trough was inverted and there was sinistral reverse movement on the NWtrending Koonenberry Fault that produced folding and thrusting in the Lambie facies rocks in the footwall (Sharp & Buckley 2003; Neef 2004). In the Thomson Orogen, deposition in the Adavale Basin was terminated by deformation in the mid-Carboniferous (Murray 1994). North-south shortening across the suture between the Lachlan and Thomson orogens resulted in basement rocks thrust south over the northern margin of the Paka Tank Trough (Fig. 12b) and in inversion of the trough itself (Alder 1999).

Deformation phase. The Kanimblan Orogeny, the last regional deformation to affect the Lachlan Orogen, is dated at c. 340 Ma. East-west Carboniferous shortening (with lesser strike-slip movement) was intense in the Eastern subprovince and led to out-of-sequence thrusting (Glen, unpublished data), with formation of high- and low-strain zones that are best recognized by deformation of the Late Devonian Lambie facies rocks. High-strain areas are reflected by formation of map-scale, narrow, elongate zones of Late Devonian strata (Fig. 12b) that occupy the cores of tight synclines (Powell 19846), commonly occupying footwalls of major N-S thrusts (Glen 1992). N-S D2 folds, faults and cleavage in the Hill End Trough are also Carboniferous in age and reflect a second phase of basin inversion by thickskinned reverse faults (Vassallo et al 2003). This intense Carboniferous deformation and

The Hunter-Bowen supercycle records the Middle Devonian to Triassic convergent margin development of East Gondwana that is expressed by the evolution of the New England Orogen and the Bowen-Gunnedah-Sydney basin system to the west. In the New England Orogen, this supercycle is built on earlier elements of the Lachlan and Tabberabberan cycles. From west to east, elements of this classical margin orogen include an arc, a forearc basin and subduction complexes together with accreted terranes (Leitch 1974; Day et al 1978; Cawood 1982; Cawood & Leitch 1985; Leitch & Scheibner 1987; Scheibner 1989; Aitchison & Flood 1992) (Figs 12a, 13a). In this paper, this Devonian-Triassic history is divided into four cycles, but the changeover between cycles 1 and 2 is not clear-cut everywhere and further work may result in the two cycles being combined.

Post-collisional phase. In the Eastern subprovince of the Lachlan Orogen New South Wales, the Kanimblan Orogeny was followed by intrusion of high-level I-type granitoids (Pogson & Watkins 1998; Shaw & Flood 1993) generated by melting of underlying Ordovician volcanic rocks (Watkins 19980) (Fig. 6). Hunter-Bowen supercycle

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Fig. 13. Map of relations in the New England Orogen: (a) deformed state geology (note the triple orocline); (b) undeformed New England Orogen, after Cawood & Leitch (1985). Late Carboniferous-Early Permian palaeogeography derived from restoring the orocline.

THE TASMANIDES OF EASTERN AUSTRALIA •

•

•

•

Cycle 1 (Late Devonian): characterized by an east-facing Late Devonian continental or intra-oceanic arc dominated by intermediate volcanism, a forearc basin (Yarrol Trough in north, Tamworth Trough in south) bounded to the east by the Yarrol Fault and Peel-Manning Fault System, and subduction complexes with accreted terranes. A possible back-arc basin occurs in the north (Drummond Basin) but not in the south (Fig. 12a). Cycle 2 (Carboniferous): characterized by an east-facing Carboniferous Andean-type arc dominated by felsic ignimbrites and granites in the west, passing eastwards into the Carboniferous part of the Yarrol and Tamworth troughs, and then into Carboniferous subduction complexes and accreted terranes. Cycle 3 (Early Permian): characterized by crustal extension associated with inception of the Bowen-Gunnedah-Sydney basin system and initial phase of the HunterBowen Orogeny. Cycle 4 (Permian to Triassic): characterized by arc magmatism, the foreland basin stage of development of the Bowen-GunnedahSydney basin system, widespread felsic igneous activity in the North Queensland Orogen to the west and terminated by the main phase of the Hunter-Bowen Orogeny, which reflects accretion of an intra-oceanic arc.

Veevers (2000d) has discussed cycles 3 and 4 in detail which he subdivided into seven stages.

Cycle 1: Late Devonian convergence In Queensland, the rhyolitic to basaltic Drummond Basin, developed on crust of the Thomson Orogen, is thought to be a back-arc basin to the New England Orogen (Day et al 1978; Murray 1986) (Fig. 4c). Note that the partly coeval development of the North Queensland Orogen to the north is interpreted to be part of the Tabberabberan cycle (Fig. 4c). This basin was initiated in the Late Devonian (Famennian) in the north but the Carboniferous (Tournaisian) in the south (Henderson et al. 1998), by NE-SW extension (Johnson & Henderson 1991) (Figs 15a, 16a). Volcanism and granite intrusion in the Anakie Inlier further west reflect coeval heating of the Thomson crust (Henderson etal 1998) (Fig. 4c). Wood & Lister (2004) suggested that the two areas were linked as lower and upper plates of an extensional system.

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The Late Devonian arc in the northern New England Orogen is best represented by the Campwyn Volcanics that were deposited on its eastern flank (Fergusson et al. 1994), although Bryan et al. (20030) have suggested these rocks are rift related. In the southern New England Orogen, an arc along the inboard part of the orogen is not preserved (Fig. 12c), but is inferred from large Late Devonian olistostromal blocks of andesitic volcanic rocks in the inboard part of the forearc basin, the Tamworth Trough (Brown 1987). This Baldwin Arc developed on oceanic or thin continental crust (Mollis 1988). The forearc basin to the east, the Tamworth Trough, was filled by marine strata that become finer grained and of deeper-water character eastwards, towards the Peel-Manning Fault System that marks the preserved outboard edge of the basin. Late Devonian trough-fill consists of mudstone packets that alternate with packets of sandstone and conglomerate that thin to the east. In the northern New England Orogen, the Yarrol Trough is the equivalent forearc basin (Figs 12a, 13a). Late Devonian strata consist of volcaniclastic sandstone and conglomerate, derived from the andesitic arc to the west, interbedded with sediments and some limestone (Yarrol Project Team 1997; Fordham et al. 1998). Minor andesitic lavas, dacites and rhyolitic ignimbrites are present. Bryan et al. (2001) suggested recently that the Yarrol Trough formed as a back-arc basin to an arc to the east, in an area now occupied by accreted terranes. While this model is based on arguments that volcanic rocks in the trough have a back-arc basin geochemical character and were derived from west- rather than eastflowing currents, on a more regional scale it implies a major decoupling between the northern and southern parts of the New England Orogen. Murray et al. (2003) rejected this back-arc model. Bryan et al. (2003&) reaffirmed their back-arc interpretation and Leitch et al. (2003) supported the forearc interpretation. East of the Peel-Manning Fault System in the southern New England Orogen, Silurian-Early Devonian subduction continued into the Late Devonian without interruption. As a result, the upper parts of thrust slices (Fig. 11 a) contain Famennian (Late Devonian) volcanolithic sandstone beds (with olistoliths), suggesting deposition not too remote from the subductionrelated arc (Aitchison et al. 19920). Ultramafic units of tholeiitic affinity along, and east of, the Yarras Fault were interpreted as a fragment of

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an accreted Middle to Late Devonian (374 Ma) arc by Aitchison et al (1994). In the northern New England Orogen, Late Devonian conodonts in the Beenleigh terrane indicate subduction in the Late Devonian (Aitchison 1988). Basalt dykes and faultbounded blocks in the Neoproterozoic Marlborough Ophiolite have a Middle Devonian Sm-Nd isochron (380 Ma) and trace element data suggestive of an intra-oceanic island arc (Bruce & Niu 20006). These relations suggest that the arc was probably built on a Neoproterozoic oceanic crust, lay only a short distance offshore (C. Murray & P. Blake unpublished geochemical data) and was accreted to Gondwana by the Early Permian (Bruce & Niu 20006). Cycle 2: Carboniferous convergence In Queensland, the back-arc Drummond Basin continued to evolve, receiving basementderived, Early Visean quartz-rich detritus (Stage 2 of Johnson & Henderson 1991; Henderson et al. 1998) followed by Mid-Late Visean ?arc-derived volcaniclastic rocks (Stage 3 of Johnson & Henderson 1991) (Fig. 4c). In the northern New England Orogen, the Carboniferous continental margin arc is represented by granitic and mafic to silicic rocks of the Connors and Auburn arches arc (Day et al 1978, but not Bryan et al 20036). These two arches lie along-strike, separated by the Gogango Overfolded Zone (Fig. 13a), which represents a part of the Permian Bo wen Basin succession strongly deformed and thrust westwards in the Late Permian to Early Triassic (Fergusson 1991; Fielding et al 1997). Although igneous activity in both arches commenced in the Tournaisian (represented by c. 350 Ma granites), the main pulse of granite formation was in the Namurian stage, from c. 324 Ma until 313 Ma (predating uplift) in the Auburn Arch, and c. 316-305 Ma in the Connors Arch, where uplift was only minor (Murray 2003). Bryan et al (2001) suggested that all these granites were extensional in character. Murray (2003), however, indicated that the older granites in the Auburn Arch and the older part of the Connors Arch are subductionrelated, and that younger granites with large volumes of rhyolitic to dacitic ignimbrites and local andesite lavas in the Connors Arch spanned the changeover from subduction to the beginning of extension, around 305 Ma, the Carboniferous-Permian boundary (Murray 2003). The Early Tournaisian-Visean cycle of fill of

the Yarrol Trough forearc basin (back-arc basin of Bryan et al (2001, 20036) was marked by the presence of felsic, rather than the intermediate, volcanic detritus of the first cycle, and by the presence of finer-grained sediments consisting mainly of fine-grained sandstone, siltstone and oolitic limestone (Yarrol Project Team 1997; Fordham et al 1998). Granite clasts appear in the Yarrol Trough around the ViseanNamurian boundary (c. 327 Ma) and reflect rapid exhumation of the arc to the west (Murray 2003). The Yarrol Trough passes outboard across the Yarrol Fault System (Figs 12,13) into ultramafic to mafic rocks of the Marlborough terrane in the north and into subduction complex rocks in the south (Fig. 13a). Subduction complex rocks in the northern New England Orogen comprise the North and South D'Aguilar and Beenleigh blocks in the south and the Wandilla and Shoalwater terranes in the north (Fig. 13a). The Beenleigh Block consists of greywacke, argillite and Early Carboniferous chert (Aitchison 1988), which are organized into packets of coherent turbidites separated by zones of broken formation (Smith 1999). Similar strata occur in the South D'Aguilar Block: both are interpreted as forming in upper levels of an accretionary complex (Holcombe et al 1997 b). The North D'Aguilar Block, in contrast, is a composite terrane, containing high-level accretionary complex rocks, as well as ophiolitic components of an accretionary complex that were subducted to more than 18 km before 315 Ma, and were then exhumed in the lower plate of a latest Carboniferous, low-angle normal fault (Little et al 1992; 1995; Holcombe et al 19970). Further north, the Early Carboniferous Wandilla terrane consists mainly of bedded and melanged deep-water sandstone and mudstone. The volcaniclastic character of sandstone beds (with clasts of basalt, andesite and rhyolite) together with rare ash-fall tuffs suggests a provenance from the coeval magmatic arc and its basement to the west (Leitch et al 2003). In contrast, the outboard Shoalwater terrane of presumed late Early Carboniferous age is dominated by cratonderived quartz-rich sandstone inferred to have been derived by longitudinal transport from the North Queensland Orogen (terminology of this paper) (Leitch et al 2003). In the southern New England Orogen, a c. 400 km belt of NW-trending Carboniferous volcanic rocks represent the outboard parts of the continental margin Currabubula arc that was developed on crust of the Lachlan Orogen (McPhie 1987) (Figs 12c, 13a). Volcanism

THE TASMANIDES OF EASTERN AUSTRALIA reached a peak in the Visean (342-327 Ma), after a minor pulse of volcanism in the Tournaisian (Fig. 4b) and was largely rhyolitic in the southern c. E-W-trending part of the arc, but mixed andesitic-rhyolitic further north. In both areas, volcanism continued locally into the early Stephanian (c. 303 Ma), but was rhyoliticdacitic in both areas (Roberts 1995; Roberts et al 2003; 2004; J. Roberts unpublished data) (Fig. 4b). The main volcanic products are lowK calc-alkaline rhyolitic ignimbrites with lesser calc-alkaline andesites (McPhie 1987; Roberts et al 2003) that show LREE enrichment, moderately negative Eu anomalies, HFSE depletion and LILE enrichment (Jenkins et al 2002). These features are all indicative of subduction zone magmatism. Highly radiogenic Sr suggests incorporation of continental crust into the melts (Jenkins et al 2002). These Carboniferous arc rocks are similar in age and chemistry to c. 340-320 Ma I-type granites in the Eastern subprovince of the Lachlan Orogen that are post-collisional to the Kanimblan cycle (Figs 6, 12a). Shaw & Flood (1993) suggested that the granites were deeper levels of a single arc that had been displaced and exhumed, whereas Jenkins et al (2002) suggested that the granites represented a second arc. An alternative view is that the two rock packages are unrelated and the arc-like signature of the Carboniferous granites reflects partial melting of the underlying Ordovician arc volcanic rocks (Watkins 19980; Glen 1998). The Carboniferous stage of the forearc basin development is marked by major segmentation of the Tamworth Trough by cross-faults, represented in some cases by submarine channel margins (Roberts & Engel 1987; Roberts & Glen unpublished). Granite clasts in the Early Visean (c. 340 Ma) and the base of Namurian (c. 327 Ma) were sourced from the exhumed arc to the west (J. Roberts pers. comm. 2004). Deposition ceased at c. 305 Ma (Roberts & Geeve 1999). Most of the accretionary complex in the New England Orogen is Carboniferous in age (Figs 4b, 4c), with ages originally based on linking oolitic detritus back to Early Carboniferous oolitic limestones in the forearc basin (Murray 1997). This complex was called the Texas Terrane by Cawood & Leitch (1985) and corresponds largely to the Anaiwan Terrane of Aitchison et al (19920). High-level, frontal accretion is recorded by imbrication of turbidite sequences and formation of melanges of Early and Late Carboniferous age (Fergusson 1985). Aitchison et al (19920) showed that the typical stratigraphy of a fault slice consisted of basal

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mafic volcanic rocks, Late Devonian cherts, Tournaisian tuffaceous siltstone and chert interbedded with sandstone passing up into the dominant Tournaisian-Visean volcanogenic sandstone (Fig. lib). The presence of these volcanogenic sandstones indicates that the younger Carboniferous part of the accretionary prism developed relatively close to the continental margin arc (Aitchison et al 19920). Subduction underplating is recorded by the presence of riebeckite-bearing terranes, now exhumed along major faults (Dirks et al 1992). K-Ar riebeckite ages of 319 Ma and 312 Ma reflect high-pressure subduction metamorphism (Watanabe et al 1988). Deformation. In the southern New England Orogen, the cessation of subduction was followed by a major c. 311-300 Ma event that culminated in uplift and emplacement of early granites at 302 Ma and 300 Ma (Hillgrove suite and Wongwibinda complex, respectively: Dirks et al 1992; Landenberger et al 1995). The early fabrics in these metamorphic complexes have been interpreted as contractional (Collins et al 1993). Approximately synchronous cessation of deposition in the forearc basin at c. 305 Ma (Roberts & Geeve 1999) was followed by major strike-slip faulting and anticlockwise rotation of crustal blocks. Three blocks in the southern part of the Tamworth Trough underwent Late Carboniferous, palaeomagneticallydetermined, anticlockwise rotation (Geeve et al 2002), although the contribution of thrusts to vertical-axis rotation has not yet been determined. A fourth block, the Hastings Block, underwent c. 250 km of left-lateral translation to the NE as well as anticlockwise rotation (Scheibner 1976; Cawood 1982; Cawood & Leitch 1985; Schmidt et al 1994; Collins et al 1993; Roberts & Geeve 1999) (Figs 12c, 13a). These rotated blocks were then folded, cleaved and faulted in the Late Carboniferous, before deposition of disconformably-overlying Early Permian strata and intrusion of the Barrington Tops Granite (which has a U-Pb zircon date of 281 Ma, Kimbrough et al 1993). Carboniferous and Devonian rocks in the northern part of the Tamworth Trough were telescoped by folding, variable cleavage development and thin-skinned thrusting into a west-vergent foreland fold thrust belt (the Tamworth Belt) that defines the external part of the New England Orogen (Glen & Brown 1995; Woodward 1995). This deformation was probably also Late Carboniferous, since Allan & Leitch (1990) recorded 4000m of prePermian uplift west of the Peel Fault System and

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pre-Permian low-grade regional metamorphism, deformation and erosion to the east. The similarity of detritus in Early Permian stratigraphic units across the Peel-Manning Fault System suggests that the forearc basin and subduction complex had been largely juxtaposed by the Early Permian (Allan & Leitch 1990; Dirks et al 1992). In Queensland, uplift in the arc is Late Carboniferous, around 305 Ma (Hutton et al 1999), and postdates mid-Late Carboniferous inversion of the back-arc Drummond Basin that produced folds and thick-skinned oblique thrusts with N-S trends in the south but which wrap around the Anakie Inlier in the north (Johnson & Henderson 1991). Deformed rocks are overlain unconformably by Late Carboniferous volcanic rocks and granites (Caritat & Braun 1992) .The cause of this Late Carboniferous deformation is uncertain: it could reflect 'hiccups' in steady-state subduction, regional deformation (e.g. Allan & Leitch 1990), or it could be extension-related, such as inferred in the northern New England Orogen.

Extension. In the New England Orogen, latest Carboniferous to Early Permian extension is recorded by (1) emplacement of granites and serpentinites into, and formation of low-angle extensional faults in, the former accretionary prism; and (2) formation of rift sedimentary basins.

sedimentary serpentinites in the Permian (Cross et al 1987; Allan & Leitch 1990). It is also consistent with Early Permian (286 and 279 Ma) K-Ar ages of nephrite from a splay fault of the Peel Fault System (Lanphere & Hockley 1976). (2) Early Permian rift basins are widespread (Fig. 14a), and were developed above deformed rocks of the forearc basins, the accretionary complexes, as well as on top of rocks of older basement (the Lachlan, Thomson and North Queensland orogens) to the west. Early Permian basins above accretionary terranes in the northern New England Orogen have been described by Sliwa et al (1993). In the southern New England Orogen, these basins are either structural remnants of one large basin (Barnard Basin of Leitch 1988) or formed as separate narrow fault-bounded transtensional basins along major faults such as the Peel-Manning Fault System (Korsch 1982; Flood & Aitchison 1992). Vickers (1999) gave one example of a local basin formed during dextral overall transpression. The Gloucester Basin in the southern New England Orogen developed above deformed rocks of the Tamworth Belt. Basal units consist of basalts and andesites with some rhyolites (the c. 275 Ma, Alum Mt Volcanics) that have MORB-like affinities with some input either from a crustal or subduction modified source (Roberts et al 1996).

(1) Granites were intruded into Devonian and Carboniferous accreted terranes in the Late Carboniferous (Fig. 13a). In the North D'Aguillar Block, in the northern New England Orogen, these granites were emplaced at c. 306 Ma during regional extension that led to generation of a major low-angle normal extensional fault. This fault separates multiply-deformed, blueschist-greenschist grade rocks in the lower plate from prehnite-pumpellyite grade, high-level accreted rocks in the upper plate (Little et al 1992; 1995). In the southern New England Orogen, the Tia Complex was uplifted at c. 311 Ma (Dirks et al 1992) and intruded by granite at 302 Ma. Uplift was thought to be contractional (Collins et al 1993) but could be extensional in nature. The emplacement of serpentinite along the Peel-Manning Fault System in the Early Permian (Leitch 1969), during regional extension is recorded by the presence of clasts of serpentinite and

Extension of crust of the North Queensland Orogen is reflected by generation of small Early Permian rift basins (Davis & Henderson 1999) and emplacement of widespread felsic volcanic rocks and granites (Bain & Draper 1997). Extension of Lachlan and Thomson crust west of the New England Orogen is represented by the Early Permian inception and growth of the Sydney, Gunnedah and Bowen basins as extensional, or more likely, transtensional rifts that are floored by volcanics or underlain by intrusive rocks such as represented by the regional elongate Meandarra gravity ridge (Qureshi 1984; Murray et al 1989). A combination of regional sag (post-rift subsidence) and cessation of loading in the late Early Permian led to establishment of marine conditions that prevailed in more distal parts of the basin system. Early deformation in the western Bowen Basin produced a gentle inversion unconformity (Korsch et al 1998). Rift volcanic rocks at the base of the Sydney Basin include the 292 Ma (Rb-Sr) Rylstone

Cycle 3: Early Permian

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63

Volcanics near the western margin (Shaw et al 1989) and the 270 Ma Werrie Basalt (with remnant magnetism) that lies at the base of the NE Sydney Basin as well as underlying most of the Gunnedah Basin (Leitch 1993; Caprarelli & Leitch 2001). This basalt has a mixed OIB and subduction-related signature, possibly reflecting derivation from melting of mantle enriched during Carboniferous subduction (Caprarelli & Leitch 2001). High level c. 270 Ma igneous rocks intrusive into the Werrie Basalt have subalkaline to alkaline chemistry. They were derived from depleted mantle with some crustal contamination and carry 340 Ma and 300 Ma detrital zircons (Teale et al 1999). Silicic flows and ignimbrites of the Boggabri Volcanics (Leitch 1993) outcrop as a curved NW belt near the eastern margin of the Gunnedah Basin and extend westwards at depth under most of the basin (Leitch 1993; Tadros 1993; Beckett et al. 1995). They have a predominant MORB-like geochemistry, but with a subduction signature inherited from Carboniferous subduction during Early Permian extension (Brownlow 1999; Brownlow & Arculus 1999). Volcanic rocks at the base of the Bo wen Basin in Queensland include the Rookwood Volcanics, bimodal rocks of the Early Permian Lizzie Creek Volcanics and the dominantly intermediate and felsic rocks of the Late Carboniferous (Stephanian) to early Permian Camboon Volcanics (Hutton et al 1999; Fielding et al 2001). Interpretations of the Early Permian extension in the Hunter-Bo wen supercycle include subduction of a spreading ridge (Murray et al 1987), slab break-off (Caprarelli & Leitch 1998), and changes in the Gondwana-proto-Pacific plate boundary configuration from retreating to advancing (Jenkins et al (2002) (Fig. 14b). Orocline formation. The southern New England Orogen is dominated structurally by three orogen-scale folds or oroclines in accretionary complex rocks; the southern one is also outlined by the curve in the Peel-Manning Fault System (e.g. Korsch & Harrington 1987; Murray 1997)

Fig. 14. (a-b) Hunter-Bowen supercycle: (a) cycle 3 showing igneous activity and Bowen, Gunnedah and Sydney basins that developed from rift (transtensional) basins into foreland basins; (b) Permian to Triassic advancing and retreating margins (after Jenkins et al 2002); (c) Mesozoic rifting and basin formation.

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(Figs 12c, 13). The southern (Manning) orocline began to form in the Late Carboniferous with rotation and left-lateral translation of blocks in the Tamworth Trough described above, and formation of radiating faults (e.g. Korsch & Harrington 1987). It continued into the Permian, since Permian basins were caught up in a N-S deformation at c. 268-267 Ma (Roberts & Geeve 1999). This sense of faulting is consistent with left-lateral movement on the Peel-Manning Fault System to the northwest, where it is recorded by oblique S2 cleavage in the eastern part of the Tamworth Belt, transecting early Dl folds (e.g. Cao & Durney 1993; Glen unpublished data), and by strike-slip fabrics in schistose serpentinite east of the Peel Fault System east of Manilla (Glen & Brown 1995). The northern (Texas) orocline also deformed Early Permian sediments (Lennox & Flood 1997) and Korsch & Harrington (1987) showed that this involved right-lateral strikeslip on the eastern limb. Together the northern and southern oroclines define a long doublyplunging regional (500 km) antiform. The eastern (or Coffs Harbour) orocline formed when accretionary complex rocks of the Anaiwan/Texas Terrane to the north were translated southwards, against the northwardsmoving Hastings Block, also resulting in east-west structures in the Early Permian Nambucca Block (Johnston et al 2002). Subsequent east-west shortening of these rocks at 266-258 Ma is dated by uplift on bounding faults to the west (Dirks et al 1992; Landenberger et al. 1995). Formation of these oroclines by synchronous yet oppositely directed strike-slip faulting seems to be a paradox. A possible explanation is suggested here, involving major out-of-plane movement, with the almost complete closure of accretionary complex rocks representing a section through a steeply plunging noncylindrical antiform. In this model, the core of the fold underwent sub-vertical extension, bringing high-grade accretionary complex rocks to the surface. This uplift corresponds to the 310-300 Ma S3/L3 event of Dirks et al (1992). The large scale of this structure (c. 500 km) is consistent with uplift of 10-15 km (Dirks et al 1992). As part of this extension, outer, lowgrade rocks of the Tamworth Belt underwent opposite-directed translation and rotation, with radiating faults focused on the southern orocline. Such a 'sheath fold' model invites comparison with the core complexes described from Queensland by Little et al (1992).

Cycle 4: Late Early Permian-Triassic Convergence. Renewed convergence in the late Early Permian to Triassic was accompanied by the extrusion of widespread ignimbrite sheets and emplacement of major I-type Late Permian-Late Triassic (255-222 Ma) granites of the New England Batholith. Most of these granites lie in the southern New England Orogen (Figs 13a, 14). Shaw et al (1991) recognized three groups of granites and associated volcanic rocks, with most emplaced between 253 Ma and 244 Ma, some in NNE-trending graben or rifts. Some granites extend into the northern New England Orogen, where they range in age between 270 Ma and 205 Ma (Murray 2003). These granites are high-level and mainly I-type, although some A-types are present, as are zoned granites and layered gabbros (Murray 2003). Cycle 4 granites have a close geochemical similarity with Mesozoic granites of the North American Cordillera (Cawood 1984; Chappell 1994; Bryant et al 1997), indicating that they are subduction-related, especially those of the Clarence Supersuite in the east (Bryant et al 1997), which extend into the northern New England Orogen. Gust et al (1993) also argued for a subduction-related origin, but suggested that there was a transition to extension in the Late Triassic due to slab rollback. Chappell (1994) noted that the I-type Moonbi suite in the southern New England Orogen was K-rich, like the Kanimblan post-collisional Bathurst granites in the eastern Lachlan Orogen. Could they also be melts of Ordovician volcanic rocks in the under-thrust Lachlan Orogen (see below)? The Gympie Terrane (or Gympie Province, Cranfield et al 1997) (Figs 13a, 14b) in the outboard part of the northern New England Orogen is a key component of cycle 4 (Harrington 1974). This terrane is in fault contact with the Carboniferous accretionary prism to the west and consists of Early-mid-Permian volcanic rocks passing up into Late Permian limestone and then into Early Triassic continental, then marine, units (Cranfield et al 1997; Sivell & McCulloch 2001). This terrane is interpreted as a primitive oceanic arc associated with enriched subcontinental lithosphere, subsequently rifted in the Early Permian. Proximity to the Gondwana margin is debatable (Harrington 1983; Cranfield et al 1997; Sivell & McCulloch 2001; Holcombe et al 1997ft). There are similarities with the offshore South Island terrane (Fig. 13a), which consists of mafic and ultramafic rocks with a maximum age of

THE TASMANIDES OF EASTERN AUSTRALIA c. 277 Ma (Early Permian) and which formed in an intra-oceanic arc (Bruce & Niu 20000). Similar-age volcanic elements in the Berserker terrane (erupted in a back-arc or intra-arc setting, Crouch 1999) and in the South Island terrane suggest relative proximity (Bruce & Niu 20000). West of the New England Orogen, cycle 4 was reflected by conversion of rift basins in the Bowen-Gunnedah-Sydney basin system to foreland basins (Figs 13a, 14). Foreland loading by the southern New England Orogen first commenced in the late Early Permian and is recorded by deposition of the oldest coal measures (Greta Coal Measures) mixed with volcanic detritus in the NE part of the Sydney Basin. Syn-sedimentary thrusting in the NE Sydney Basin led to the formation of growth anticlines and westward-propagating thrust fronts (Glen & Beckett 1997) and produced southward-draining axial palaeodrainage ahead of the thrust front. In the Gunnedah Basin, equivalent deposits in the eastern Maules Creek sub-basin were derived from basal volcanic rocks upthrust to the west (Beckett et al 1995). Minor Late Permian (c. 252 Ma) anticlinal growth further west is reflected by flatlying earliest Triassic rocks above dipping Permian reflectors (Tadros 1993). In the Bowen Basin, thrusting produced the Gogango Overfolded Zone (Fergusson 1991; Fielding et al 2001). By the end of the Late Permian (Tartarian), all three basins had been converted to coalbearing foreland basins, fed in pulses from first cycle volcanic detritus and uplifted detritus (e.g. jaspers from the accretionary complex) from the New England Orogen in the north and east (Fielding et al. 2001), where crustal loading was synchronous with active volcanism and granite formation. Tuffs in Late Permian coal measures reflect coeval volcanism in the southern New England Orogen (Brownlow 1979; Shaw et al. 1991), with the oldest tuffs dated at c. 264 and c. 265 Ma. Mixed volcanic and quartz-rich detritus in some stratigraphic units, especially in the Gunnedah Basin, reflects input from both the New England Orogen and a westwardreceding crustal bulge in the older Lachlan Orogen to the west. A major environmental change just above the Permian-Triassic boundary (c. 250 Ma) coincided with the end of tuff and coal deposition and the onset of red-bed oxidized fluviatile sedimentation of the Rewan Group in the Bowen Basin and the Narrabeen Group in the Sydney Basin (Veevervood). The last unit especially was derived from the New England

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Orogen and was synchronous with major periods of volcanism and granite emplacement. The thrust front reached the eastern margin of the Bowen Basin during this event. This Early Triassic phase of basin filling lasted only c. 14 million years (250-244 Ma) and was followed by relative uplift in the craton from 244 Ma to 235 Ma (Fielding et al. 2001), leading to the shedding of large amounts of craton-derived detritus into the Bowen, Gunnedah and Sydney basins to form the Clematis and lower Napperby formations and the Hawkesbury Sandstone, respectively. Final filling of these basins (235-230 Ma) was dominated again by detritus shed from the New England Orogen (which included coeval tephra deposits, Shaw et al 1991), although the finer-grained nature of sediments, commonly lacustrine, point to lesser topographic expression and reduced thrusting. Collision. In the New England Orogen, cycle 4 was terminated by the collision of the Gympie Terrane/Province with Gondwana in the Early to Middle Triassic (Sivell & Arnold 1999) or Middle to Late Triassic (Cranfield et al 1997). This deformation is inferred to have caused the Bowen phase of the Hunter-Bowen Orogeny, a major crustal loading event in the New England Orogen (Cranfield et al 1997). The Sydney, Gunnedah and Bowen basins became converted into fold-thrust belts as the deformation fronts migrated westwards. An Early Triassic contractional deformation is recorded from the Bowen Basin and a Middle Triassic event from the Drummond Basin to the west (Johnson & Henderson 1991), before the terminal Late Triassic deformation at c. 233 Ma (Korsch et al. 1998), which also involved rocks of the Tamworth Belt being thrust westward over Middle Triassic and unconformably-overlying Early Jurassic rocks of the Surat Basin (Wartenberg et al 1999). In central Queensland, widespread thrusting and open folding in the central and western parts of the Bowen Basin suggest greater amounts of crustal shortening than in the Gunnedah and Sydney basins. Indeed, Malone (1964) and Fielding et al (1997) suggested that the basin extended up to 200 km to the east, reaching the Queensland coast before deformation, with only small relics preserved on the surface and in the subsurface. In contrast, the Gunnedah Basin underwent less shortening below the bounding 15° E-dipping Mooki Thrust, so that only the eastern parts (the Maules Creek Sub-basin and Boggabri Ridge) in the footwall underwent significant thrust-related deformation (Beckett

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et al 1995). The western part of the Gunnedah Basin still retains elements of its early rift geometry, being divided into north-south blocks by major cross-faults (Tadros 1993). Further south, the northeastern margin of the Sydney Basin lies in the lower plate of the major east-dipping Hunter Thrust, which has a complex history of greater shortening in the north (Glen & Beckett 1997; Glen & Roberts unpublished data). Thrusting in the offshore part of the New England Orogen was westdirected and is probably responsible for formation of major N-S folds and blind faults in the main part of the Sydney Basin. Hunter-Bo wen cycle 4 also affected basement west of the New England Orogen, with the emplacement of large amounts of granite and felsic volcanic rocks in the North Queensland Orogen (Bain & Draper 1997). In the North Queensland Orogen, the Hodgkinson subprovince that underwent D3 vertical shortening during cycle 3 extension (Davis & Henderson 1999) underwent D4 horizontal crustal shortening in cycle 4, which is best constrained by the syn-deformational intrusion of S-type granites (dated at 280-265 Ma, Davis et al. 2002). Subsequent formation of small Early-Late Permian coal basins were deformed in the Late Permian-Early Triassic by the Bo wen phase of the Hunter-Bo wen Orogeny (Bain & Draper 1997). Mesozoic rifting After the end of the Hunter-Bo wen supercycle, the plate margin was relocated to the east. As a consequence of the opening of the Tasman Sea, beginning at c. 90 Ma, the record of this younger Mesozoic interaction between Gondwana and the proto-Pacific plate is preserved in the islands of the southwest Pacific and in New Zealand. The Australian sector of Gondwana underwent Late Triassic rift-related volcanism and slow subsidence that led to the formation of widespread Jurassic-Cretaceous shallow-water basins. The prominent Jurassic Whitsunday Rift developed along the eastern margin of the continent from c. 135 Ma to 90 Ma (Bryan et al. 1997) (Fig. 14c). Part 3: Discussion In this third part of the paper, some of the more keenly-debated parts of the history of Delamerian and Lachlan cycles are discussed from the point of view of both the author as well as other models.

Significance of cycles Recognition of cycles in the Tasmanides emphasizes the wide-scale nature of stages, both extensional/magmatic/sedimentary and contractional/strike-slip, in the evolution of the Australian sector of East Gondwana. By drawing attention to linkages between orogenic belts, the recognition of cycles forces a widerscale look at the growth of the Tasmanides. Without being too doctrinaire in this approach, recognizing the possibility of diachronous events and lack of correlations, several points can be made. Recognition of elements of the Delamerian cycle in the New England Orogen leads to questions on what happened in the outboard part of the Tasmanides during Rodinia break-up, and when subduction began (e.g. Cawood & Leitch 1985; Cawood 2002; Crawford et al 20030). The recognition of Lachlan cycles in the Delamerian Orogen suggests that cratonized crust and upper mantle can still respond to deformational and heating events generated at the plate margin further outboard. The application of Lachlan cycles to the North Queensland Orogen implies broad similarities in plate events, despite the fact that, as discussed below, one area underwent major back-arc spreading not evident in the other. If these two orogens were part of the same, albeit segmented, plate margin, where is the intervening part of this convergent margin? Was it overthrust by the outboard (Delamerian) part of the Thomson Orogen? Although Devonian elements of the Hunter-Bo wen supercycle are interpreted as having affected the Thomson and North Queensland orogens, it is not clear how to separate them from elements of the Tabberabberan cycle. Finally, the presence of Delamerian to Tabberabberan cycles in the New England Orogen bears on the question of how allochthonous that orogen was in relation to the Gondwana margin. It suggests formation not too far from the Lachlan Orogen in the Ordovician-Late Devonian, an issue addressed most recently from palaeomagnetic data by Klootwijk (2002). Rifting in the Delamerian cycle The rift history of the Delamerian cycle represents the response to the break-up of Rodinia and the separation of Laurentia (e.g. Dalziel et al 1994; Powell et al 1994) although Direen & Crawford (20030) preferred a model in which microcontinental ribbons were calved off. Most of our knowledge of the Delamerian cycle conies from the inboard parts of the

THE TASMANIDES OF EASTERN AUSTRALIA Tasmanides. Several issues can be addressed from these areas, such as the timing of rift-drift transition, the geometry of rifting and the relationship between Tasmania and the mainland before c. 700 Ma. However, clues as to what happened further outboard during the c. 200 millions of years of rifting are less easy to find. Neoproterozoic rift volcanic rocks in the Anakie Inlier lie thousands of kilometres east of the Gondwana margin, as now accepted, and indicate that either our ideas of the shape of the East Gondwana margin were wrong or that its current location reflects subsequent rifting to the west. Similarly, the Marlborough Ophiolite in the New England Orogen indicates that seafloor spreading was underway in the Neoproterozoic. This is consistent with the presence of old lithosphere under the southern New England Orogen. This lithosphere could have rifted away from cratonic Australia, to be subducted early in Tasmanides history, or was brought in during subsequent Palaeozoic to Mesozoic subduction or strike-slip faulting. Part of the answer to this question may come from western Tasmania, which shows an early history not recorded on the mainland (e.g. significant rift tholeiites >780 Ma) and which appears to have lain further outboard than the Adelaide Rift Complex before deposition of c. 700 Ma Sturtian tillites (e.g. Elliott & Gray 1992; Powell et al 1994). Berry et al (2001) showed that the detrital zircon pattern in Tasmanian Cambrian rocks differed from those of the Kanmantoo Group on the mainland and suggested that western Tasmania rifted from Australia and remained as a separate terrane or promontory until the Delamerian Orogeny. Returning to the inboard part of Gondwana, the old edge of Rodinia can be reconstructed partially as a steep east-dipping normal fault (the Torrens Hinge Zone) on the western side of the Adelaide Rift Complex in the southern part of the Tasmanides, because of the limited reactivation of this zone by subsequent thrusting. Rifting occurred over 300 million years (827-527 Ma), and was taken up mainly by rift cycle 1 of this paper. It is generally agreed that the Australian craton was part of the Rodinia supercontinent and that rifting marked the opening of the proto-Pacific Ocean and separation of Laurentia (as well as the South China Block, Li et al. 1995) from Gondwana. Powell et al (19935) used palaeomagnetic data to suggest that this separation began after 725 Ma; newer data (Wingate & Giddings 2000) indicated that rifting occurred earlier, by c. 755 Ma. If so, this event was not only amagmatic, it was not reflected in the

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Adelaide Rift Complex for another 50 million years. The Tapley Hill Shale (post-Sturtian tillite) is the first stratigraphic unit that was deposited right across the rift system and interpreted by Powell et al. (1994) as reflecting the rift-drift transition. If separation occurred by 755 Ma the rift margin must have lain further east. What happened before 600-580 Ma, a period of major igneous rifting with production of continental rift alkaline-tholeiitic basalt and picritic volcanism, documented by Crawford and co-workers (e.g. Direen & Crawford 20030)? Was the 600-580 Ma event linked to subsequent rifting of smaller continental blocks in newly formed Gondwana, calving off ribbons, or was it the main rift event, as argued by Direen & Crawford (20030) and by Veevers et al. (1997), who suggested that the break-up of Rodinia occurred around 560 Ma. The direction of rifting in event 1 is provided by the orientation of the Gairdner Dyke Swarm and the palaeogeographical reconstructions of the Adelaide Rift Complex by Preiss (2000). Both imply a NE direction of maximum extension from c. 827 Ma, in cycle 1 events 1-2,4 (see the section on 'Delamerian cycle: extension and passive margin phase') until the Cambrian. Mixed NE and E extension characterized events 3 and 5, with east-west extension especially common in the southern part of the Adelaide Rift Complex, with the Torrens Hinge Zone initiated in cycle Ic as the western edge of rifting (Fig. 3b).

Significance of the Tasman Line The Tasman Line was originally defined by Hill (1951) as the line marking the eastern margin of outcropping Precambrian rocks in Australia. Harrington (1974) extended the line to the south, and locations of these lines and other authors' versions are shown in Figure 1. Subsequently, the concept of the Tasman Line has been broadened to represent the western boundary of Tasmanides, with the line considered to mark the place of break-up of a Mesoproterozoic supercontinent of which Australia was part (e.g. Veevers & Powell 1984; Powell et al. 1994; Scheibner & Basden 1998; Scheibner & Veevers 2000; but not Direen & Crawford 20035) (Fig. 1). While the Tasman Line coincides with the western margin of the Tasmanides in North Queensland, it is a young contractional fault that adds little hard data to the debate about supercontinent break-up geometry or kinematics. South of the Diamantina River Lineament, most authors run the Tasman Line east of, and

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around, the Curnamona craton, which includes the Palaeo-Mesoproterozoic Willyama Supergroup rocks of Olary and Broken Hill, and east of the Neoproterozoic Adelaide Rift Complex (Fig. 1). However, these differing definitions of the Tasman Line do not differentiate between Neoproterozoic fill of a rift system (and its thinned basement of older thinned Palaeo- and Mesoproterozoic rocks) and unthinned basement to the west. It has been argued above that the true boundary lies further west, along the Torrens Hinge Zone (Fig. 1) (also Mills 1992). The Tasman Line has also been recognized in Tasmania as the Tamar Fracture along the NNW-trending Tamar River, separating eastern and western Tasmania (Williams 1978). This fault has been linked to the mainland by a WNW tear fault in Bass Strait (e.g. Scheibner 1974; Veevers & Powell 1984). Because geophysical modelling suggested that the Tamar Fracture has no crustal significance (Leaman 1994), Reed et al (2002) recognized a new boundary, the west Tamar Fault, 20 km to the west, as an Early-Middle Devonian (Tabberabberan) NE-dipping thrust system (Fig. 1). As a result of these points, it is suggested here that south of the Diamantina River Lineament, the Tasman Line has little tectonic significance (see also Direen & Crawford 2003£). This is consistent with an analysis of recent earthquake SKIPPY data that found little correspondence between the Tasman Line and the eastern edge of Proterozoic Australia, except along the Palmerville Fault System (van der Hilst et al 1998). Several authors have used the zig-zag shape of the southern part of the Tasman Line that wraps east around the Curnamona craton (Fig. 1) to suggest that either NE-trending segments formed as extensional faults (Veevers & Powell 1984; Powell et al 1994) or that NW-trending segments formed as extensional faults (Powell 1998; Gibson 1998). These interpretations assume that the zig-zag in the Tasman Line around the Curnamona craton east of Broken Hill is an original feature. However, this shape may reflect later deformation that progressively plastered the Delamerian Orogen onto the older craton (Scheibner & Basden 1998; Scheibner & Veevers 2000), in the same way as the Fleurieu arc south of Adelaide was compressed onto the SW corner of the Gawler craton (most recently, Marshak & Flottman 1996). If correct, the original configuration of the southern part of the Tasman Line was approximately planar without zig-zags. In this case, the locations of the 600-580 Ma rift

volcanic rocks of Direen & Crawford (20030) would restore to an approximately linear trend.

Convergence, collision and post-collision in the Delamerian cycle The Delamerian/Tyennan Orogeny in western Tasmania and western Victoria was triggered by arc-continent collision around 510-505 Ma (Berry & Crawford 1988; Crawford 20030), followed by deformation of 500 Ma postcollisional volcanic rocks and rift basins around 495 Ma. Outboard parts of the Delamerian Orogen in western Tasmania and western Victoria preserve evidence of accreted Cambrian forearc boninitic crust - the Tasmanian mafic^ultramafic complex (Crawford & Berry 1992) and the lower part of the Stavely Volcanic Complex, respectively. These underwent post-collisional extension leading to emplacement of andesitic rocks (Crawford 20030). Using West Pacific analogues, Crawford (20030) argued that rocks of the same association in hanging walls of major thrusts in the Lachlan Orogen represent forearc crust that was distal to the collision. The actual colliding arc originally lay further east, according to the Crawford model, although several authors (Scheibner 1989; VandenBerg et al 2000; Cayley et al 2002) have suggested this accreted arc lies in western Victoria. Calc-alkaline volcanic rocks in the hanging wall of the Mt Wellington Fault are more ambiguous. They could represent part of the (largely subducted) colliding arc, but they are probably post-collisional, like the Mt Read Volcanics in western Tasmania that they resemble geochemically (Crawford et al 1996). However, that means there must have been a Delamerian/Tyennan deformation that extended into Victoria, despite the lack of recognition of an earlier 'ophiolite' obduction event (VandenBerg et al 1995; 2000; Crawford 20036). VandenBerg et al (1995, 2000) argued that these calc-alkaline rocks are windows into underlying west Tasmania crust - the Selywn Block of Cayley et al (2002). There are two other possible interpretations. The least likely accreted is that these are calc-alkaline rocks of another arc. This is effectively precluded by the similarity in geochemistry with the Mt Read Volcanics and the Stavely Volcanic Complex (Crawford et al 1996; A. Crawford pers. comm. 2004). The preferred view here is that while the andesites are geochemically post-collisional, using the arguments of Crawford, their position so far east of the Delamerian margin reflects

THE TASMANIDES OF EASTERN AUSTRALIA

post-Delamerian rifting. It is thus envisaged that pieces of a forearc crust and post-collisional volcanic rocks, accreted to the East Gondwana Orogen during and just after the terminal Cambrian deformation, were rifted away during Early Ordovician rollback of the plate boundary. Using the terrane model of Glen & Percival (2003) and Glen (2004), these calcalkaline Cambrian volcanic rocks form local substrate to the Bendigo Terrane, which accumulated as submarine fan systems off Antarctica and which was translated north to be accreted to the Australian part of East Gondwana in the Middle Devonian (see below). In this interpretation, there would be no continental crust (Selywn Block) under central Victoria, and no widespread Cambrian oceanic crust under the Southwestern subprovince: only locally stranded rifted segments embedded in younger Ordovician oceanic crust. The tectonic significance of the geophysically defined western volcanic belt, closer to the Gondwana margin, is more uncertain. It may represent an arc, possibly related to westdipping subduction and accreted at c. 530 Ma,

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before formation of the Kanmantoo Trough wich underwent deformation beginning at 515 Ma and followed by post-collisional volcanism, just as inferred from the eastern volcanic belt by Crawford et al. (1996). Such a history is consistent with the early cleavage and foreland basin formation cited earlier. The 'rift or arc' interpretation of the c. 526 Ma Koonenberry andesites is an important part in reconstructing this story and, thus, the nature of the Delamerian margin. The curvilinear, largely concealed, mafic and ultramafic rocks along the southern margin of the Thomson Orogen may indicate the presence of a third Delamerian arc, possibly mixed in with ocean crust igneous rocks from the inferred suture between the two orogens. If these rocks are Cambrian, then a major part of the convergent East Gondwana margin had a latitudinal trend and a north-south vector of subduction. The Cambrian ultramafic and related volcaniclastic rocks along the Peel-Manning Fault System in northeastern New South Wales are also enigmatic, since their geochemistry

Fig. 15. Seismic section from Lachian Orogen in the west into New England Orogen in the east (acquired by Geoscience Australia). Blueschists are found as knockers in schistose serpentinite along major faults such as the Peel Fault, part of the Peel-Manning Fault System. Based on Glen & Brown (1995) after Glen et al (1993).

R. A. GLEN

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suggests strong affinities with forearc rocks in western Tasmania and Victoria. Their location along a major Permian fault system implies that they have been thrust up from beneath the forearc basin of the New England Orogen. It is suggested here that they are allochthonous and part of Cambrian igneous substrate to the Lachlan Orogen (Fig. 15) that was translated northwards as part of major plate boundary rearrangement around the Ordovician-Silurian boundary (see below). Key features in the Lachlan supercycle Benambran cycle After the accretion of the old Delamerian Cambrian forearc crust, the Australian protoPacific plate boundary underwent rollback of c. 1000 km. This rollback is reflected by: 1. 2.

3.

post-collisional extension, volcanism and uplift in the Delamerian Orogen; splitting of the segments of the accreted Cambrian arc and forearc system that stretched from Antarctica to the New England Orogen (Miinker & Crawford 2000). One rifted segment is inferred from geochemical data to constitute substrate to the Macquarie Arc (Glen et al 2003; Crawford et al 2005). It was suggested above that a second segment is represented by Cambrian calc-alkaline volcanic rocks of the Mt Wellington Fault zone in central Victoria. These volcanic rocks were covered progressively by clastic sediments, then deep-sea cherts, as they became isolated from the Gondwana margin; formation of a wide back-arc basin (the Wagga Basin of Packham & Falvey 1971) in the Central and Western subprovinces filled by Early-Middle Ordovician turbidites. Earliest Ordovician turbidites are deposited on Cambrian igneous crust of the Delamerian cycle. The new Gondwana-proto-Pacific plate boundary was segmented (Fig. 16a). (a) The 'strongly' convergent northern segment, c. 1000 km long, formed opposite the Macquarie Arc. Behind the arc, the Wagga Basin developed as a back-arc basin floored by igneous crust and linked to the Macquarie Arc by the presence of Ordovician backarc basin basalts. Further development of this basin led to deposition of Early-Middle Ordovician turbidites. These are overlain by a starved Late

Ordovician black shale sequence (Colquhoun etal 2004) that developed as the depocentre moved out of range of terrigenous sedimentation from the continental margin (Glen 2004). These black shales are interpreted as responses to renewed phase(s) of back-arc basin opening (Fig. 16b). New igneous substrate formed during this spreading is manifest as Ordovician MORB-like volcanic rocks - the Narragudgil Volcanics (Duggan 2000), the Nacka Nacka Complex (Basden 1990; Meffre & Glen unpublished data) and tholeiitic basaltic schists of the Tottenham Group with MORB chemistry (Muir 1999). These volcanic rocks now occupy high-strain zones or fault blocks with Early-Middle Ordovician turbidites. Narrow belts of schistose serpentinite are inferred to mark faults cutting down into this substrate. (b) The longer, southern segment was probably strike-slip or highly oblique (Glen & Percival 2003; Glen 2004; Glen etal 2004) (Fig. 16a). It was marked by the presence of several large turbidite fan systems containing Early-Middle Ordovician craton-derived detritus. The Bega terrane was shunted to the north in the Late Ordovician to lie 'outboard' of the Macquarie Arc, and the Bendigo terrane was shunted northwards in the Silurian-Early Devonian (Fig. 16). The Bega terrane now lies 'outboard' of the Macquarie Arc (Fig. 16f). Most of the terrane is occupied by Early-Middle Ordovician cratonderived quartz-rich turbidites that are faulted against volcanic rocks of the Macquarie Arc and show no sign of mixing of provenance (Glen et al 1998; Meffre et al 2005). While submarine systems can bypass arcs to deposit material on the incoming oceanic plate, as canvassed by Glen et al (1998), this would still produce mixed provenance. Evidence from the Western Pacific suggests that this lack of provenance mixing constrains the two terranes to have formed at least hundreds of kilometres apart (Meffre et al 2005). The Bega terrane is thus inferred to have been shunted northward along the plate boundary (Figs 16b-c), with the incoming of the condensed Late Ordovician black shale sequence reflecting the drift of the terrane away from the Antarctic margin towards the Lachlan Orogen (Fig. 16b-c) (Glen 2004). The Bendigo

THE TASMANIDES OF EASTERN AUSTRALIA terrane remained anchored off Antarctica until the Silurian; alone of all the turbidite fan systems, it shows a west-east thinning of grain size and fining of rock packets consistent with increasing distance from the Gondwana margin. The view that the Late Ordovician condensed black shales sequence in the GirilamboneWagga terrane and Bega terrane represents periods of back-arc spreading and strike-slip along the Gondwana margin differs from that of Cas (1983), who suggested that the black shales represented deposition on submarine highs bypassed by turbidite distributary systems. Now that more is known about the areal extent of these shales, other origins must be sought VandenBerg & Stewart (1992) suggested sealevel rise. However, Late Ordovician turbidites occur just east of Melbourne, so this explanation does not work. Fergusson & Fanning (2002) suggested that the black shales in the Bega terrane represent background sedimentation when terrigenous deposition was blocked by subduction of the Wagga Basin. However, black shales also represent the Late Ordovician fill of that basin, the Girilambone-Wagga terrane. In this strike-slip model (Glen & Percival 2003; Glen 2004), Ordovician volcaniclastic rocks and blueschists along the Peel-Manning Fault System (and other faults) represent the missing forearc and accretionary prism rocks that originally formed outboard of the Lachlan Macquarie Arc (Fig. 16a). Cambrian ophiolites of the Delamerian cycle in the New England Orogen are similar chemically to those in the Lachlan Orogen (Aitchison & Ireland 1995) and were also caught up in this shunting, to be exhumed later along major faults (Fig. 16). This is somewhat similar to the idea of Cawood & Leitch (1985) that Ordovician volcaniclastic rocks along the Peel-Manning Fault System were part of the Ordovician Macquarie Arc, subsequently rifted off in the SilurianDevonian. Such an interpretation is tenable, since the southern part of the New England Orogen lies offshore and outboard of the Lachlan Orogen (Fig. 2), but it is argued here that strike-slip is still needed to explain these Lachlan-type rocks in the northern part of the southern New England Orogen. In Part 1 of this paper, evidence was cited that the New England Orogen was underlain by old, fast lithosphere. Here, it is suggested that this lithosphere may correspond to the Cambrian-Ordovician parts of the Lachlan Orogen and its older substrate. The two-stage Benambran Orogeny is explained by the Bega terrane entering the northern part of Gondwana-proto-Pacific plate boundary (Fig. 16c) and causing it to jam (Glen

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2004). As part of this deformation, the Bega terrane was accreted to the Macquarie Arc, which was driven into the back-arc basin, the Girilambone-Wagga terrane (Fig. 16d). Most of the deformation was taken up in these turbidite terranes by formation of multiple cleavages, folds and oblique thrusts. Earliest Silurian sinistral strike-slip faulting inserted some back-arc Ordovician craton-derived turbidites into an 'inter-arc' position between the western and central belts of Ordovician volcanic rocks (R. Scott unpublished data). After extension, a further deformation pulse resulted in the Girilambone-Wagga terrane being thrust obliquely over the arc while undergoing southwards translation. This second pulse also resulted in deformation of Llandovery turbidites and the syn-deformational emplacement of granites. In the strike-slip model, the Late Ordovician-Early Silurian 'Benambran' deformation in the Bendigo terrane occurred while that terrane was still 'moored' off West Antarctica. Deformation intensity decreased towards the ocean, with distal parts of the fan system overlain conformably by Silurian strata filling the Melbourne Trough in the Tabberabberan cycle.

Tabberabberan cycle Jamming of the subduction zone by the Bega terrane at the end of the Late Ordovician (Fig. 16d) caused rollback of the proto-Pacific plate until a new arc and inferred west-dipping subduction zone were established in the Late Silurian, further east in the present New England Orogen. The Tabberabberan cycle thus represents convergent margin relationships in the New England Orogen and back-arc basin extension in the Lachlan Orogen behind the arc. Several previous authors have argued for an extensional setting for the Tabberabberan cycle (e.g. Powell 1983, who envisaged extension in a dextral transform setting; Cas 1983; Scheibner 1989; Glen 1992; Scheibner & Veevers 2000; Collins 2002). The Lachlan and the North Queensland orogens were characterized by widespread rifting that was marked by the formation of rift/extensional basins and the emplacement of vast amounts of granite. In North Queensland, many of these granites were emplaced into Precambrian basement west of the North Queensland Orogen, separated from it by an extensional fault detachment (Blewett & Black 1998; P. Donchak pers. comm. Fig.; 9). Rifting in the Lachlan Orogen was terminated by the Mid Devonian Tabberabberan Orogeny: in the

THE TASMANIDES OF EASTERN AUSTRALIA

North Queensland Orogen by a less clearly defined Devonian-Carboniferous deformation. The lack of a prominent Mid-Devonian orogeny is not surprising if this deformation was driven by strike-slip tectonics in Victoria (see below). Rift basins of the Tabberabberan cycle in the Lachlan Orogen were discussed earlier. The large-scale crustal structure of the Tabberabberan cycle is still uncertain. Some of the shallowest basins (e.g. Cowra Trough) contain large amounts of volcanic rocks and intrusive granites that do not imply large extension, since the fill is only a few kilometres of generally shallow water to subaerial material (David & Glen unpublished data). Either the trough margins have been strongly overthrust or the heating is related to asymmetrical extension across the Lachlan Orogen (e.g. Lister el al 1991). One major question is how the different basement rocks responded to extension. The deformed Ordovician Macquarie Arc seems to have been pulled apart by extensional/oblique faults, with rifted parts forming basement to the deep-water Hill End Trough, the shallower Jemalong Trough and part of the Cowra Trough, as well as shallow-water flanking shelves (Glen et al 2002; Vassallo et al 2003). Unrifted parts occur as dispersed structural belts (Glen etal 1998). Seismic reflection profiling suggests that these extensional faults are largely planar (Glen et al 2002). However, the Tumut Trough, underlain by a complex basement of accreted Macquarie Arc juxtaposed against Ordovician turbidites and MORB-type volcanic rocks (Meffre & Glen unpublished data), formed by north-south transtension, and Stuart-Smith (1990) described a low-angle extensional detachment at the basement-cover interface. Basement consisting of Ordovician turbidites

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also underwent widespread mid-Silurian to mid-Devonian subsidence. This led to deposition of deep-water troughs but also shallowwater (to subaerial) shelves, commonly containing largely felsic volcanic and volcaniclastic rocks with lesser amounts of basalts and andesites. A key difference is the presence of large granitic batholiths intruding these Ordovician turbidites. It appears that clues to extension mechanisms and geometries in the middle crust may come from study of the widespread granites, which reflect major transfer of material from the lower crust and upper mantle into an actively extending middle crust. In the Lachlan Orogen, many authors noted that these voluminous granites were generated and emplaced in a back-arc extensional environment (e.g. Scheibner 1987; Fergusson 19920; Glen 1992; Scheibner & Basden 1998). Onedimensional modelling by Zen (1995) indicated that S-type melts could be generated in thinning crust. He also suggested that I-type granites east of the I-S line might reflect deeper-level melting in less extended crust. Chappell & co-workers have maintained that granite generation was unrelated to subduction: only the easternmost suite of I-type granites possesses geochemistry resembling East Pacific-type cordilleran granites (Chappell 1984). Other I-type granites were derived from 500-600 Ma tonalitic igneous crust (Chappell & Stephens 1988; Williams & Chappell 1998). S-type (and I-type) granites are now thought to be sourced from Ordovician turbidites underthrust during the Benambran Orogeny (Glen 1992) and mixed in various ratios with a mafic basaltic component (Gray 1984) or with two components - a mantlederived component and a component geochemically similar to Cambrian mafic volcanic rocks (Keay et al 1997; Collins 1998). Detrital zircon

Fig. 16. The strike-slip model, (a) Early and Middle Ordovician showing convergent and strike-slip nature of plate boundary and distribution of Adaminaby superterrane as series of fans along the east Gondwana margin. The Macquarie Arc, opposite the convergent plate margin in north, and the Bendigo terrane were built on rifted pieces of Delamerian forearc and post-collisional (p.c.) volcanic rocks. Western Tasmania (WT) was built on the Delamerian Orogen, and was the site of platform sedimentation. If the Bega and Bendigo terranes were sourced from the Ross Orogen, they would lie on the Gondwana Plate. If they were sourced from the Mozambique Belts they would occupy a diffuse transform plate margin, (b) Movement of the Bega terrane away from the Gondwana margin and its northward translation is reflected by the Late Ordovician condensed black shale sequence. The onset of this translation was synchronous with a hiatus in volcanism in the Macquarie Arc (due to seamount impinging on the trench) and with inferred back-arc basin spreading in the Girilambone-Wagga terrane. (c), (d) The Bega terrane was translated north along the plate margin to lie outboard of the Macquarie Arc, where it blocked oblique subduction and was then accreted obliquely to the Macquarie Arc, which was driven west into the Girilambone-Wagga terrane. Bendigo terrane was deformed off the Antarctic sector of the Gondwana margin, (e) Northwards translation of the Bendigo terrane during fill of the Melbourne Zone from sources in the south, west and southwest until just before accretion when an eastern source became available. Accretion involved underthrusting of the Girilambone-Wagga terrane which moved south, (f) Cretaceous reconstruction after (Lawver & Gahagan 1994) showing final distribution of terranes.

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work (Williams & Chappell 1998) confirms that the source need be no older than Early Ordovician, and the fertile nature of these rocks is confirmed by the recognition that older parts of the turbidite pile are more lithic and felspathic than the upper parts (M. Scott & O.Thomas, pers. comm.). Melting to produce these Silurian-Devonian granites probably began during thrust thickening of Ordovician turbidites in the Benambran Orogeny (Pagan 1979; Pogson 1982; Glen 1992) and continued during crustal extension. Collins & Hobbs (2001) showed that input from mantle melts was also needed to generate granitic magmas. Mantle input is also supported by the association of intrusive gabbro to dioritic plutons and associated dykes around granites and also by the presence of mafic-intermediate enclaves (e.g. Gray 1984; Soesoo & Nicholls 1999; Collins & Hobbs 2001). The subduction signature of some Silurian-Devonian granites and volcanic rocks is ascribed to inheritance from Ordovician subduction (e.g. Watkins 1998b; Glen 1998), rather than indicating the roots of an arc (Soessoo et al 1997; Collins & Hobbs 2001 see below), although geochemical modelling is needed to confirm this possibility. This argument becomes stronger when examined in the light of the change from Late Carboniferous convergence to Early Permian extension in the New England Orogen. Several authors cited in Part 2 of this paper have pointed out that the first melts in the Early Permian carry subduction signatures inherited from Carboniferous subduction (e.g. Brownlow & Arculus 1999; Caprarelli & Leitch 2001; Jenkins et al 2002). The Tabberabberan Orogeny at the end of the Tabberabberan cycle reflects accretion of the northward-shunting Bendigo terrane to Gondwana (Glen et al 19926; VandenBerg et al 2000; Willman et al 2002) (Fig. 16e). This terrane, structurally represented by the Southwestern subprovince, was inserted between the Delamerian Orogen on the west in the Late Siluarian-Middle Devonian by a combination of mild north-south shortening, coupled with strike-slip deformation and strong east-west shortening (Glen et al 1992; Miller et al 2001) and the Eastern and Central subprovinces of the Lachlan Orogen on the east in the Middle Devonian. As it moved north, it received a mixture of sandstone and shale from sources in the west and south: easterly sources only became apparent as the terrane approached the Gondwana margin in the late Early Devonian. The strike-slip model of VandenBerg et al (2000) and Willman et al (2002) has opposite

dynamics to that of this paper, with Eastern Australia (Eastern and Central subprovinces) moving dextrally SSE rather than the Bendigo terrane moving to the north.

Other models Plate tectonic models in the early 1970s envisaged only one subduction zone at the western margin of the proto-Pacific plate (e.g. Oversby 1971; Solomon & Griffiths 1972). Updates of these ideas were presented by several authors (e.g. Powell 19840; Scheibner 1985; Coney 1992). Models invoking multiple subduction zones were also proposed, e.g. the multiple back-arc model'of Scheibner (1973, but not later), the sequential forearc model of Crook (1980) and back-arc model of Collins & Vernon (1992). The last suggested two subduction zones behind the plate boundary: an arc related to east-dipping subduction along the boundary between the Central and Southwestern subprovinces at 435 Ma, and a west-dipping subduction zone at c. 400 Ma within the Southwestern subprovince to close the Melbourne Trough, interpreted as a back-arc basin. Gray and co-workers further developed the concept of multiple subduction zones, based on Ar-Ar dating in the Southwestern subprovince that showed that the oldest cooling ages were Ordovician and became younger (midDevonian) in the east (Gray et al 1997; Foster et al 1999; Foster & Gray 2000) (Fig. 6). This eastward younging in cooling ages was interpreted as a single 'diachronous' transgressive deformation extending over 50 million years (Gray et al 1997). This concept of continuous deformation led Gray & Foster (1997: 880-881) to suggest that 'the time from what was previously defined as Benambran to Tabberabberan is redefined as one progressive orogenic episode that we now call the Lachlan Orogeny after Cas (1983)'. This prograding deformation was likened to that occurring in an accretionary wedge above a subduction complex, with migrating deformation related to 'subduction accretion during plate convergence' in an oceanic setting with three subduction zones active in the mid-Palaeozoic (Gray et al 1997: 497). Fault vergence was thus regarded as antithetic to the dips of subduction zones. However, modelling by Keep (2003) showed that oceanward-verging thrusts could form without subduction of oceanic lithosphere. Whereas Gray & coworkers suggested that the Benambran Orogeny reflected progressive growth of an accretionary wedge, Collins & Hobbs (2001) suggested that the two phases

THE TASMANIDES OF EASTERN AUSTRALIA identified above reflect two separate, synchronous subduction-acrcetion/magmatic arc complexes, with Early Silurian S-type granites forming the roots of magmatic arcs. Both are in contrast to the model of this paper. Subduction zone 1 in the west (Gray & Foster 1997) was shallowly west-dipping and existed from 460 Ma to 420 Ma (Foster et al 1999). No arc was associated with this zone. CambrianOrdovician turbidites in the Stawell and Bendigo structural zones, deformed into eastvergent structures, were interpreted as having formed in the associated accretionary prism that was advancing progressively eastward and closing a marginal basin (Foster et al 1999). Subduction zone 2 was short lived (440^20 Ma, Foster et al 1999), was east-dipping and lay at the boundary between the Melbourne and Tabberabbera zones (between the Central and Southwestern subprovinces). There is no arc associated with this subduction zone, although some syn- to post-tectonic granites in the Central subprovince have subduction-related signatures. Ordovician turbidites with southwest-verging folds (Fergusson 1987) in the Tabberabbera and Omeo zones were interpreted as accretionary prism rocks and reflections of a magmatic arc. Subduction zone 3 was the Gondwana-proto-Pacific plate boundary. The associated arc was the Macquarie Arc of Glen et al (1998). The oceanic Narooma Terrane and surrounding Ordovician turbidites were interpreted as the related accretionary prism, after Miller & Gray (1996), although this was argued against by Glen et al (2004). The multiple subduction model was taken further by Soesoo et al (1997), who suggested that subduction zone 1 persisted until the Middle Devonian (c. 380 Ma) and that the Tabberabbera zone was underlain by two opposing subduction zones - the east-dipping subduction zone 2 of Foster et al (1999) and an extra west-dipping zone - from 420 to c. 380 Ma. The existence of this subduction zone is supported by the presence of c. 450 Ma blueschist knockers along the boundary between the two structural zones (Spaggiari et al 2002) and by the subduction signatures of granitic rocks (Nicholls et al 1996). The 'Lachlan Orogeny' was terminated by the late Early to Middle Devonian closure of the marginal basin (original width of c. 750 km, now the Melbourne Trough) and by docking of an island arc/forearc system to the east (Gray & Foster 1997). Negative Nb and Ti anomalies in some Early Devonian granites have been used to infer the influence of subduction on mantle magma

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sources, suggesting that these tracts of Late Silurian and Early Devonian granites are the roots of coeval subduction-related arcs (Soesoo et al 1997). In contrast, Middle and Late Devonian mafic rocks lack significant Nb anomalies and were interpreted as emplaced in continental-rift extensional settings (Soesoo & Nicholls 1999). Collins & Hobbs (2001) proposed the existence of two coeval subduction zones from 435 Ma to 425 Ma, with arcs reflected by the distribution of NW- and N-trending belts of largely S-type granites. The multiple subduction model was extended by Fergusson (2003) who argued for four subduction zones In contrast to the multiple subduction models, Collins & Vernon (1994) suggested a model involving east-west sequential delamination of lithospheric slabs, with granitoids generated from basaltic underplating rather than from subduction melts as in the earlier model. VandenBerg et al (2000) and Cayley et al (2002) argued for a vice model, in which the deformation of the Southwestern subprovince was driven by compression between the Delamerian Orogen backstop in the west and a block of Proterozoic continental crust under the Melbourne Zone in the east, driven westwards by plate boundary forces. This 'rifted-off' piece of continental crust has been incorporated as an alternative into the multiple subduction model by Gray et al (2003). Subduction zones, however, require special conditions to form and do not turn on and off easily. Strike-slip 'transpressional' tectonics, coupled with models in which the subduction signatures of Silurian-Early Devonian granites are inherited from Ordovician subduction, are alternatives to the multiple subduction zone models. The problem with strike-slip tectonics is that it is based largely on palaeogeographical reconstructions, since kinematic evidence of hundreds of kilometres of strike-slip movement has not been established. However, it is argued that strike-slip deformation must be the norm in convergent margin orogens (e.g. Teyssier & Tikoff 1995), except in rare cases where convergence is head-on to a perfectly planar plate boundary, and that much of it can be achieved by summing up smaller (tens of kilometres) displacements on a number of smaller faults.

Concluding discussion The Tasmanides of eastern Australia have been part of Gondwana, and then part of Pangaea before it began to break apart at c. 230 Ma. The Tasmanides faced the proto-Pacific Ocean following supercontinent break-up at c. 750 Ma.

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There is no evidence of continent-continent collision, or that the Wilson Cycle of ocean opening and closing ever operated (Crawford et al 20030; Cawood 2002, 2005). Crook (1969) first pointed out this difference between Atlantic and Pacific geosynclines. Most of the history of the Tasmanides records extension or rifting. With the exception of passive margin development during rifting of Rodinia in the Neoproterozoic, it is envisaged that all subsequent evolution of the Tasmanides occurred in a convergent margin setting along the proto-Pacific plate. Phases of extension and rifting are separated by deformation events that occupy only short intervals during the development of the Tasmanides (Fig. 4). They were probably more complicated than current models suggest, and probably diachronous along-strike. It was suggested previously (Glen 1995) that the Lachlan Orogen, in particular, possessed many of the features of a convergent margin in which the rate of convergence of the two plates was less than the rate of subduction (cf. Royden & Burchfiel 1989; Royden 1993). Such an orogen is characterized by extension, lack of major collision and resultant high topography, and an absence of faults that exhume high-grade basement. Today, one would call this a retreating or accretionary orogen, defined by rollback of the proto-Pacific plate (Cawood 2002; Collins 2002). To some extent, this concept can be applied to the other parts of the Tasmanides (e.g. Cawood 2002). However, there is little evidence for rollback of the proto-Pacific Plate in the North Queensland Orogen, and the plate boundary in the New England Orogen was fixed in position effectively from the Late Devonian to the Late Permian. This evolutionary model for the Tasmanides - accretionary in nature, with distinct, short periods of advance or accretion/collision events (e.g. Collins 2002; Crawford 20030) - differs from that of Gray and co-workers for the Lachlan Orogen, in which most of the deformation of the Lachlan Orogen occurred in propagating accretionary wedges, with deformation reflecting times when wedges were thickened by internal deformation. Despite occupying small intervals of time, some orogeny (at least) seems to have been complex and either protracted or multiphase. The Delamerian Orogeny in the outboard parts of the Delamerian Orogen (western Tasmania and western Victoria) consisted of several phases, but seems to be more simple than the same orogeny in the inboard part of the orogen, where dating of foliated granites suggest that it lasted for 26 million years. The Benambran

Orogeny is another protracted or multiphase event lasting c. 10 million years. In accretionary orogens, what produces deformations? There are several possibilities. (a) Changes in plate motions. The evolution of modern SW Pacific tectonics is generally seen as due to responses to changes in plate motions (e.g. Crawford et al 20030). A possible candidate in the Tasmanides is the inferred Carboniferous accretion of the New England Orogen to the Lachlan Orogen as the driver for the Kanimblan Orogeny, for which no other cause is clear. Veevers (20006) pointed out that this deformation could have been a far-field result of the collision between Gondwana and Laurussia, although it seems to be a little older than the main collision phase in the Variscan belt (c. 340 Ma cf. 330-320 Ma, Veevers 20006). (b) Change in plate boundary dynamics or in coupling, from retreating to advancing, possibly in response to changes in spreading rates in (a). Carboniferous-Triassic deformation and extension in the HunterBo wen super cycle orogen were attributed to this mechanism by Jenkins et al (2002). On a larger scale, differences between the Delamerian and New England orogens, on one hand, and the Lachlan Orogen, on the other, might reflect formation along largely advancing, as opposed to largely retreating, parts of the plate margin. Both the Delamerian and New England orogens, despite their differences, are classical orogens in the sense of having highly deformed and metamorphosed internal parts, containing accreted complexes and less deformed external parts consisting of fold-thrust belts. The Lachlan Orogen is more of a retreating orogen with deformations that did not produce high topography (Glen 1992). (c) Collision between the Gondwana margin and a collider such as a large seamount or submarine plateau or island arc on the proto-Pacific plate. Examples in the Tasmanides include the arc-continent collision that produced the Delamerian Orogeny and, perhaps, also the accretion of the Gympie Terrane with Gondwana to produce the Triassic part of the HunterBowen Orogen. (d) Shuffling of terranes along-strike in a 'transpressional' regime that also involved shortening. Examples include the Benambran Orogeny as a response to shuffling of

THE TASMANIDES OF EASTERN AUSTRALIA the Bega terrane to a location outboard of the Macquarie Arc, located on the Gondwana plate, and its resulting accretion, which drove the arc into its back-arc basin. The Tabberabberan Orogeny, which developed as a result of accretion of the Bendigo terrane with other parts of the Lachlan Orogen, is another example. The end-Silurian Bindi deformation seems to be a smaller-scale version of this, reflecting strike-slip-driven movements on linked faults. There are several other key features that can be distilled from the synthesis of the Tasmanides. (1) There are significant differences in the evolution of the Lachlan and North Queensland orogens, despite the general recognition of elements of the Lachlan supercycle in the North Queensland Orogen, indicating that both underwent Ordovician convergence followed by Silurian to Middle (or Late) Devonian extension. Their subsequent evolution differs: there is no Kanimblan cycle, represented by major foreland-style fluviatile deposits, in the North Queensland Orogen. The primary difference between the two orogens is that the North Queensland Orogen was developed on continental crust of the Delamerian cycle, whereas the Lachlan Orogen formed on oceanic crust east of the Delamerian margin. The second difference flows from this the North Queensland Orogen lies immediately east of the inboard boundary of the Tasmanides, whereas the Lachlan Orogen formed well to the east of the Tasmanides margin. These differences suggest that the rapid rollback of the Gondwana-proto-Pacific plate boundary in southeastern Australia - after the Delamerian and Benambran orogenies - was not mirrored in North Queensland, where the plate boundary was more or less anchored in time from the Cambrian through to the Triassic. This lack of major back-arc spreading is, in turn, reflected in the continental margin nature of inferred arc(s) rather than the intra-oceanic nature of the Lachlan arcs. The evolution of the North Queensland Orogen is, thus, to a limited extent, more akin to the evolution of Palaeozoic South America (e.g. Ramos 1988; Rapela et al. 1998). The morphology of the proto-Pacific margin of Gondwana must have reflected this disparity, either by segmentation or by variation of spreading. (2) The relationship between New England Orogen and orogens to the west is uncertain. Although there is strong continuity and parallel development between the northern and

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southern parts of the New England Orogen, there are striking differences between the New England Orogen and the Lachlan, Thomson and North Queensland orogens to the west, all of which occupied back-arc positions from the Middle Silurian to the Triassic. Similarities between the Cambrian to Ordovician geology of the Lachlan Orogen and that of Cambrian-Ordovician fault-bounded blocks along the Peel-Manning Fault System in the southern New England Orogen were pointed out by Cawood & Leitch (1987) and Glen & Scheibner (1993). These similarities are interpreted here in terms of the Lachlan rocks being translated north in the Llandovery, to form substrate to the developing New England Orogen, and exhumed along the Peel-Manning Fault System. (Alternatively, Aitchison & Ireland (1995) suggested the New England Orogen overthrust the Lachlan Orogen in the Late Carboniferous.) In the Silurian and Devonian, the two orogens had very different histories, despite the southern part of the New England Orogen lying outboard of the Lachlan Orogen after restoration of the Permian oroclines. This contrast is most evident in the Late Devonian, when fluviatile siliciclastic sediments extended from the Delamerian Orogen to the present NSW south coast, and when the western part of the New England Orogen was the site of a continental arc. The presence of volcanic detritus in the Lambie facies east of the Hill End Trough (Powell et al. 1984) suggests proximity, as does the presence of Lambie-facies type clasts in the forearc basin of the New England Orogen in the Late Visean-Namurian (Cawood & Leitch 1985). If the New England Orogen was allochthonous (Klootwijk 2002), it was approaching the Lachlan Orogen in the Late Devonian, and docked in the Early Carboniferous, at about the time of the Kanimblan Orogeny. The lack of a Lambie facies in the Thomson and North Queensland orogens reduces this contrast with the northern New England Orogen. Both the Thomson and North Queensland orogens have undergone extension, contraction and granite emplacement, extending into the Triassic, all consistent with their being back-arc systems to the New England Orogen. Several questions still remained to be answered, not the least of which is more closely tying granite genesis into upper crustal histories. Another question not touched on here is how to reactivate the orogen-normal and orogenoblique structures that seem to persist through these cycles and through the change from oceanic to continental crust.

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The author wishes to thank the many people who have tried to help him understand, over many years, the formation of the Tasmanides. Among these in particular are Jeff Beckett, Tony Crawford, Vlad David, Paul Donchak, Barry Drummond, David Gray, Russell Korsch, Sebastien Meffre, Robin Offler, Ian Percival, Dennis Pogson, Chris Powell, Wolfgang Preiss, John Roberts, Rob Scott, Erwin Scheibner, Ian Stewart, John Walshe, Jeff Vassallo and Fons VandenBerg. Ideas on inheritance of geochemical signatures involved discussions with John Watkins and Larry Barren. Paul Donchak (Geological Survey of Queensland) kindly permitted the use of his, as yet unpublished, figures on the evolution of the Hodgkinson subprovince in the North Queensland Orogen. Thanks also go to the management of the Geological Survey for support over the years. Special thanks are due to Robyn Sharpe, Phil Carter, Margaret McLaren and others in the Geospatial Group of the Survey for their cartographic work. The paper benefited greatly from the reading and criticisms of Paul Donchak, Cec Murray, Jeff Vassallo, John Roberts and Ian Percival, and from reviewers Peter Cawood and Tony Crawford who had to deal with a manuscript over and above the normal call of duty. In addition, Bob Pankhurst is thanked for editorial guidance and critical comments. Financial support from the Australian UNESCO committee for the International Geological Correlation Programme and from the Tapmog Conference organizers is much appreciated. This is a contribution to IGCP project 436. Published with the permission of the Director-General, New South Wales Department of Primary Industries.

References ABELL, R.S. 1993. Geology of the Canberra 1:100000 sheet area. Bureau of Mineral Resources, Canberra, Bulletin 233. ACENOLAZA, F. & TOSELLI, A. 1976. Consideraciones estratigraficas y tectonicas sobre el Paleozoico inferior del Noroeste Argentine. 2nd Congreso Latinoamericano de Geologia, Adas. 2, Caracas, 755-763. AITCHISON, J.C. 1988. Early Carboniferous radiolaria from the Neranleigh-Fernvale beds near Brisbane. Queensland Government Mining Journal, 89, 240-241. AITCHISON, J.C. & FLOOD, P.G. 1990. Geochemical constraints on the depositional setting of Palaeozoic cherts from the New England Orogen, NSW, eastern Australia. Marine Geology, 94, 79-95. AITCHISON, J.C. & FLOOD, P.G. 1992. Implications of radiolarian research for analysis of subduction complex terranes in the New England Orogen, NSW, Australia. Palaeogeography, Palaeoclimatology, Palaeoecology, 96, 89-102. AITCHISON, J.C. & FLOOD, P.G. 1995. Gamilaroi terrane: a Devonian rifted intra-oceanic island arc assemblage, NSW Australia. In: SMELLIE, J.L. (ed.) Volcanism Associated with Extension at Consuming Plate Margins. Geological Society, London, Special Publications, 81,155-168.

AITCHISON, J.C. & IRELAND, T.R. 1995. Age profile of ophiolitic rocks across the Late Palaeozoic New England Orogen, New South Wales: implications for tectonic models. Australian Journal of Earth Sciences, 42, 11-23. AITCHISON, J.C., FLOOD, P.G. & SPILLER, F.C.P. 19920. Tectonic setting and palaeoenvironment of terranes in the southern New England Orogen, eastern Australia as constrained by radiolarian biostratigraphy. Palaeogeography, Palaeoclimatology, Palaeoecology, 94, 31-54. AITCHISON, J.C., IRELAND, T.R., BLAKE JR, M.C. & FLOOD, P.G. 19926. 530 Ma zircon age for ophiolite from the New England Orogen: oldest rocks known from eastern Australia. Geology, 20, 125-128. AITCHISON, J.C., BLAKE JR, M.C., FLOOD, P.G. & JAYKO, A.S. 1994. Paleozoic ophiolite assemblages within the southern New England Orogen of eastern Australia: implications for the growth of the Gondwana margin. Tectonics, 13,1135-1149. ALDER, D. 1999. Drilling to follow seismic survey in the west. Minfo (New South Wales Mining and Exploration Quarterly), 64, 29-30. ALDER, J.D., HAWLEY, S.P., MAUNG, T., SCOTT, J., SHAW, R.D., SINELNIKOV, A. & KOUZMINA, G. 1998. Prospectivity of the Offshore Sydney Basin: a new perspective. Appea Journal, 38, 68-91. ALLAN, A.D. & LEITCH, E.C. 1990. The tectonic significance of unconformable contacts at the base of Early Permian Sequences, southern New England Fold Belt. Australian Journal of Earth Sciences, 37, 43-49. ARNOLD, G.O. & FAWCKNER, J.F. 1980. The Broken River and Hodgkinson Provinces. In: HENDERSON, R.A. & STEPHENSON, P.J. (eds) The Geology and Geophysics of Northeastern Australia. Geological Society of Australia, Queensland Division, Brisbane, 175-189. BAIN, J.H.C. & DRAPER, J. J. (eds) 1997. North Queensland Geology. Bulletin of the Australian Geological Survey Organisation 240, and Queensland Department of Mines and Energy Geology 9. BAIN, J.H.C., WYBORN, D., HENDERSON, G.A.M., STUART-SMITH, P.G, ABELL, R.S. & SOLOMON, M. 1987. Regional geologic setting of the Woodlawn and Captains Flat massive sulphide deposits, N.S.W., Australia. Abstracts of the Pacific Rim Congress 87, Australasian Institute of Mining and Metallurgy, Surfers Paradise, 25-28. BARRON, L.M. 1999. Metamorphism. Dubbo 1:250 000 Geological Sheet SI/55-4, 2nd edition. Explanatory notes. Geological Survey of New South Wales, Sydney. BARTON, TJ. 1999. Crustal structure of northern Tasmania based upon a deep seismic transect. Abstracts of the Geological Society of Australia, 53, 3-4. BASDEN, H. (ed.) 1990. Geology of the Tumut 1:100 000 sheet 8527. Geological Survey of New South Wales, Sydney. BECKETT, J., GLEN, R.A., KORSCH, RJ. & BOREHAM, C. 1995. Hydrocarbons below the Hunter Thrust:

THE TASMANIDES OF EASTERN AUSTRALIA a Cordilleran model for petroleum potential in the Maules Creek Sub Basin, Gunnedah New South Wales. Proceedings of the 1995 NSW Petroleum Symposium. Petroleum Exploration Association of Australia, NSW Branch, Sydney. BELPERIO, A.P., PREISS, W.V., FAIRCLOUGH, M.C., GATEHOUSE, C.G., GUM, J., HOUGH, J. & BURTT, A.C. 1998. Tectonic and metallogenic framework of the Cambrian Stansbury Basin-Kanmantoo Trough, South Australia. AGSO Journal of Australian Geology and Geophysics, 17,183-200. BEMBRICK, C.S. 1997. A re-appraisal of the Darling Basin Devonian sequence. Quarterly Notes of the Geological Survey of New South Wales, 105,1-16. BENNETT, V.C., KENNETT, B.L.N. & CARLSON, R.W. 2002. Combined Osmium isotopic and seismic evidence for orphaned early Proterozoic mantle beneath Phanerozoic crust in the New England Fold Belt, eastern Australia. Abstracts of the Geological Society of Australia, 67, 100. BERRY, R.F. 1994. Tectonics of western Tasmania: Late Precambrian-Devonian. Abstracts of the Geological Society of Australia, 39, 6-8. BERRY, R.F. 1995. Palaeozoic tectonics of western Tasmania. Abstracts of the Geological Society of Australia, 40, 11-12. BERRY, R.F. & CRAWFORD, AJ. 1988. The tectonic significance of Cambrian allochthonous maficultramafic complexes in Tasmania. Australian Journal of Earth Sciences, 35, 523-533. BERRY, R.F, HUSTON, D.L., STOLZ, A.X, HILL, A.P., BEAMS, S.D., KURONEN, U. & TAUBE, A. 1992. Stratigraphy, structure, and volcanic-hosted mineralization of the Mount Windsor Subprovince, North Queensland, Australia. Economic Geology, 87, 739-763. BERRY, R.F, JENNER, G.A., MEFFRE, S. & TUBRETT, M.N. 2001. A North American provenance for Neoproterozoic to Cambrian sandstones in Tasmania? Earth and Planetary Science Letters, 192, 207-222. BETTS, P.G., GILES, D., LISTER, G.S. & FRICK, L.R. 2002. Evolution of the Australian lithosphere. Australian Journal of Earth Sciences, 49,661-695. BIERLEIN, F.P., HUGHES, M.J., DUNPHY, J. & MCKNIGHT, S. 2001. Tectonic and economic implications of trace element, 40Ar/39Ar and Sm-Nd data from mafic dykes associated with erogenic gold minerals in central Victoria, Australia. Lithos, 58,1-31. BLEVIN, P.L. 1998. Paleozoic igneous and metallogenic evolution of the LFB. In: LEWIS, P.C. (ed.) 'Lachlan Fold belt conference'. Bulletin of the Australian Institute of Geoscientists, 23, 25-26. BLEWETT, R.S. & BLACK, L.P. 1998. Structural and temporal framework of the Coen region, north Queensland: implications for major tectonothermal events in east and north Australia. Australian Journal of Earth Sciences, 45, 597-609. BLEWETT, R.S., BLACK, L.P, SUN, S-S., KNUTSON, J., HUTTON, LJ. & BAIN, J.H.C. 1998. U-Pb zircon and Sm-Nd geochronology of the Mesopoterozoic of north Queensland: implications for a Rodinian connection with the Belt Supergroup of

79

North America. Precambrian Research, 89, 101-127. BRADLEY, G. 1993. Evolution and hydrocarbon prospectivity of the Offshore Sydney basinNSW/P10. In: SWARBRICK, G.J. & MORTON, D.J. (eds) Proceedings of the NSW Petroleum Symposium. Petroleum Exploration Society of Australia, NSW Branch, Sydney. BRILL, B.A. 1989. Deformation and recrystallization microstructures in deformed ores from the CSA mine, Cobar, N.S.W., Australia. Journal of Structural Geology, 11, 591-601. BROWN, D.A., CAMPBELL, K.S.W. & CROOK, K.A.W. 1968. The Geological Evolution of Australia and New Zealand. Pergamon Press, Oxford. BROWN, R.E. 1987. Newly defined stratigraphic units from the western New England Fold Belt, Manilla 1:250000 sheet area. Quarterly Notes of the Geological Survey of New South Wales, 69, 1-9. BROWNLOW, J.W. 1979. Discussion: The Reids Mistake Formation at Swansea Heads, New South Wales. Journal of the Geological Society of Australia, 26, 319-322. BROWNLOW, J.W. 1999. Geological setting and the origin of the Early Permian Boggabri Volcanics, Gunnedah Basin, NSW. In: FLOOD, P.G. (ed.) Conference on Regional Geology, Tectonics and Metallogenesis, New England Orogen. University of New England, Armidale, 325-333. BROWNLOW, J.W. & ARCULUS, R. 1999. A doublydepleted-double enriched element abundance pattern from a basalt in the Early Permian Boggabri Volcanics, Gunnedah Basin, NSW. In: FLOOD, P.G. (ed.) Conference on Regional Geology, Tectonics and Metallogenesis, New England Orogen, University of New England, Armidale, 335-339. BRUCE, M.C. & Niu, Y. 20000. Early Permian suprasubduction assemblage of the South Island terrane, Percy Isles, New England Fold Belt, Queensland. Australian Journal of Earth Sciences, 47, 1077-1085. BRUCE, M.C. & Niu, Y. 2000ft. Evidence for Palaeozoic magmatism recorded in the Late Neoproterozoic Marlborough ophiolite, New England Fold Belt, central Queensland. Australian Journal of Earth Sciences, 47, 1065-1076. BRUCE, M.C., Niu, Y., HARBORT, T.A. & HOLCOMBE, R.J. 2000. Petrological, geochemical and geochronological evidence for a Neoproterozoic ocean basin recorded in the Marlborough terrane of the northern New England Fold Belt. Australian Journal of Earth Sciences, 47, 1053-1064. BRYAN, S.E., CONSTANTINE, A.E., STEPHENS, C.J., EWART, A., SCHOEN, R.W. & PARIANOS, J. 1997.

Early Cretaceous volcano-sedimentary successions along the eastern Australian continental margin: Implications for the breakup of eastern Gondwana. Earth and Planetary Science Letters, 153, 85-102. BRYAN, S.E., HOLCOMBE, R.J. & FIELDING, C.R. 2001. Yarrol Terrane of the northern New England

R. A. GLEN Fold Belt: forearc or backarc? Australian Journal of Earth Sciences, 48, 293-316. BRYAN, S.E., FIELDING, C.R., HOLCOMBE, R.J., COOK, A. & MOFFITT, C.A. 20030. Stratigraphy, fades architecture and tectonic implications of the Upper Devonian to Lower Carboniferous Campwyn Volcanics of the northern New England Fold Belt. Australian Journal of Earth Sciences, 50, 377-401. BRYAN, S.E., HOLCOMBE, RJ. & FIELDING, C.R. 20036. Reply: Yarrol Terrane of the northern New England Fold Belt, forearc or back arc? Australian Journal of Earth Sciences, 50, 278-293. BRYANT, C.J., ARCULUS, RJ. & CHAPPELL, B.W. 1997. Clarence River Supersuite: 250 Ma Cordilleran tonalitic I-type intrusions in eastern Australia. Journal of Petrology, 38, 975-1001. BUCKLEY, P.M. 2003. The Mount Daubeny Basin: mineralised Late Silurian to Early Devonian pullapart structure. In: PELJO, M. (ed.) Broken Hill Exploration Initiative. Geoscience Australia Record, 2003/13,16-17. BULTITUDE, RJ., DONCHAK, PJ.T., DOMAGALA, J. &

FORDHAM, B.G. 1993. The pre-Mesozoic stratigraphy and structure of the western Hodgkinson Province and environs. Queensland Geological Record, 1993/29. BULTITUDE, RJ. & GARRAD, P.D. (compilers). 1997. Geology of the Innisfail Region. Record of the Department of Mines and Energy Queensland, Brisbane, 1997/2. CALVER, C.R. & WALTERS, M.R. 2000. The Late Neoproterozoic Grassy Group of King Island, Tasmania: correlation and palaeogeographic significance. Precambrian Research, 100, 299-312. CAO, X.N. & DURNEY, D.W. 1993. Folding and cleavage in the eastern Tamworth belt, Manilla, NSW. In: FLOOD, P.G. & AITCHISON, J.C. (eds) Abstract volume, Conference on the New England Orogen, eastern Australia. University of New England, Armidale, 255-256. CAPRARELLI, G. & LEITCH, E.C. 1998. Magmatic changes during the stabilisation of a cordilleran fold belt: the Late Carboniferous-Early Triassic igneous history of eastern New South Wales, Australia. Lithos, 45, 413-430. CAPRARELLI, G. & LEITCH, E.C. 2001. Geochemical evidence from Lower Permian volcanic rocks of northeast New South Wales for asthenospheric upwelling following slab breakoff. Australian Journal of Earth Sciences, 48, 151-166. CARITAT, P.D. & BRAUN, J. 1992. Cyclic development of sedimentary basins at plate margins-1. Structural and tectono-thermal evolution of some Gondwana basins of eastern Australia. Journal of Geodynamics, 16, 241-282. CARR, G., DEAN, J.A., SUPPEL, D.W. & HEITHERSAY, P.S. 1995. Precise lead isotope fingerprinting of hydrothermal activity associated with Ordovician to Carboniferous metallogenic events in the Lachlan Fold Belt of New South Wales. Economic Geology, 90,1467-1505. CAS, R.A.F. 1983. A review of the fades patterns

palaeogeographic development and tectonic context of the Palaeozoic Lachlan Fold Belt of southeastern Australia. Geological Society of Australia, Special Publication, 10. CAS, R.A.F & VANDENBERG, A.H.M. 1988. Ordovician. In: DOUGLAS, J.G. & FERGUSON, J.A. (eds) Geology of Victoria. Geological Society of Australia, Victorian Division, Melbourne, 63-102. CAWOOD, P.A. 1976. Cambro-Ordovician strata from northern New South Wales. Search, 7, 317-318. CAWOOD, P.A. 1980. The geological development of the New England Fold Belt in the WoolominNemingha and Wisemans Arm regions: the evolution of a Palaeozoic fore-arc terrain. PhD thesis, University of Sydney, Sydney. CAWOOD, P.A. 19820. Tectonic reconstruction of the New England Fold Belt in the Early Permian: an example of development of an oblique-slip margin. In: FLOOD, P.G. & RUNNEGAR, B. (eds) New England Geology, Proceedings, Voisey Symposium. University of New England, Armidale, 25-34. CAWOOD, P.A. 1982£. Structural relations in the subduction complex of the New England Fold Belt, eastern Australia. Journal of Geology, 90, 381-392. CAWOOD, P.A. 1984. The development of the SW Pacific margin of Gondwana: correlations between the Rangitata and the New England Orogens. Tectonics, 3, 539-553. CAWOOD, P.A. 2005. Terra Australia Orogen: Rodinian breakup and development of the Pacific and lapetus margins of Gondwana during the Neoproterozic and Paleozoic. Earth Science Reviews, 69, 249-279. CAWOOD, P.A. & FLOOD, R.H. 1989. Geochemical character and tectonic significance of Early Devonian keratophyres in the New England Fold Belt, eastern Australia. Australian Journal of Earth Sciences, 36, 297-311. CAWOOD, P.A. & LEITCH, E.C. 1985. Accretion and dispersal tectonics of the Southern New England Fold Belt, eastern Australia. In: HOWELL, D.G. (ed.) Tectonostratigraphic Terr ones of the Circum-Pacific Region. Circum-Pacific Council for Energy and Mineral Resources, Houston, 1, 481-492. CAYLEY, R.A. & TAYLOR, D.H. 1996. Geological evolution and economic potential of the Grampians area, Victoria. In: HUGHES, M.J., Ho, S.E. & HUGHES, C.E. (eds) Recent Developments in Victorian Geology and Mineralisation. Bulletin of the Australian Institute of Geoscientists, 6, 11-18. CAYLEY, R.A., TAYLOR, D.H., VANDENBERG, A.H.M. & MOORE, D.H. 2002. Proterozoic-early Palaeozoic rocks and the Tyennan Orogeny in central Victoria: the Selwyn Block and its tectonic implications. Australian Journal of Earth Sciences, 49, 225-254. CHAPPELL, B.W. 1984. Source rocks of I- and S-type granites in the Lachlan Fold Belt southeastern Australia. Philosophical Transactions of the Royal Society, London, A310, 693-707.

THE TASMANIDES OF EASTERN AUSTRALIA CHAPPELL, B.W. 1994. Lachlan and New England: fold belts of contrasting magmatic and tectonic development. Journal & Proceedings of the Royal Society of New South Wales, 127, 47-59. CHAPPELL, B.W. & STEPHENS, W.E. 1988. Origin of infracrustal (I-type) granite magmas. Transactions of the Royal Society of Edinburgh, Earth Sciences, 79, 71-86. CHAPPELL, B.W. & WHITE, A.J.R. 1974. Two contrasting granite types. Pacific Geology, 8,173-174. CHAPPELL, B.W., WHITE, A.J.R. & HINE, R. 1988. Granite provinces and basement terranes in the Lachlan Fold Belt southeastern Australia. Australian Journal of Earth Sciences, 35, 505-521. CHAPPELL, B.W., BRYANT, C.J., WYBORN, D., WHITE, A.J.R. & WILLIAMS, I.S. 1998. High- and lowtemperature I-type granites. Resource Geology, 48, 225-235. CHAPPELL, B.W., WHITE, A.J.R., WILLIAMS, I.S., WYBORN, D. & WYBORN, L.A.I. 2000. Lachlan Fold belt granites revisited: high and low temperature granites and their implications. Australian Journal of Earth Sciences, 47, 123-138. COLLINS, W.J. 1998. Evaluation of petrogenetic models for Lachlan Fold Belt granitoids: implications for crustal architecture and tectonic models. Australian Journal of Earth Sciences, 45, 483-500. COLLINS, W.J. 2002. Nature of extensional accretionary orogens. Tectonics, 21, 10.1029/ 2000TC001272 001276.001271-001276.001279. COLLINS, W.J. & HOBBS, B.E. 2001. What caused the Early Silurian change from mafic to silicic (Stype) magmatism in the eastern Lachlan Fold Belt? Australian Journal of Earth Sciences, 48, 25-41. COLLINS, W.J. & VERNON, R.H. 1992. Palaeozoic arc growth, deformation and migration across the Lachlan Fold Belt, southeastern Australia. Tectonophysics, 214, 381-400. COLLINS, W.J. & VERNON, R.H. 1994. A rift-driftdelamination model of continental evolution: Palaeozoic tectonic development of eastern Australia. Tectonophysics, 235, 249-275. COLLINS, W.J., OFFLER, R., FARRELL, J.R. & LANDENBERGER, B. 1993. A revised Late PalaeozoicEarly Mesozoic tectonic history for the southern New England Fold Belt. In: FLOOD, P.O. & AITCHISON, J.C. (eds) Abstracts of the Conference on the New England Orogen, eastern Australia. University of New England, Armidale, 69-84. COLLINS, C.D.N., DRUMMOND, B.J. & NICOLL, M.G. 2003. Crustal thickness patterns in the Australian continent In: HILLIS, R.R. & MULLER, R.D. (eds.) Evolution and Dynamics of the Australian Plate. Geological Society of Australia, Special Publications, 22 and Geological Society of America, Special Paper, 372,121-128 COLQUHOUN, G.P., FERGUSSON, C.L. & TYE, S.C. 1999. Provenance of early Palaeozoic sandstones, southeastern Australia, Part 2: cratonic to arc switching. Sedimentary Geology, 125, 153-163. COLQUHOUN, G., CAMERON, R., MEAKIN, S., HENDRICKX, M. & VASSALLO, J. 2004. Stratigraphy

81

and structure of the Cargelligo region, central Lachlan Orogen, NSW. Abstracts of the Geological Society of Australia, 73, 150. CONEY, P.J. 1992. The Lachlan belt of eastern Australia and the Circum-Pacific tectonic evolution. Tectonophysics, 214, 1-25. COOPER, J.A., JENKINS, R.J.F., COMPSTON, W. & WILLIAMS, I.S. 1992. lon-microprobe zircon dating of a mid-Early Cambrian tuff in South Australia. Journal of the Geological Society, London, 149,185-192. COOPER, PF & TUCKWELL, K.D. 1971. The Upper PreCambrian Adelaidean of the Broken Hill area - a new subdivision. Quarterly Notes of the Geological Survey of New South Wales, 3, 8-16. COOPER, P.P., TUCKWELL, K.D., GILLIGAN, L.B. & MEARES, R.M.D. 1975. Torrowangee-Fowlers Gap 1:100000 Geological Sheet, Geological Survey of New South Wales, Sydney. COOPER, R.A. 1999. The Ordovician time scale: calibration of graptolite and conodont zones. Acta Universitatis Carolinae, Geologica, 43, 1-4. Cox, S.F., ETHERIDGE, M.A., CAS, R.A.F. & CLIFFORD, B.A. 1991. Deformational style of the Castlemaine area, Bendigo-Ballarat Zone: Implications for evolution of crustal structure in central Victoria. Australian Journal of Earth Sciences, 38, 151-170. CRANFIELD, L.C., SHORTEN, G., SCOTT, M. & BARKER, R.M. 1997. Geology and mineralisation of the Gympie Province. In: ASHLEY, P.M. & FLOOD, P.G. (eds) Tectonics and Metallogenesis of the New England Orogen. Geological Society of Australia, Special Publications, 19,128-147. CRAWFORD, A.J. 1988. Cambrian. In: DOUGLAS, J.G. & FERGUSON, J.A. (eds) Geology of Victoria, 2nd edition. Geological Society of Australia, Victorian Division, Melbourne, 37-62. CRAWFORD, A.J. & BERRY, R.F. 1992. Tectonic implications of Late Proterozoic-early Palaeozoic igneous rock associations in western Tasmania. Tectonophysics, 214, 37-56. CRAWFORD, A.J., CAMERON, W.E. & KEAYS, R.R. 1984. The association boninite low-Ti andesitetholeiite in the Heathcote greenstone belt, Victoria; ensimatic setting for the early Lachlan Fold Belt. Australian Journal of Earth Sciences, 31, 161-174. CRAWFORD, A.J., DONAGHY, A.G., BLACK, L. & STUART-SMITH, P. 1996. Mt. Read Volcanics correlates in western Victoria: a new exploration opportunity. Bulletin Australian Institute of Geoscientists, 20, 97-102. CRAWFORD, A.J., STEVENS, B.P.J. & FANNING, M. 1997. Geochemistry and tectonic setting of some Neoproterozoic and Early Cambrian volcanics in western New South Wales. Australian Journal of Earth Sciences, 44, 831-852. CRAWFORD, A.J., MEFFRE, S. & SYMONDS, P.A. 20030. 120 to 0 Ma tectonic evolution of the southwest Pacific and analogous geological evolution of the 600 to 220 Ma Tasman Fold Belt System. In: HILLIS, R.R. & MULLER, R.D. (eds) Evolution and Dynamics of the Australian Plate. Geological

82

R. A. GLEN

Society of Australia, Special Publications, 22 and Geological Society of America Special Paper, 372, 383-103. CRAWFORD, AJ. (co-ordinator) and 14 others. 20036. Neoproterozoic and Cambrian continental rifting, continent-arc collision and post-collisional magmatism. In: BIRCH, W.D. (ed.) Geology of Victoria. Geological Society of Australia, Special Publications, 23, 73-92. CRAWFORD, AJ. ET AL. 2005. Tectonomagmatic evolution of the Lachlan Fold Belt in the Late Cambrian-Ordovician: west-Pacific type tectonics above a rolling back slab. Unpublished. CROOK, K.A.W. 1960. Petrology of the Tamworth Group, Lower and Middle Devonian, TamworthNundle district New South Wales. Journal of Sedimentary Petrology, 30, 538-552. CROOK, K.A.W. 1969. Contrasts between Atlantic and Pacific geosynclines. Earth and Planetary Science Letters, 5, 429-438. CROOK, K.A.W. 1979. Tectonic implications of some field relationships of the Adelaidean Cooee Dolerite, Tasmania. Journal of the Geological Society of Australia, 26, 353-361. CROOK, K.A.W. 1980. Fore-arc evolution in the Tasman Geosyncline: the origin of southeast Australian continental crust. Journal of the Geological Society of Australia, 27, 215-232. CROOK, K.A.W. & POWELL, C.M. 1976. The evolution of the southeastern part of the Tasman Geosyncline. International Geological Congress 25th Session Sydney Excursion Guide 17A. CROSS, K.C., FERGUSSON, C.L. & FLOOD, P.G. 1987. Contrasting structural styles in the Paleozoic subduction complex of the southern New England Orogen, eastern Australia. In: LEITCH, E.C. & SCHEIBNER, E. (eds) Terrane Accretion and Orogenic Belts. American Geophysical Union Geodynamics Series, 19, 83-92. CROUCH, S.B.S. 1999. Geology, tectonic setting and metallogenesis of the Berserker subprovince, northern New England Orogen. Queensland Government Mining Journal, 100, 6-14. DADD, K.A. 1992. The Middle to Late Devonian Eden-Comerong-Yalwal Volcanic Zone of southeastern Australia: an ancient analogue of the Yellowstone-Snake River Plain region of the USA. Tectonophysics, 214, 277-291. DADD, K.A. 1998. Incipient backarc magmatism in the Silurian Tumut Trough, New South Wales: an ancient analogue of the early Lau Basin. Australian Journal of Earth Sciences, 44,109-121. DALZIEL, I.W.D., DALLA SALDA, L.H. & GAHAGAN, L.M. 1994. Paleozoic Laurentia-Gondwana interaction and the origin of the AppalachianAndean mountain system. Bulletin of the Geological Society of America, 106, 243-252. DAVID, V., GLEN, R.A. & SPENCER, R. 2003. Threedimension extension of the Hill End Trough based on the modelling of the regional gravity data. In: VASSALLO, JJ.A. & GLEN R.A. (eds.) Evolution of the Hill End Trough. Geological Survey of New South Wales, Sydney, Report GS2003/291, 43-48.

DAVIS, B.K. & HENDERSON, R.A. 1999. Syn-orogenic extensional and contractional deformation related to granite emplacement in the northern Tasman Orogenic Zone, Australia. Tectonophysics, 305, 453^75. DAVIS, B.K. HENDERSON, A. & WYSOCZANSKI, R. 1998. Timing of granite emplacement under conditions of low strain in the northern Tasman Orogenic Zone, Australia. Tectonophysics, 284,179-202. DAVIS, B.K., HENDERSON, R.A., LINDSAY, M. & WYSOCZANSKI, R. 2002. Deformation history and timing of granite emplacement in the Butchers Hill-Helenvale region, northern Hodgkinson Province, Queensland. Australian Journal of Earth Sciences, 49, 775-785. DAY, R.W., MURRAY, C.G. & WHITTAKER, W.G. 1978. The eastern part of the Tasman Orogenic Zone. Tectonophysics, 48, 327-364. DIREEN, N.G. 1997. Geophysical investigations of the Koonenberry Belt, western NSW. Record Australian Geological Survey Organisation, 1997/49,10-13. DIREEN, N.G. & CRAWFORD, AJ. 20030. Fossil seaward dipping reflector sequences preserved in southeastern Australia. Journal of the Geological Society, London, 160, 985-990. DIREEN, N.G. & CRAWFORD, AJ. 2003ft. The Tasman Line: where is it, what is it, and is it Australia's Rodinian breakup boundary? Australian Journal of Earth Sciences, 50, 491-502. DIRKS, P.H., HAND, M., COLLINS, WJ. & OFFLER, R. 1992. Structural-metamorphic evolution of the Tia Complex, New England fold belt: thermal overprint of an accretion-subduction complex in a compressional back arc setting. Journal of Structural Geology, 14, 669-688. DOMAGALA, J., BULTITUDE, R.J., DONCHAK, P.J.T.,

FORDHAM, B.G,. JELL, J. & FANNING, M. 1998. Tectonic setting implications from greywacke provenance data from the Hodgkinson province. Abstracts of the Geological Society of Australia, 49,116. DONCHAK, P.J.T. 1993. Tectonic Evolution. In: BULTITUDE, R.J., DONCHAK, P.J.T., DOMAGALA J. & FORDHAM, B.G. (eds) The Pre-Mesozoic Stratigraphy and Structure of the Western Hodgkinson Province and Environs. Queensland Geological Record, Department of Minerals and Energy Queensland, 1993/20,176-185. DRAPER, J.J., BAIN, J.H.V., WITHNALL, I.W., BLEWETT, R.S., BLACK, L.P., BULTITUDE, RJ. & WELLMAN, P. 1998. Proterozoic rocks of north Queensland. Abstracts of the Geological Society of Australia, 49,123. DUGGAN, M.B. 2000. Narragudgil Volcanics. Forbes 1:250 000 Geological Sheet SI 55-7, 2nd Edition, Explanatory Notes. Records of the Australian Geological Survey Organisation, Canberra, 20, 37-39. ELLIOTT, C.G. & GRAY, D.R. 1992. Correlations between Tasmania and the Tasman-Transantarctic Orogen; evidence for easterly derivation of Tasmania relative to mainland Australia. Geology, 20, 621-623.

THE TASMANIDES OF EASTERN AUSTRALIA ELLIOTT, SJ.A. & MARTIN, A.R. 19910. Geology and mineralisation of the Fifield Platinum Province New South Wales. 6th International Platinum Symposium. Guidebook for the pre-symposium field excursion. Geological Society of Australia and Vanguard Press, Sydney. EVANS, P.R., HOFFMAN, K.L., REMUS, D.A. & PASSMORE, V.L. 1990. Geology of the Eromanga sector of the Eromanga-Brisbane Geoscience Transect. In: FINLAYSON, D.M. (ed.) The Eromanga-Brisbane Geoscience Transect: a guide to Basin Development across Phanerozoic Australia in Southern Queensland. Bulletin of the Bureau of Mineral Resources, Canberra, 232, 83-104. FABRIS, A. 2003. Northwestern Murray basin basement. MESA Journal (Quarterly Earth Resources Journal of Primary Industries and Resources South Australia), 30, July 2003, 39-45. FAGAN, R.K. 1979. S-type granite genesis and emplacement in N.E. Victoria and its implications. Records of the Bureau of Mineral Resources, 1972/2. FANNING, C.M. 1996. Geochronological synthesis of southern Australia. Envelope, Department of Mines and Energy South Australia, 8918. FANNING, C.M., LUDWIG, K.R., FORBES, B.G. & PREISS, W.V. 1986. Single and multiple grain U-Pb zircon analyses for the early Adelaidean Rook Tuff, Willouran Ranges, South Australia. Abstracts of the Geological Society of Australia, 15, 71-72. FANNING, C.M., LEITCH, E.G. & WATANABE, T. 2002. An updated assessment of the SHRIMP U-Pb Zircon dating of the Attunga eclogite in New South Wales, Australia: Relevance to the Pacific margin of Gondwana. In: International Symposium on the Amalgamation of Precambrian Blocks and the Role of the Paleozoic Orogens in Asia. Hokkaido, Japan, Abstract Volume with supplementary data. FEARY, D.A., CHAMPION, D.C., BULTITUDE, R.J. & DAVIES, P.J. 1993. Igneous and metasedimentary basement lithofacies of the Queensland Plateau. Proceedings of the Ocean Drilling Program, Scientific Results, 133, 535-540. FELTON, E.A. 1976. Geology of the Braidwood 1:100 000 sheet. Geological Survey of New South Wales, Sydney. FERGUSSON, C.L. 1985. Trench floor sedimentary sequences in a Palaeozoic subduction complex. Sedimentary Geology, 42, 181-200. FERGUSSON, C.L. 1987. Early Paleozoic back-arc deformation in the Lachlan Fold Belt southeastern Australia: Implications for terrane translations in Eastern Gondwanaland. In: LEITCH, E.G. & SCHEIBNER, E. (eds) Terrane Accretion and Orogenic Belts. American Geophysical Union, Geodynamic Series, 19, 39-56. FERGUSSON, C.L. 1991. Thin-skinned thrusting in the northern New England Orogen, central Queensland, Australia. Tectonics, 10, 797-806. FERGUSSON, C.L. 1998. Cambrian-Silurian oceanic rocks, upper Howqua River, eastern Victoria: tectonic implications. Australian Journal of Earth Sciences, 45, 633-644.

83

FERGUSSON, C.L. 2003. Ordovician-Silurian accretion tectonics of the Lachlan Fold Belt, southeastern Australia. Australian Journal of Earth Sciences, 50, 475^90. FERGUSSON, C.L. & CONEY, P.J. 19920. Convergence and intraplate deformation in the Lachlan Foldbelt of southeastern Australia. Tectonophysics, 214, 411-439. FERGUSSON, C.L. & CONEY, P.J. 1992Z?. Implications of a Bengal Fan-type deposit in the Paleozoic Lachlan fold belt of southeastern Australia. Geology, 20, 1047-1049. FERGUSSON, C.L. & FANNING, C.M. 2002. Late Ordovician stratigraphy, zircon provenance and tectonics, Lachlan Fold Belt, southeastern Australia. Australian Journal of Earth Sciences, 49, 423-436. FERGUSSON, C.L. & GLEN, R.A. (eds). 1992. The Palaeozoic eastern margin of Gondwanaland: tectonics of the Lachlan Fold Belt, southeastern Australia and related orogens. Tectonophysics, 214. FERGUSSON, C.L. & FRIKKEN, D. 2003. Diapirism and structural thickening in an Early Palaeozoic subduction complex, southern New South Wales, Australia. Journal of Structural Geology, 25, 43-58. FERGUSSON, C.L., HENDERSON, R.A., LEITCH, E.G. & ISHIGA, H. 1993. Lithology and structure of the Wandilla terrane, central Queensland, and an overview of the Palaeozic subduction complex of the New England Fold Belt. Australian Journal of Earth Sciences, 40, 403-414. FERGUSSON, C.L., HENDERSON, R.A. & WRIGHT, J.V. 1994. Facies in a Devonian-Carboniferous volcanic fore-arc succession, Campywn Volcanics, Mackay district, central Queensland. Australian Journal of Earth Sciences, 41, 287-300. FERGUSSON, C.L., CARR, P.F., FANNING, C.M. & GREEN, T.J. 2001. Proterozoic-Cambrian detrital zircons and monzonite ages from the Anakie Inlier central Queensland: Grenville and PacificGondwana signatures. Australian Journal of Earth Sciences, 48, 857-866. FERM, J.C. & STAUB, J.R. 1984. Depositional controls of mineable coal bodies. In: RAHMANI, R.A. (ed.) Sedimentology of Coal and Coal-bearing Sequences. Special Publication, International Association of Sedimentologists, 7, 275-289. FIELDING, C.R., STEPHENS, C.J. & HOLCOMBE, R.J. 1997. Permian stratigraphy and palaeogeography of the eastern Bowen Basin, Gogango Overfolded Zone and the Strathmuir Synclinorium in the Rockhampton-Mackay region, central Queensland. In: ASHLEY, P.M. & FLOOD, P.G. (eds) Conference on the Tectonics and Metallogenesis of the New England Orogen. Geological Society of Australia, Special Publication, 19, 80-95. FIELDING, C.R., SLIWA, R., HOLCOMBE, R.J. & JONES, J.T. 2001. A new palaeogeographic synthesis for the Bowen, Gunnedah and Sydney basins of eastern Australia. In: BERECHER, T. & HILL, K. (eds) PESA Eastern Australasian Basins

84

R. A. GLEN

symposium. Australasian Institute of Mining and Metallurgy, Melbourne, 269-279. FINLAYSON, D.M. 1990. The Eromanga-Brisbane geoscience transect: a guide to basin development across Phanerozoic Australia in southern Queensland. Bulletin 232, Bureau of Mineral Resources, Canberra. FINLAYSON, D.M., WAKE-DYSTER, K.D., LEVEN, J.H. ET AL. 1990. Seismic imaging of major tectonic features in the crust of Phanerozoic Australia. Tectonophysics, 173, 211-230. FLOOD, P.G. 1999. Exotic seamounts within Gondwanan accretionary complexes, Eastern Australia. In: FLOOD, P.G. (ed.) Conference on Regional Geology, Tectonics and Metallogenesis, New England Orogen. University of New England, Armidale, 23-29. FLOOD, P.G. & AITCHISON, J.C. 1992. Late Devonian accretion of the Gamilaroi Terrane to eastern Gondwana: provenance linkage suggested by the first appearance of Lachlan Fold Belt-derived quartzarenite. Australian Journal of Earth Sciences, 39, 539-544. FLOTTMANN, T. & JAMES, P.R. 1997. Influence on basin architecture on the style of inversion and fold-thrust belt tectonics - the southern Adelaide Fold-Thrust Belt, South Australia. Journal of Structural Geology, 19,1093-1110. FLOTTMAN, T., JAMES, P., ROGERS, J. & JOHNSON, T. 1994. Early Palaeozic foreland thrusting and basin reactivation at the palaeo-Pacific margin of the southeastern Australian Precambrian craton: a reappraisal of the structural evolution of the southern Adelaide Fold-Thrust Belt. Tectonophysics, 234, 95-116. FLOTTMANN, T., HAINES, P.W., COCKSHELL, C.D. & PREISS, W.V. 1997. An early Palaeozoic foreland basin succession beneath Gulf St Vincent, South Australia? - implications for petroleum plays. PESA Journal, 25, 33-40. FLOTTMANN, T., HAINES, P., JAGO, J., JAMES, P., BELPERIO, A. & GUM, J. 1998. Formation and reactivation of the Cambrian Kanmantoo Trough, SE Australia: implications for early Palaeozoic tectonics at eastern Gondwana's plate margin. Journal of the Geological Society, London, 155, 525-539. FODEN, J.D., ELBURG, M.A., TURNER, S.P., SANDIFORD, M., O'CALLAGAHAN, J.O. & MITCHELL, S.

2002. Granite production in the Delamerian Orogen, South Australia. Journal of the Geological Society, London, 159, 557-575. FORDHAM, B.G. 1993. Chillagoe Formation. In: BULTITUDE, R.J., DONCHAK, P.J.T., DOMAGALA,

J. & FORDHAM, B.G. (eds) The Pre-Mesozoic Stratigraphy and Structure of the Western Hodgkinson Province and Environs. Queensland Geological Record, 1993/29, 40-51. FORDHAM, B.G., MURRAY, C.G., HAYWARD, M.A., SIMPSON, G.A. & BLAKE, P.R. 1998. Basin-history of the Yarrol Block: steady tectonic subsidence and repeated forearc development from Silurian to Permian. Abstracts of the Geological Society of Australia, 49, 151.

FOSTER, D.A. & GRAY, D.R. 2000. Evolution and structure of the Lachlan Fold Belt (Orogen) of eastern Australia. Annual Review of Earth and Planetary Sciences, 28, 47-80. FOSTER, D.A., GRAY, D.R. & BUCHER, M. 1999. Chronology of deformation within the turbiditedominated Lachlan Orogen: Implications for the tectonic evolution of eastern Australia and Gondwana. Tectonics, 18, 452^85. FUKUI, S., WATANABE, T., ITAYA, T. & LEITCH, E.C. 1995. Middle Ordovician high PT metamorphic rocks in eastern Australia: evidence from K-Ar ages. Tectonics, 14,1014-1020. FUREY-GREIG, T. 1996. The age of forearc basin and subduction complex rocks of the New England fold belt: new conodont data and revised alignments. In: KEENE, J. (ed.) The Geological Evolution of Eastern Australia. Earth Resources Foundation, University of Sydney, Sydney, 18-19. FUREY-GREIG, T. 1999. Late Ordovician conodonts from the olistostromal Wisemans Arm Formation (New England Region, Australia). Abhandlungen der Geologischen Bundesanstalt, 54, 303-321. GARRATT, MJ. 1983. Silurian to Early Devonian facies and biofacies patterns for the Melbourne Trough, central Victoria. Journal of the Geological Society of Australia, 30,121-147. GATEHOUSE, C.G. 1986. The geology of the Warburton Basin in South Australia. Australian Journal of Earth Sciences, 33, 161-180. GEEVE, R.I, SCHMIDT, P.W. & ROBERTS, J. 2002. Paleomagnetic results indicate pre-Permian counterclockwise rotation of the southern Tamworth Belt, southern New England Orogen, Australia. Journal of Geophysical Research, 107(B9), 2196 doi:10.1029/2000JB000037. GIBSON, G.M. 1998. Initiation of Rodinia breakup in southeast Australia by SW-NE extension: evidence from shear zone kinematics and dyke intrusion. Abstracts of the Geological Society of Australia, 50, 37-38. GIBSON, G.M. & NIHILL, D.N. 1992. Glenelg River Complex: western margin of the Lachlan Fold Belt or extension of the Delamerian Orogen into western Victoria. Tectonophysics, 214, 69-91. GLAZNER, A.F., BARTLEY, J.M., COLEMAN, D.S., GRAY, W. & TAYLOR, R.Z. 2004. Are plutons assembled over millions of years by amalgamation from small magma chambers? GSA Today, 14, 4-11. GLEN, R.A. 1991. Inverted transtensional basin setting for gold and copper and base metal deposits at Cobar, New South Wales. BMR Journal of Geology and Geophysics, 12, 13-24. GLEN, R.A. 1992. Thrust, extensional and strike-slip tectonics in an evolving Palaeozoic orogen - a structural synthesis of the Lachlan Orogen of southeastern Australia. Tectonophysics, 214, 341-380. GLEN, R.A. 1993. Palaeomagnetism and terranes in the Lachlan Orogen. Exploration Geophysics, 24, 247-256. GLEN, R.A. 1994. Ordovician-earliest Silurian

THE TASMANIDES OF EASTERN AUSTRALIA sedimentary rocks. In: LEWIS, P.C., GLEN, R.A., PRATT, G.W. & CLARKE, I. (eds) Bega-Mallacoota 1:250000 Geological Sheet SI/55-4, SJ/55-8: Explanatory Notes. Geological Survey of New South Wales, Sydney, 17-27. GLEN, R.A. 1995. Thrust and thrust-associated mineralization in the Lachlan Orogen. Economic Geology, 90,1402-1429. GLEN, R.A. 1998. Tectonic development of the Lachlan Orogen: a framework for mineral exploration. In: LEWIS, P.C. (ed.) Lachlan Fold Belt Conference '98. Bulletin of the Australian Institute of Geoscientists, 23,1-6. GLEN, R.A. 1999. Structure. In: MEAKIN, N.S. & MORGAN, EJ. (eds) Dubbo 1:250 000 Geological Sheet SI/55-4, 2nd Edn. Explanatory Notes, Geological Survey of New South Wales, Sydney, 266-392. GLEN, R.A. 2004. Plate tectonics of the Lachlan Orogen: a framework for understanding its metallogenesis. Abstracts of the Geological Society of Australia, 74, 33-36. GLEN, R.A. & BECKETT, J. 1997. Structure and tectonics along the inner edge of a foreland basin: the Hunter Coalfield in the northern Sydney Basin, New South Wales. Australian Journal of Earth Sciences, 44, 853-877. GLEN, R.A. & BROWN, R.E. 1995. Fault-related mineralisation in the Teatree Creek goldfield, Tamworth Belt, New England Orogen. Quarterly Notes of the Geological Survey of New South Wales, 97,1-13. GLEN, R.A. & PERCIVAL, I.G. 2002. In: STOREY, B. (ed.) Ordovician Terranes in the Lachlan sector of Gondwana. Abstracts of Gondwana 11, University of Canterbury, Christchurch, 41. GLEN, R.A. & PERCIVAL, I.G. 2003. Convergent and transform boundaries along the SE Australian sector of the proto-Pacific Gondwana margin in the Ordovician. In: FUETTERER, D.K. (ed.) 9th International Symposium on Antarctic Earth Sciences, Potsdam, 120-121. GLEN, R.A. & PERCIVAL, I.G. 2005. Convergent and transform boundaries along the SE Australian sector of the proto-Pacific Gondwana margin in the Ordovician. Unpublished. GLEN, R.A. & SCHEIBNER, E. 1993. Lachlan rocks in the New England Orogen. In: FLOOD, P.G. & AITCHISON, J.C. (eds) Abstracts of the Conference on the New England Orogen, Eastern Australia, University of New England, Armidale, 123-126. GLEN, R.A. & VANDENBERG, A.H.M. 1996. Assessing the thin-skinned paradigm in the Lachlan Orogen: tectonic and economic implications of 13 years of thrusting. Bulletin of the Australian Institute of Geoscientists, 20, 25-31. GLEN, R.A. & VANDENBERG, A.H.M. 1987. Thinskinned tectonics in part of the Lachlan Fold Belt near Delegate, southeastern Australia. Geology, 15, 1070-1073. GLEN, R.A. & WALSHE, J.L. 1999. Cross structures in the Lachlan Orogen: The Lachlan Transverse Zone example. Australian Journal of Earth Sciences, 46, 641-658.

85

GLEN, R.A. & WATKINS, J.J. 1999. Implications of Middle Devonian deformation of the eastern part of the Hill End Trough, Lachlan Orogen. Australian Journal of Earth Sciences, 46, 35-52. GLEN, R.A. & WYBORN, D. 1997. Inferred thrust imbrication, deformation gradients and the Lachlan Transverse Zone in the eastern belt of the Lachlan Orogen, New South Wales. Australian Journal of Earth Sciences, 44, 49-68. GLEN, R.A., STEWART, I.R. & VANDENBERG, A.H.M. 1990. Imbrication of a reference section: Reevaluation of the Adaminaby Beds at El Paso, Dalgety, New South Wales. Journal & Proceedings Royal Society of New South Wales, 123,15-26. GLEN, R.A., DALLMEYER, R.D. & BLACK, L.P. 19920. Isotopic dating of basin inversion - the Palaeozoic Cobar Basin in the Lachlan Orogen, Australia. Tectonophysics, 214, 249-268. GLEN, R.A., SCHEIBNER, E. & VANDENBERG, A.H.M. 19926. Paleozoic intraplate escape tectonics in Gondwanaland and major strike-slip duplication in the Lachlan Orogen of southeastern Australia. Geology, 20, 795-798. GLEN, R.A., KORSCH, RJ. & WAKE-DYSTER, K.D. 1993. A deep seismic cross section through the Tamworth belt: preliminary interpretation of 4 seconds two-way time data. In: FLOOD, P.G. & AITCHISON, J.C. (eds) Abstracts of the Conference on the New England Orogen, Eastern Australia. University of New England, Armidale, 102-104. GLEN, R.A., CLARE A. & SPENCER, R. 1996. Extrapolating the Cobar Basin model to the regional scale: Devonian basin-formation and inversion in western New South Wales. In: COOK, W.G., FORD, A.J.H., MCDERMOTT, J.J., STANDISH, P.N., STEGMAN, C.L. & STEGMAN, T.M. (eds) The Cobar Mineral Field - 1996. Australasian Institute of Mining and Metallurgy, Melbourne, Spectrum Series 3/96, 43-83. GLEN, R.A., WALSHE, J.L., BARRON, L.M. & WATKINS, J.J. 1998. Ordovician convergent-margin volcanism and tectonism in the Lachlan sector of east Gondwana. Geology, 26, 751-754. GLEN, R.A., LENNOX, P.G. & FOSTER, D.A. 1999. 40 Ar-39Ar dating of deformations west of the Hill End Trough, Lachlan Orogen, New South Wales. Quarterly Notes of the Geological Survey of New South Wales, 110, 13-22. GLEN, R.A., KORSCH, R.J., DIREEN, N.G. ET AL. 2002. Crustal structure of the Ordovician Macquarie Arc, Eastern Lachlan Orogen, based on seismicreflection profiling. Australian Journal of Earth Sciences, 49, 323-348. GLEN, R.A., CRAWFORD, AJ. & COOKE, D.R. 2003. Tectonic setting of porphyry copper-gold mineralisation in the Macquarie Arc. In: BLEVIN, P.L., JONES, M. & CHAPPELL, B. (eds) Magmas to Mineralisation: the Ishihara Symposium. Geoscience Australia Record, 2003/14, 65-68. GLEN, R.A., STEWART, I.R. & PERCIVAL, I.G. 20040. Narooma Terrane: implications for the construction of the outboard part of the Lachlan Orogen. Australian Journal of Earth Sciences, 51, 859-884. GLEN, R.A., ZHANG, W., DAWSON, M. & NEEDHAM, J.

86

R. A. GLEN

20046. Solid geology of the Eastern Lachlan Orogen, New South Wales, Australia. GIS compilation. In: Geological Survey of New South Wales. Geological Survey of New South Wales, Sydney. GLEN, R.A., SCOTT, RJ. & MEFFRE, S. 2005. The Benambran Orogeny: plate rearrangement and collision of the Macquarie Arc to its backarc basin. Unpublished. GOODGE, J.W. 2002. Neoproterozoic to Early Paleozoic record of Rodinia breakup and Gondwanamargin convergence in east Antarctica. Abstracts of the Geological Society of Australia, 67, 120. GRAVESTOCK, D.I. & GATEHOUSE, C.G. 1995. Eastern Warburton Basin. In: DREXEL, J.F. & PREISS, W.V. (eds) The Geology of South Australia. Volume 2: The Phanerozoic. Bulletin of the Geological Survey of South Australia, 54, 31-34. GRAY, C.M. 1984. An isotopic mixing model for the origin of granitic rocks in southeastern Australia. Earth and Planetary Science Letters, 70, 47-60. GRAY, CM., KEMP, A.I.S., ANDERSON, J.A.C.,BUSHELL, D.J., FERGUSON, D.J., FITZHERBERT, J. & STEVENSON, M.D. 2002. Delamerian Glenelg tectonic zone, western Victoria: geology and metamorphism of stratiform rocks. Australian Journal of Earth Sciences, 49, 187-200. GRAY, D.R. 1988. Structure and Tectonics. In: DOUGLAS, J.D. & FERGUSON, J.A. (eds) Geology of Victoria. Geological Society of Victoria, Victorian Division, Melbourne, 1-36. GRAY, D.R. 1995. Thrust kinematics and transposition fabrics from a basal detachment zone, eastern Australia. Journal of Structural Geology, 17, 1637-1654. GRAY, D.R. & FOSTER, D.A. 1997. Orogenic concepts - application and definition: Lachlan Fold Belt, eastern Australia. American Journal of Science, 297, 859-891. GRAY, D.R. & FOSTER, D.A. 1998. Character and kinematics of faults within the turbidite-dominated Lachlan Orogen: implications for the structural evolution of eastern Australia. Journal of Structural Geology, 20,1691-1720. GRAY, D.R. & FOSTER, D.A. 2000. Character and kinematics of faults within the turbidite-dominated Lachlan Orogen: implications for the tectonic evolution of eastern Australia. Reply. Journal of Structural Geology, 22, 529-535. GRAY, D.R. & WILLMAN, C.E. 19910. Deformation in the Ballarat Slate Belt, central Victoria, and implications for crustal structure across southeast Australia. Australian Journal of Earth Sciences, 38, 171-201. GRAY, D.R. & WILLMAN, C.E. 1991ft. Thrust-related strain gradients and thrusting mechanisms in a chevron-folded sequence, southeastern Australia. Journal of Structural Geology, 13, 691-710. GRAY, D.R., FOSTER, D.A. & BUCHER, M. 1997. Recognition and definition of orogenic events in the Lachlan Fold Belt. Australian Journal of Earth Sciences, 44, 489-501. GRAY, D.R. (co-ordinator) and ten others. 2003. Structure, metamorphism, geochronology and tectonics of Palaeozoic rocks. Geology of

Victoria. In: BIRCH, W.D. (ed.) Geology of Victoria. Geological Society of Australia, Special Publication (Victoria division), 23, 17-70. GUNN, PI,

MlLLIGAN, PR., MACKEY, T, LlU, S.,

MURRAY, A., MAIDMENT, D. & HAREN, R. 1997. Geophysical mapping using the national airborne and gravity datasets: an example focusing on Broken Hill. AGSO Journal of Australian Geology and Geophysics, 17, 127-136. GUST, D.A., STEPHENS, C.J. & GRENFELL, A.T. 1993. Granitoids of the northern NEO: their distribution in time and space and their tectonic implications. In: FLOOD, P.G. & AITCHISON, J.C. (eds) Conference on New England Orogen, eastern Australia. University of New England, Armidale, 557-564. HAINES, P.W. & FLOTTMANN, T. 1998. Delamerian Orogeny and potential foreland sedimentation: a review of age and stratigraphic constraints. Australian Journal of Earth Sciences, 45, 559-570. HAINES, P.W, JAGO, J.B. & GUM, J.C. 2001. Turbidite deposition in the Cambrian Kanmantoo Group, South Australia. Australian Journal of Earth Sciences, 48, 465^78. HANDLER, M.R., BENNETT, V.C. & ESAT, T.M. 1997. The persistence of off-cratonic lithospheric mantle: OS isotopic systematics of variably metasomatised southeast Australian xenoliths. Earth and Planetary Science Letters, 151, 61-75. HARRINGTON, H.J. 1974. The Tasman Geosyncline in Australia. In: DENMEAD, A.K., TWEEDALE, G.W. & WILSON, A.F. (eds) The Tasman geosyncline- a symposium in honour of Professor Dorothy Hill. Geological Society of Australia, Queensland Division, Brisbane, 383-407. HARRINGTON, H.J. 1983. Correlation of the Permian and Triassic Gympie Terrane of Queensland with the Brook Street and Maitai terranes of New Zealand. Permian Geology of Queensland. Geological Society of Queensland, Queensland Division, Brisbane, 431—436. HARRINGTON, H.J., BRAKEL, A.T, HUNT, J.W, WELLS, A.T, MIDDLETON, M.F.,O'BRIEN, P.E., HAMILTON, D.S., BECKETT, J., WEBER, C.R., RADKE, S., TOTTERDELL, J.M., SWAINE, D.J. & SCHMIDT, P.W.

1989. Permian Coals of Eastern Australia. Bureau of Mineral Resources Bulletin, 231. HENDERSON, R.A. 1986. Geology of the Mt. Windsor Subprovince - a Lower Palaeozoic volcanosedimentary terrane in the northern Tasman Orogenic Zone. Australian Journal of Earth Sciences, 33, 343-364. HENDERSON, R.A. 1987. An oblique subduction and transform faulting model for the evolution of the Broken River Province, northern Tasman Orogenic Zone. Australian Journal of Earth Sciences, 34, 237-249. HENDERSON, R.A., DAVIS, B.K. & FANNING, C.M. 1998. Stratigraphy, age relationships and tectonic setting of rift-phase infill in the Drummond Basin, central Queensland. Australian Journal of Earth Sciences, 45, 579-595. HENDERSON, R., FERGUSSON, G, WITHNALL, I., LEWTHWAITE, K. & FANNING, M. 2004. Early

THE TASMANIDES OF EASTERN AUSTRALIA Palaeozoic extension in the Northern Tasmanides. Abstracts of the Geological Society of Australia, 73,165. HILL, D. 1951. Geology. In: MACK, G. (ed.) Handbook of Queensland. Australian Association for the Advancement of Science, Brisbane, 13-24. HILL, K.C. & HALL, R. 2003. Mesozoic-Cenozoic evolution of Australia's New Guinea margin in a west Pacific context. In: HILLIS, R J. & MULLER, R.D. (eds) Evolution and Dynamics of the Australian Plate. Geological Society of Australia, Special Publication, 22 and Geological Society of America, Special Paper, 372, 265-290. HOLCOMBE, R.J., STEPHENS, C.J., FIELDING, C.R. ETAL. 1991 a. Tectonic evolution of the northern New England Fold Belt: the Permian-Triassic HunterBowen event. In: ASHLEY, P.M. & FLOOD, P.G. (eds) Tectonics and Metallogenesis of the New England Orogen. Geological Society of Australia Special Publication, 19, 52-65. HOLCOMBE, R.J., STEPHENS, C.J., FIELDING, C.R. ETAL. I991b. Tectonic evolution of the northern New England Fold Belt: Carboniferous to Early Permian transition from active accretion to extension. In: ASHLEY, P.M. & FLOOD, P.G. (eds) Tectonics and Metallo genesis of the New England Orogen. Geological Society of Australia Special Publication, 19, 66-79. HOLM, O.H. & BERRY, R.S. 2002. structural history of the after lineament, northwestern Tasmania: an analysis of critical outcrops. Australian Journal of Earth Sciences, 49,167-185. HOLM, O.H., CRAWFORD, A.J. & BERRY, R.F. 2003. Geochemistry and tectonic setting of metaigneous rocks in the Arthur Lineament and surrounding area, northwest Tasmania. Australian Journal of Earth Sciences, 50, 903-918. HUTTON, L.J., RlENKS, I.P., TENISON-WOODS, K.L.,

HARTLEY, J.S. & CROUCH, S.B. 1994. Geology of the Ravenswood Batholith, North Queensland. Queensland Geological Record, 1994/4. HUTTON, L., FANNING, C.M. & WITHNALL, I.W. 1998. the Cape River area-evidence for Late Mesoproterozoic and Neoproterozoic crust in north Queensland. Abstracts of the Geological Society of Australia, 48, 216. HUTTON, L.J., WITHNALL, I.W., RIENKS, I.P. ETAL. 1999. A preliminary Carboniferous to Permian magmatic framework for the Auburn and Connors Arches, central Queensland. In: FLOOD, P.G. (ed.) Regional Geology, Tectonics and Metallo genesis, New England Orogen. University of New England, Armidale, 223-232. IRELAND, T.R., Mum, R.J., WEAVER, S.D. & ADAMS, C.J. 1994. Zircon age distributions in granites, greywackes and gneisses from the southwest Pacific Gondwana region. US Geological Survey Circular, 1107, 151. IRELAND, T.R., FLOTTMANNN, T, FANNING, C.M., GIBSON, G.M. & PREISS, W.V. 1998. Development of the early Paleozoic Pacific margin of Gondwana from detrital zircon ages across the Delamerian Orogen. Geology, 26, 243-246. ISHIGA, H., LEITCH, E.C., WATANABE, T, NAKA, T. &

87

IWASAKI, M. 1988. Radiolarian and conodont biostratigraphy of siliceous rocks from the New England Fold Belt. Australian Journal of Earth Sciences, 35, 73-80. JAGO, J., GUM, J.C., BURTT, A.C & HAINES, P.W. 2003. Stratigraphy of the Kanmantoo Group: a critical element of the Adelaide Fold Belt and the Palaeo-Pacific plate margin, Eastern Gondwana. Australian Journal of Earth Sciences, 50, 343-363. JENKINS, R.B., LANDENBERGER, B. & COLLINS, WJ. 2002. Late Palaeozoic retreating and advancing subduction boundary in the New England Fold Belt, New South Wales. Australian Journal of Earth Sciences, 49, 467^89. JENKINS, R.J.F. & SANDIFORD, M. 1992. Observations on the tectonic evolution of the Southern Adelaide Fold Belt. Tectonophysics, 214, 27-36. JOHNSON, S.E. & HENDERSON, R.A. 1991. Tectonic development of the Drummond Basin, eastern Australia: backarc extension and inversion in a Late Palaeozoic active margin setting. Basin Research, 3, 197-213. JOHNSTON, A.J., OFFLER, R. & Liu, S. 2002. Structural fabric evidence for indentation tectonics in the Nambuca Block, southern New England Fold Belt, New South Wales. Australian Journal of Earth Sciences, 49, 407-421. JONES, B.G., FERGUSSON, C.L. & ZAMBELLI, P.F. 1993. Ordovician contourites in the Lachlan Fold Belt, eastern Australia. Sedimentary Geology, 82, 257-270. JONES, J.A., CARR, P.S., FERGUSSON, C.L. & McDouGALL, I. 1995. Silurian volcanism in the Wollondilly Basin, eastern Lachlan Fold Belt, New South Wales. Australian Journal of Earth Sciences, 42, 25-34. JONES, J.G. & MCDONNELL, K.L. 1981. Papua New Guinea analogue for the late Permian environment of northeastern New South Wales. Palaeogeography, Palaeoclimatology, Palaeoecology, 34, 191-205. JONES, J.G, CONAGHAN, P.J., MCDONNELL, K.L., FLOOD, R.H. & SHAW, S.E. 1984. Papuan Basin analogue and a foreland basin model for the Bowen-Sydney Basin. In: VEEVERS, J.J. (ed.) Phanerozoic Earth History of Australia. Oxford University Press, Oxford, 243-262. KAY, J.R. 1998. The Palaeozoic tectonic evolution of north Queensland, and associated metalliferous mineralisation. Abstracts of the Geological Society of Australia, 48, 235. KEAY, S., COLLINS, WJ. & MCCULLOCH, M.T. 1997. A three-component Sr-Nd isotopic mixing model for granitoid genesis, Lachlan fold belt, eastern Australia. Geology, 25, 307-310. KEEP, M. 2003. Physical modelling of deformation in the Tasman Orogenic Zone. Tectonophysics, 375, 37-47. KEMP, A.I.S., GRAY, CM., ELLIS, D.J., ANDERSON, J.A.C. & FERGUSON, D.J. 2002. Delamerian Glenelg tectonic zone, western Victoria: characterisation and synthesis of igneous rocks. Australian Journal of Earth Sciences, 49,201-224. KIMBROUGH, D.L., CROSS, K.C. & KORSCH, R.J. 1993.

88

R. A. GLEN

U-Pb ages for zircons from the Pola Fogal and Nundle granite suites, Southern New England Orogen. In: FLOOD P.O. & AITCHISON, J.C. (eds) Conference on the New England Orogen, eastern Australia. University of New England, Armidale, 403-412. KLOOTWIJK, C.T. 2002. Carboniferous plaeomagnetism of the Rocky Creek Block, northern Tamworth Belt and the New England pole path. Australian Journal of Earth Sciences, 49, 375^1-05. KORSCH, RJ. 1982. Early Permian tectonic events in the New England Orogen. In: RUNNEGAR, B. (ed.) Voisey Symposium. University of New England, Armidale, 35^2. KORSCH, RJ. & HARRINGTON, H.J. 1981. Stratigraphic and structural synthesis of the New England Orogen. Australian Journal of Earth Sciences, 28, 205-226. KORSCH, RJ. & HARRINGTON, H.J. 1987. Oroclinal bending, fragmentation and deformation of terranes in the New England Orogen. In: LEITCH, E.G. & SCHEIBNER, E. (eds) Terrane Accretion and Orogenic Belts. American Geophysical Union Geodynamics Series, 19,129-139. KORSCH, RJ., WAKE-DYSTER, K.D. & JOHNSTONE, D.W. 1993. The Gunnedah Basin - New England Orogen deep seismic reflection profile: implications for New England tectonics. In: FLOOD, P.G. & AITCHISON, J.C. (eds) Abstract Volume, Conference on the New England Orogen, Eastern Australia. University of New England, Armidale, 85-100. KORSCH, RJ., JOHNSTONE, D.W. & WAKE-DYSTER, K.D. 1997. Crustal architecture of the New England Orogen based on deep seismic reflection profiling, In: ASHLEY, P.M. & FLOOD, P.G. (eds) Tectonics and Metallogenesis of the New England Orogen. Geological Society of Australia, Special Publication, 19, 29-51. KORSCH, R J, BARTON, T.J., GRAY, D.R., OWEN, AJ. & FOSTER, D.A. 2002. Geological interpretation of the deep seismic-reflection transect across the boundary between the Delamerian and Lachlan Orogens, in the vicinity of the Grampians, Western Victoria. Australian Journal of Earth Sciences, 49,1057-1075. LAMBERT, I.B. & WHITE, AJ.R. 1965. The Berridale Wrench Fault: a major structure in the Snowy Mountains of New South Wales. Journal of the Geological Society of Australia, 12, 25-33. LANDENBERGER, B., FARRELL, J.R., OFFLER, R., COLLINS, WJ. & WHITFORD, DJ. 1995. Tectonic implications of Rb-Sr biotite ages for the Hillgrove Plutonic Suite, New England Fold Belt, NSW, Australia. Precambrian Research, 71, 251-263. LANPHERE, M.A. & HOCKLEY, J. 1976. The age of nephrite occurrences in the Great Serpentine Belt of New South Wales. Journal of the Geological Society of Australia, 21,133-156. LAWYER, L.A. & GAHAGAN, L.M. 1994. Constraints on timing of extension in the Ross sea region. Terra Antarctica, 1, 545-552. LEAMAN, D.E. 1994. The Tamar fracture system: does

it exist? Australian Journal of Earth Sciences, 41, 73-74. LEITCH, E.G. 1969. Igneous activity and diastrophism in the Permian of New South Wales. Special Publication of the Geological Society of Australia, 2, 21-37. LEITCH, E.G. 1974. The geological development of the southern part of the New England Fold Belt. Journal of the Geological Society of Australia, 21, 133-156. LEITCH, E.G. 1988. The Barnard Basin and Early Permian development of the southern part of the New England Fold Belt. In: KLEEMAN, J.D. (ed.) New England Orogen - Tectonics and Metallogenesis. University of New England, Armidale 61-67. LEITCH, E.G. 1993. The floor of the Gunnedah Basin north of the Liverpool Range. In: TADROS, N.Z. (ed.) The Gunnedah Basin, New South Wales. Geological Survey of New South Wales, Sydney, Memoir Geology, 12, 335-348. LEITCH, E.G. & SCHEIBNER, E. 1987. Stratotectonic terranes of the eastern Australian Tasmanides. In: LEITCH, E.G. & SCHEIBNER, E. (eds) Terrane Accretion and Orogenic Belts. American Geophysical Union, Washington, Geodynamic Series, 19,1-19. LEITCH, E.C., FERGUSSON, C.L. & HENDERSON, R.A. 1992. Geological Note: The intra-Devonian angular unconformity at Mt Gelobera, south of Rockhampton, central Queensland. Australian Journal of Earth Sciences, 39, 121-122. LEITCH, E.C., FERGUSSON, C.L. & HENDERSON, R.A. 2003. Arc to craton provenance switching in a Late Palaeozoic subduction complex, Wandilla and Shoalwater terranes, New England Fold Belt, eastern Australia. Australian Journal of Earth Sciences, 50, 919-929. LENNOX, P.G. & FLOOD, P.G. 1997. Age and structural characterisation of the Texas megafold, southern New England Orogen, eastern Australia. In: ASHLEY, P.M. & FLOOD, P.G. (eds) Tectonics and Metallogenesis of the New England Orogen. Geological Society of Australia, Special Publication, 19, 161-177. LENNOX, P.G, FOWLER, T. & FOSTER, D. 1998. Barry Granodiorite and Sunset Hills Granite: Wyangala-style intrusion at the margin of a regional ductile shear zone, northern Lachlan Fold Belt, New South Wales. Australian Journal of Earth Sciences, 45, 849-863. Li, Z.X., ZHANG L. & POWELL, C.M.A. 1995. South China in Rodinia: part of the missing link between Australia-East Antarctica and Laurentia? Geology, 23, 407^10. LISTER, G.S., ETHERIDGE, M.A. & SYMONDS, PA. 1991. Detachment models for the formation of passive continental margins. Tectonics, 10,1038-1064. LITTLE, T.A., HOLCOMBE, R.J., GIBSON, G.M., OFFLER, R.L., GANS, P. & MCWILLIAMS, M.O. 1992. Exhumation of late Paleozoic blue schists in Queensland, Australia by extensional faulting. Geology, 22, 231-234. LITTLE, T.A., MCWILLIAMS, M.O. & HOLCOMBE, R.J.

THE TASMANIDES OF EASTERN AUSTRALIA 1995. 40Ar/39Ar Thermochronology of epidote blueschists from the North D'Aguilar block, Queensland Australia: timing and kinematics of subduction complex unroofing. Geological Society of America Bulletin, 107, 520-535. LYONS, P. 2000. Structural Geology. Forbes 1:250 000 Geological Sheet SI 55-7, Explanatory Notes (second edition), Record 2000/20, Australian Geological Survey Organisation, Canberra. LYONS, P. & PERCIVAL, I.G. 2002. Middle to Late Ordovician age for the Jindalee Group of the Lachlan Fold Belt, New South Wales: conodont evidence and some tectonic implications. Australian Journal of Earth Sciences, 49, 801-808. MALONE, E.J. 1964. Depositional evolution of the Bowen Basin. Journal of the Geological Society of Australia, 11,263-282. MARSDEN, M.A.H. 1972. The Devonian history of northeastern Australia. Journal of the Geological Society of Australia, 19,125-162. MARSHAK, S. & FLOTTMANN, T. 1996. Structure and origin of the Fleurieu and Nackara Arcs in the Adelaide fold-thrust belt, South Australia. Journal of Structural Geology, 18, 891-908. MCELHINNY, M.W., POWELL, C.M. & PISAREVSKY, S.A. 2003. Paleozoic terranes of eastern Australia and the drift history of Gondwana. Tectonophysics, 362, 41-65. MclLVEEN, G.R. 1974. The Eden-Comerong-Yalwal Rift zone and the contained gold mineralisation. Geological Survey of New South Wales Bulletin, 16, 245-277. McPniE, J. 1987. Andean analogue for Late Carboniferous volcanic arc and arc flank environments of the western New England Orogen, New South Wales. Tectonophysics, 138, 269-288. MEFFRE, S., BERRY, R.F. & HALL, M. 2000. Cambrian metamorphic complexes in Tasmania: tectonic implications. Australian Journal of Earth Sciences, 47, 971-985. MEFFRE, S., SCOTT, R.J., GLEN, R.A. & SQUIRE, R. 2005. Re-evaluation of contacts between Ordovician volcanic belts and the quartz-rich turbidites of the Lachlan Orogen. Unpublished. MESSENGER, P.R. 1996. Relationships between Devonian magmatism and Au-Cu mineralisation at Mt Morgan, central Queensland. PhD thesis, University of Queensland, Brisbane. MILLER, J.M. & GRAY, D.R. 1996. Structural signature of sediment subduction-accretion in a Palaeozoic accretionary complex, southeastern Australia. Journal of Structural Geology, 18,1245-1258. MILLER, J.M., DUGDALOE, L.J. & WILSON, C.J.L. 2001. Variable hangingwall palaeotransport during Silurian and Devonian thrusting in the western Lachlan Fold Belt: missing gold lodes, synchronous Melbourne Trough sedimentation and Grampians Group fold interference. Australian Journal of Earth Sciences, 48, 901-909. MILLER, J.M., PHILLIPS, D., WILSON, C. & DUGDALE, J. 2003. 40Ar/39Ar dating in western Victoria: implications for the evolution of the Lachlan and Delamerian orogens. Abstracts of the Geological Society of Australia, 72, 73.

89

MILLER, J.M., PHILLIPS, D., WILSON, C.J.L. & DUGDALE, L.J. 2004. A new tectonic model for the Delamerian and western Lachlan orogens. Abstracts of the Geological Society of Australia, 73, 174. MILLIGAN, PR., PETKOVIC, P. & DRUMMOND, B.J. 2003. Potential-field datasets for the Australian region: their significance in mapping basement architecture. In: HILLIS, R.R. & MULLER, R.D. (eds) Evolution and Dynamics of the Australian Plate. Special Publication Geological Society of Australia, 22 and Special Paper Geological Society of America, 372,129-139. MILLS, K.J. 1992. Geological evolution of the Wonaminta Block. Tectonophysics, 214, 57-68. MILLS, K. 2001. New Broken Hill Exploration seeks copper-gold. Minfo, New South Wales Mining Industry Quarterly, 69, 27. MILLS, K.J. 2002. Relationships between the Delamerian and Lachlan fold belts in western New South Wales. Abstracts of the Geological Society of Australia, 67,136. MILLS, K.J. 2003. Composition and structure of the Koonenberry region. In: PELJO, M. (ed.) Broken Hill Exploration Initiative. Geoscience Australia, Canberra, Record 2003/13, 116-119. MILLS, KJ. & DAVID, V. 2003. The Koonenberry deep seismic reflection line and geological modelling of the Koonenberry region. Report of the Geological Survey of New South Wales, 2003/289. MORAND, V.J. 1988. Emplacement and deformation of the Wyangala Batholith, New South Wales. Australian Journal of Earth Sciences, 35,339-353. MORAND, V.J. 1993. Stratigraphy and tectonic setting of the Calliope Volcanic Assemblage, Rockhampton area, Queensland. Australian Journal of Earth Sciences, 40, 15-30. MORAND, V.J. & GRAY, D.R. 1991. Major fault zones related to the Omeo Metamorphic Complex northeastern Victoria. Australian Journal of Earth Sciences, 38, 203-221. MUNKER, C. & CRAWFORD, A.J. 2000. Cambrian arc evolution along the SE Gondwana active margin: a synthesis from Tasmania-New ZealandAustralia-Antarctica correlations. Tectonics, 19, 415-432. MUIR, B.J. 1999. Structure of the Girilambone Group at Tottenham, central western N.S.W.; implications for copper-pyrite mineralisation. BSc thesis, School of Earth Sciences. Hobart, University of Tasmania. MURRAY, C.G. 1986. Metallogeny and tectonic development of the Tasman Fold Belt System in Queensland. Ore Geology Reviews, 1, 315-400. MURRAY, C.G. 1990. Tectonic evolution and metallogeny of the Bowen Basin. Abstracts of the Bowen Basin Symposium, Geological Society of Australia, Queensland Division, 201-212. MURRAY, C.G. 1994. Basement cores from the Tasman Fold Belt System beneath the Great Artesian Basin in Queensland. Geological record, Queensland Department of Minerals And Energy, 1994/10. MURRAY, C.G. 1997. From geosyncline to fold belt: a personal perspective on the development of ideas

90

R. A. GLEN

regarding the tectonic evolution of the New England Orogen. In: ASHLEY, P.M. & FLOOD, P.G. (eds) Tectonics and Metallogenesis of the New England Orogen. Geological Society of Australia, Special Publication, 19,1-28. MURRAY, C.G. 2003. Granites of the northern New England Orogen. In: BLEVIN, P.L., JONES, M. & CHAPPELL, B. (eds) Magmas to Mineralisation: The Ishihara Symposium. Geoscience Australia Record, 2003/14, 101-108. MURRAY, C.G. & KIRKEGAARD A.G. 1978. The Thomson Orogen of the Tasman Orogenic Zone. Tectonophysics, 48, 299-325. MURRAY, C.G, FERGUSSON, C.L., FLOOD, P.G., WHITTAKER, W.G. & KORSCH, RJ. 1987. Plate tectonic model for the Carboniferous evolution of the New England fold Belt. Australian Journal of Earth Sciences, 34, 213-236. MURRAY, C.G, SCHEIBNER, E. & WALKER, R.N. 1989. Regional geological interpretation of a digital coloured residual Bouguer gravity image of eastern Australia with a wavelength cut-off of 250 km. Australian Journal of Earth Sciences, 36, 423-449. MURRAY, C.G, BLAKE, PR., HUTTON, L.J., WITHNALL, I.W., HAYWARD, M.A., SIMPSON, G.A. & FORDHAM, B.G. 2003. Discussion: Yarrol Terrane of the northern New England Fold Belt: forearc or back arc? Australian Journal of Earth Sciences, 50, 271-278. NEEF, G. 2004. Stratigraphy, sedimentology, structure and tectonics of Lower Ordovician and Devonian strata of South Mootwingee, Darling Basin, western New South Wales. Australian Journal of Earth Sciences, 51, 15-29. NEEF, G., EDWARDS, A.C., BOTTRILL, R.S., HATTY, J., HOLZBERGER, I., KELLY, R. & VAUGHAN, J. 1989.

The Mt. Daubeny Formation: Arenite-rich ?Late Silurian-Early Devonian (Gedinnian) strata in far western New South Wales. Journal and Proceedings of the Royal Society of New South Wales, 122, 97-106. NOLL, C.A. & HALL, M. 2003. Stratigraphic architecture and depositional setting of the coarsegrained Upper Cambrian Owen Conglomerate, West Coast Range, western Tasmania. Australian Journal of Earth Sciences, 50, 835-852. NOLL, C.A. & HALL, M. 2004. The growth history of an ancient extensional border fault system, and the role of normal fault growth on sedimentation during basin formation: a case study from the Late Cambrian Owen Conglomerate, West Coast Range, western Tasmania, Australia. Abstracts of the Geological Society of Australia, 73, 106. OFFLER, R. 1999. Origin and significance of blueschist 'knockers', Glenrock Station, NSW. In: FLOOD, P.G. (ed.) Conference on New England Orogen, Regional Geology, Tectonics and Metallogenesis. University of New England, Armidale, 35-44. OFFLER, R. & GAMBLE, J. 2002. Evolution of an intraoceanic island arc during the Late Silurian to Late Devonian, New England Fold Belt. Australian Journal of Earth Sciences, 49, 349-366.

O'HALLORAN, G.J. & CAS, R.A.F. 1995. Evidence for syndepositional deformation from the Late Devonian of the Mansfield Basin, east-central Victoria. Australian Journal of Earth Sciences, 42, 581-596. O'REILLY, S. 1990. Discussion The Sydney Basin: Composition of basement. Australian Journal of Earth Sciences, 37, 485-486. OVERSBY, B. 1971. Palaeozoic plate tectonics in the southern Tasman Geosyncline. Nature, Physical Sciences, 234, 45-48. OWEN, M. & WYBORN, D. 1979. Geology and geochemistry of the Tantangara and Brindabella 1:100 000 sheet areas, New South Wales and Australian Capital Territory. Bulletin, Bureau of Mineral Resources, 204. PACKHAM, G.H. 1969. Tectonics and Sedimentation. In: PACKHAM, G.H. (ed.) The Geology of New South Wales. Journal of the Geological Society of Australia, 16, 216-226. PACKHAM, G.H. 1973. A speculative Phanerozoic history of the south-west Pacific. In: COLEMAN, PJ. (ed.) The Western Pacific, Island arcs, marginal seas, geochemistry. University of Western Australia Press, Perth, 369-388. PACKHAM, G.H. & FALVEY, D.A. 1971. An hypothesis for the formation of marginal seas in the western Pacific. Tectonophysics, 11, 79-109. PATERSON, S.R., TOBISCH, O.T. & MORAND, VJ. 1990. The influence of large ductile shear zones on the emplacement and deformation of the Wyangala Batholith, southeast Australia. Journal of Structural Geology, 12, 639-650. PAULICK, H. & McPniE, J. 1999. Facies architecture of the felsic lava-dominated first sequence to the Towanga massive sulfide deposit, Lower Ordovician, Northern Queensland. Australian Journal of Earth Sciences, 46, 391-405. PERCIVAL, I.G & GLEN, R. A. 2006. Stratigraphy of the Ordovician Macquarie Arc. Australian Journal of Earth Sciences, 53, in press. PERCIVAL, I.G, ZHEN, Y.-Y. & WEBBY, B.D. 2003. Early Ordovician conodont distribution from craton to basin and island terranes in East Gondwana. In: ALBANESE, G.L., BERESI, M.S. & PERALTA, S.H. (eds) Ordovician from the Andes. Institute Superior de Correlacion Geologicca Insugeo, Series Correlacion Geologica, 533-537. PERKINS, C, MCDOUGALL, I. & WALSHE, J.L. 1993. Isotopic dating of precious and base metal deposits and their rocks in the eastern Australia. Final report. Project P. 334, AMIRA, Melbourne. PERKINS, C. & WALSHE, J.L. 1993. Geochronology of the Mount Read Volcanics, Tasmania, Australia. Economic Geology, 88, 1176-1197. PERKINS, C, WALSHE, J.L. & MORRISON, G. 1995. Metallogenic episodes of the Tasman Fold Belt system, eastern Australia. Economic Geology, 90, 1443-1466. PHILIP, G.M. 1966. The occurrence and palaeogeographic significance of Ordovician strata in northern New South Wales. Australian Journal of Science, 29, 112-114. PHILLIPS, G, MILLER, J.M. & WILSON, C.J.L. 2002.

THE TASMANIDES OF EASTERN AUSTRALIA Structural and metamorphic evolution of the Moornambool Metamorphic Complex, western Lachlan Fold Belt, southeastern Australia. Australian Journal of Earth Sciences, 49, 891-913. PICKETT, J.W. 1982. The Silurian System in New South Wales. Bulletin of the Geological Survey of New South Wales, 29. POGSON, DJ. 1982. Stratigraphy, structure and tectonics: Nymagee-Melrose, central western New South Wales. M. Applied Sc. thesis, New South Wales Institute of Technology, Sydney. POGSON, DJ. & PERCIVAL, I.G. 2003. Hill End Trough time: integrating biostratigraphic zonation and isotopic age dating. In: VASSALLO, JJ. & GLEN, R.A. (eds) Evolution of the Hill End Trough: Hill End Trough Geoscience Database. Geological Survey of New South Wales, Sydney, 21-25. POGSON, DJ. & WATKINS, JJ.C. 1998. Bathurst 1:250 000 Geological Sheet SI/55-8: Explanatory Notes. Geological Survey of New South Wales. Sydney. POWELL, C.M. 1983. Tectonic relationships between the Late Ordovician and Late Silurian palaeogeographies of southeastern Australia. Journal of the Geological Society of Australia, 30, 353-373. POWELL, C.M. 19840. Ordovician to earliest Silurian: Marginal sea and island arc; Silurian to midDevonian dextral transtensional margin. In: VEEVERS, JJ. (ed.) Phanerozoic Earth history of Australia. Oxford Science Publications, Oxford, 290-309. POWELL, C.M. 19846. (h) Late Devonian and early Carboniferous: continental magmatic arc along the eastern edge of the Lachlan Fold belt. In: VEEVERS, JJ. (ed.) Phanerozoic Earth History of Australia. Oxford Science Publications, Oxford, 329-240. POWELL, C.M., CONAGHAN, PJ. & PRENDERGAST, E.I. 1984. Tectonostratigraphy of the Devonian? Carboniferous Lambian Foreland Basin. Abstracts of the Geological Society of Australia, 12, 438-439. POWELL, C.M., BAILLIE, PA., CONAGHAN, PJ. & TURNER, NJ. 19930. The mid-Palaeozoic turbiditic Mathinna Group, northeast Tasmania. Australian Journal of Earth Sciences, 40, 169-196. POWELL, C.M., MCELHINNY, M.W, MEERT, J.G. & PARK, J.K. 1993/?. Paleomagnetic constraints on timing of the Neoproterozoic breakup of Rodinia and the Cambrian formation of Gondwana. Geology, 21, 889-892. POWELL, C.M., PREISS, W.V., GATEHOUSE, C.G., KRAPEZ, B. & Li, Z.X. 1994. South Australian record of a Rodinia epicontinental basin and its mid-Neoproterozoic breakup (c.700 Ma) to form the Paleo Pacific Ocean. Tectonophysics, 237, 113-140. POWELL, C.M., BAILLIE, P.W., & VANDENBERG, A.H.M. 2003. Silurian to mid-Devonian basin development of the Melbourne Zone, Lachlan Fold belt, southeastern Australia. Tectonophysics, 375, 9-36. POWELL, W. & O'REILLY, S. 2001. Ancient mantle

91

beneath New England. GEMOC (National Key Centre for Geochemical Evolution and Metallogeny of Continents, Macquarie University, Sydney) Annual Report, 32-33. PREISS, W.V. 2000. The Adelaide Geosyncline of South Australia and its significance in Neoproterozoic continental reconstruction. Precambrian Research, 100, 21-63. QURESHI, I.R. 1984. Wollondilly-Blue Mountains gradient and its bearing on the origin of the Sydney Basin. Australian Journal of Earth Sciences, 31, 293-302. RAMOS, V.A. 1988. Late Proterozoic-Early Paleozoic of South America - a collisional history. Episodes, 11,168-174. RAMOS, V.A. 1999. Ciclos Orogenicos y Evolution Tectonica. In: ZAPPATTINI, E.O. (ed.) Anales Recursos Minerales de la Republica Argentina. Institute de Geologia y Recursos Minerales SEGEMAR, Buenos Aires, 35, 29^9. RANKIN, L.R., CLOUGH, BJ. & GATEHOUSE, C.G. 1991. Early Palaeozoic mafic suites of the western Tasman fold belt System. South Australian Department of Mines and Energy Report. DME 281789. RAPELA, C.W, PANKHURST, R.J., CASQUET, C, BALDO, E., SAAVEDRA, J. & GALINDO, C. 1998. Early evolution of the proto-Andean margin of South America. Geology, 26, 707-710. RAYMOND, O.L. 1996. The Dulladerry VolcanicsMiddle Devonian A-type volcanics from centralwest New South Wales. Abstracts of the Geological Society of Australia, 41, 359. RAYMOND, O., SHERWIN, L., LYONS, P. & SCOTT, M. 2001. The Jemalong Trough: further evidence of Silurian-Devonian rifting in the eastern Lachlan Orogen. In: KORSCH, RJ. & LYONS, P. (eds) Integrated geophysical and geological studies of the northeastern Lachlan Orogen. Australian Geological Survey Organisation Record, 20017 009, 69-76. REED, A.R. 2001. Pre-Tabberabberan deformation in eastern Tasmania: a southern extension of the Benambran Orogeny. Australian Journal of Earth Sciences, 48, 785-796. REED, A.R., CALVER, C. & BOTTRILL, R.S. 2002. Palaeozoic suturing of eastern and western Tasmania in the west Tamar region: implications for the tectonic evolution of southeastern Australia. Australian Journal of Earth Sciences, 49, 809-830. RICHARDS, J.R. & SINGLETON, O.P. 1981. Palaeozoic Victoria, Australia: igneous rocks, ages and their interpretation. Journal of the Geological Society of Australia, 28, 395-421. ROBERTS, J. & ENGEL, B.A. 1987. Depositional and tectonic history of the southern New England Orogen. Australian Journal of Earth Sciences, 34, 1-20. ROBERTS, J. & GEEVE, RJ. 1999. Allochthonous forearc blocks and their influence on an orogenic timetable for the southern New England Orogen. In: FLOOD, P.G. (ed.) Conference on Regional Geology, Tectonics and Metallo genesis, New

92

R. A. GLEN

England Orogen. University of New England, Armidale, 105-114. ROBERTS, J., CLAOUE-LONG, J.C. & JONES, PJ. 1995. SHRIMP zircon age control of Gondwanan sequences in Late Carboniferous and Early Permian Australia. In: DUNNAY, R.E. & HAILWOOD, E.A. (eds) Dating and Correlating Biostratigraphically Barren Strata. Geological Society, London, Special Publications, 89, 145-174. ROBERTS, X, CLAOUE-LONG, J.C. & FOSTER, C.B. 1996. SHRIMP zircon dating of the Permian System of eastern Australia. Australian Journal of Earth Sciences, 43, 401-421. ROBERTS, J., WANG, X. & FANNING, M. 2003. Stratigraphy and correlation of Carboniferous ignimbrites, Rocky Creek region, Tamworth Belt, southern New England Orogen, New South Wales. Australian Journal of Earth Sciences, 50, 931-954. ROBERTS, I, OFFLER, R. & FANNING, M. 2004. Upper Carboniferous to Lower Permian volcanic successions of the Carroll-Nandewar region, northern Tamworth Belt, southern New England Orogen, Australia. Australian Journal of Earth Sciences, 51, 205-232. ROYDEN, L.H. 1993. The tectonic expression of slab pull at continental convergent boundaries. Tectonics, 12, 303-325. ROYDEN, L.H. & BURCHFIEL, B.C. 1989. Are systematic variations in thrust belt style related to plate boundary processes? (The Western Alps versus the Carpathians). Tectonics, 8, 51-61. RUSSELL, T.G. 1979. A reappraisal of the Late Devonian Bective Unconformity. Journal and Proceedings of the Royal Society of New South Wales, 112, 63-69. RUTLAND, R.W.R. 1976. Orogenic evolution of Australia. Earth Science Reviews, 12, 161-196. SAVAGE, A.E.P.L. 2000. EL 5039 'Louth' third annual report on exploration for the 12 months ending 19th June 1999 and final report. Report Geological Survey of New South Wales GS 2000/129, unpublished. SCHEIBNER, E. 1973. A plate tectonic model of the Palaeozoic tectonic history of New South Wales. Journal of the Geological Society of Australia, 20, 405-426. SCHEIBNER, E. 1974. Fossil fracture zones (transform faults), segmentation and correlation problems in the Tasman Fold Belt System. In: DENMEAD, A.K., TWEEDALE, G.W. & WILSON, A.F. (eds) The Tasman Geosyncline - a symposium in honour of Professor Dorothy Hill. Geological Society of Australia, Queensland Division, Brisbane, 65-98. SCHEIBNER, E. 1976. Explanatory Notes on the Tectonic Map of New South Wales. Geological Survey of New South Wales, Sydney. SCHEIBNER, E. 1985. Suspect terranes in the Tasman Fold Belt System, eastern Australia. In: HOWELL, D.J. (ed.) Tectonostratigraphic Terranes of the Circum-Pacific region. Circum-Pacific Council for Energy and Mineral Resources, Earth Science Series, Houston, 1, 493-514. SCHEIBNER, E. 1987. Paleozoic development of

eastern Australia in relation to the Pacific region. In: MONGER, J.W.H. & FRANCHETAU, J.J. (eds) Circum-Pacific orogenic belts and evolution of the Pacific ocean basin. American Geophysical Union Geodynamic Series, 18,133-165. SCHEIBNER, E. 1989. The tectonics of New South Wales in the second decade of application of the plate tectonic paradigm. Journal and Proceedings of the Royal Society of New South Wales, 122, 35-74. SCHEIBNER, E. & BASDEN, H. (eds) 1996. Geology of New South Wales - Synthesis. Volume 1: Structural framework. Memoir of the Geological Survey of New South Wales, Sydney, 13. SCHEIBNER, E. & BASDEN, H. (eds) 1998. Geology of New South Wales - synthesis. Volume 2: Geological evolution. Memoir of the Geological Survey of New South Wales, 13. SCHEIBNER, E. & VEEVERS, J.J. 2000. Tasman Fold Belt System. Billion-year earth history of Australia and neighbours in Gondwanaland. In: VEEVERS. J.J. (ed.) Billion-year earth history of Australia and neighbours in Gondwanaland. GEMOC Press, Sydney, 154-234. SCHMIDT, P.W., AUBOURG, G, LENNOX, P.G. & ROBERTS, J. 1994. Palaeomagnetism and tectonic rotation of the Hastings Terrane, eastern Australia. Australian Journal of Earth Sciences, 41, 547-560. SDROLIAS, M., MULLER, R.D. & GAINA, C. 2003. Tectonic evolution of the southwest Pacific using constraints from backarc basins. In: HILLIS, R.J. & MULLER, R.D. (eds) Evolution and Dynamics of the Australian Plate. Geological Society of Australia Special Publication 22 and Geological Society of America Special Paper 372, 343-359. SEYMOUR, D.B. & CALVER, C.R. 1995. Explanatory notes for the time-space diagram and stratotectonic elements map of Tasmania. Record of the Tasmanian Geological Survey, 1. SHARP, T. & BUCKLEY, P. 2003. Palaeozoic history of the Bancannia Trough & resurrection of the Mt. Wright Volcanic Arc. In: PELJO, M. (ed.) Broken Hill Exploration Initiative. Geoscience Australia Record, 2003/13, 154-157. SHAW, R.D., FAWKNER, J.F. & BULTITUDE, R.J. 1987. The Palmerville fault system: a major imbricate thrust system in the Northern Tasmanides. Australian Journal of Earth Sciences, 34, 69-93. SHAW, R.D., WELLMAN, P., GUNN, P., WHITAKER, A.I, TARLOWSKI, C. & MORSE, M. 1996. Australian Crustal Elements Map. Australian Geological Survey Organisation, Canberra, Record 1996/30. SHAW, S.E. & FLOOD, R.H. 1993. Carboniferous magmatic activity in the Lachlan and New England Fold Belts. In: FLOOD, P.G. & AITCHISON, J.C. (eds) Conference on New England Orogen, eastern Australia. University of New England, Armidale, 113-121. SHAW, S.E., FLOOD, R.H. & LANGWORTHY, PJ. 1989. Age and association of the Rylstone Volcanics: new isotopic evidence. 23rd Newcastle Symposium on Advances in the Study of the Sydney Basin. University of Newcastle, Newcastle, 45-51.

THE TASMANIDES OF EASTERN AUSTRALIA SHAW, S.E., CONAGHAN, PJ. & FLOOD, R.H. 1991. Late Permian and Triassic igneous activity in the New England Batholith and contemporaneous tephra in the Sydney and Gunnedah Basins. 25th Newcastle Symposium on Advances in the Study of the Sydney Basin. University of Newcastle, Newcastle, 44-51. SIVELL, W.J. & ARNOLD, G.O. 1999. Geochemical and volcanic stratigraphy of a segment of an island arc-Gondwana rim accretion zone, Gympie Province, southeast Queensland. In: FLOOD, P.G. (ed.) Conference on Regional Geology, Tectonics and Metallo genesis, New England Orogen. University of New England, Armidale, 255-266. SIVELL, WJ. & MCCULLOCH, M.T. 2001. Geochemical and Nd-isotopic systematics of the PermoTriassic Gympie Group, southeast Queensland. Australian Journal of Earth Sciences, 48, 377-393. SLIWA, R., HOLCOMBE, R.J., FIELDING, C.R., LITTLE, T.A., BRYAN, S.E. & FIFOOT, A. 1993. Early Permian marine fault basins formed during exhumation of the New England Orogen subduction complex in southeastern Queensland. In: FLOOD, P.G. & AITCHISON, J.C. (eds) Conference on New England Orogen, eastern Australia. University of New England, Armidale, 557-564. SMITH, J.V. 1999. Structure of the southern Beenleigh block, northeast New South Wales. In: FLOOD, P.G. (ed.) Conference on Regional Geology, Tectonics and Metallogenesis, New England Orogen. University of New England, Armidale, 181-187. SMITH, R.E. 1969. Zones of progressive regional burial metamorphism in part of the Tasman Geosyncline, eastern Australia. Journal of Petrology, 10,143-163. SOESOO, A. & NICHOLLS, 1.1999. Mafic rocks spatially associated with Devonian felsic intrusions of the southern Lachlan Fold Belt: a possible mantle contribution to crustal evolution processes. Australian Journal of Earth Sciences, 46, 725-734. SOESOO, A., BONS, P.D., GRAY, D.A. & FOSTER, D.A. 1997. Divergent double subduction: tectonic and petrologic consequences. Geology, 25, 755-758. SOLOMON, M. & GRIFFITHS, J.R. 1972. Tectonic evolution of the Tasman Orogenic Zone, eastern Australia. Nature, Physical Sciences, 237, 3-6. SPAGGIARI, C.V., GRAY, D.R., FOSTER, D.A. & FANNING, M. 2002. Occurrence and significance of blueschists in the southern Lachlan Orogen. Australian Journal of Earth Sciences, 49, 255-269. SPAGGIARI, C.V., GRAY, D.R. & FOSTER, D.A. 20030. Tethyan- and Cordilleran-type ophiolites of eastern Australia: implications for the evolution of the Tasmanides. In: DILEK, Y. & ROBINSON, P.T. (eds) Ophiolites in Earth History. Geological Society, London, Special Publications, 218, 517-539. SPAGGIARI, C.V., GRAY, D.R., FOSTER, D.A. & MCKNIGHT, S. 20036. Evolution of the boundary between the western and central Lachlan

93

Orogen; implications for Tasmanide tectonics. Australian Journal of Earth Sciences, 50,725-749. SQUIRE, R.J., CAMPBELL, I.H., ALLEN, C. & WILSON, CJ.L. 2003c. U-Pb detrital-zircon LA-ICP-MS age dating of early Palaeozoic sandstones from the Lachlan Orogen: sediment provenance and implications for tectonic evolution. Abstracts Geological Society of Australia, 72, 72. STEVENS, B. 1985. Preliminary interpretation of regional basement geology in northwestern New South Wales. Quarterly Notes, Geological Survey of New South Wales, 61, 9-22. STEVENS, B.PJ. 1991. Northwestern New South Wales and its relationship to the Lachlan Fold Belt. Abstracts of the Geological Society of Australia, 29, 50. STEWART, I.R. 1988. Conodonts. In: DOUGLAS, J.G. & FERGUSON, J.A. (eds) Geology of Victoria. Geological Society of Australia, Victorian Division, 79-81. STEWART, I. 1995. Cambrian age for the Pipeclay Creek Formation, Tamworth Belt, northern New South Wales. Courier Forschung Institut Senckenbergiana, 182, 565-566. STEWART, I.R. & GLEN, R.A. 1991. New Cambrian and Early Ordovician ages from the New South Wales south coast. Quarterly Notes of the Geological Survey of New South Wales, 85, 1-8. STOLZ, A.J. 1995. Geochemistry of the Mount Windsor volcanics: implications for the tectonic setting of Canberra-Ordovician volcanic-hosted massive sulfide mineralisation in northeastern Australia. Economic Geology, 90,1080-1097. STRATFORD, J.M.C. 1993. Lithostratigraphy of the Gamilaroi terrane at Glenrock Station, Scone, NSW. In: FLOOD, P.G. & AITCHISON, J.C. (eds) Conference on New England Orogen, eastern Australia. University of New England, Armidale, 247-253. STRATFORD, J.M.C. & AITCHISON, J.C. 1996. Devonian intra-oceanic island arc rift sedimentation: facies development in the Gamilaroi terrane, New England Orogen, eastern Australia. Sedimentary Geology, 101, 173-193. STUART-SMITH, P.G. 1990. Evidence for extensional tectonics in the Tumut Trough, Lachlan Fold Belt, NSW. Australian Journal of Earth Sciences, 37, 147-167. SUPPEL, D.W & BARRON, L.M. 1986. Platinum in basic to ultrabasic intrusive complexes at Fifield: a preliminary report. Quarterly Notes of the Geological Survey of New South Wales, 65, 1-8. SUTHERLAND, R. 1999. Basement geology and tectonic development of the greater New Zealand region: an interpretation from regional magnetic data. Tectonophysics, 308, 341-362. TADROS, N.Z. 1993. Structural elements. The Gunnedah Basin. In: TADROS, N.Z. (ed.) The Gunnedah Basin, New South Wales. Geological Survey of New South Wales, Sydney, Memoir Geology, 12, 55-86. TAYLOR, D.H. & CAYLEY, R.A. 2000. Character and kinematics of faults within the turbiditedominated Lachlan Orogen: implications for

94

R. A. GLEN

tectonic evolution of eastern Australia. Discussion. Journal of Structural Geology, 22, 523-528. TEALE, G.S., FANNING, CM., FLOOD, R.H. & PURVIS, A.C. 1999. The geology, geochemistry and mineralisation of the Mt. Terrible Volcanic Complex. In: FLOOD, P.G. (ed.) Conference on Regional Geology, Tectonics and Metallogenesis, New England Orogen. University of New England, Armidale, 403-407. TEYSSIER, C. & TIKOFF, B. 1995. Oblique plate motion and continental tectonics. Geology, 23, 447-450. THOMAS, O.D., JOHNSTON, A.J., SCOTT, M.M., POGSON, D.J., SHERWIN, L. & MACRAE, G.P. 2002. Goulburn 1:100 000 Geological Sheet 8828, First Edition. Geological Survey of New South Wales, Orange and Sydney. THOMSON, B.P. 1969. Palaeozoic Era. In: PARKIN, L.W. (ed.) Handbook of South Australian Geology. Geological Survey of South Australia, Adelaide, 84-132. TURNER, N.J., BLACK, L.P. & KAMPERMAN, M. 1998. Dating of Neoproterozoic and Cambrian orogenies in Tasmania. Australian Journal of Earth Sciences, 45, 789-806. TURNER, N.J., BOTTRILL, R.S., CRAWFORD, AJ. & VILLA, I. 1992. Geology and prospectivity of the Arthur Mobile Belt. Bulletin of the Geological Survey of Tasmania, 70, 227-233. TURNER, S.P., ADAMS, C.J., FLOTTMANN, T. & FODEN, J.D. 1993. Geochemical and geochronological constraints on the Glenelg River Complex, western Victoria. Australian Journal of Earth Sciences, 40, 275-292. TURNER, S.P., SANDIFORD, M. & FODEN, J. 1992. Some geodynamic and compositional constraints on 'postorogenic' magmatism. Geology, 20, 931-934. TURNER, S.P., KELLEY, S.P., VANDENBERG, A.H.M., FODEN, ID., SANDIFORD, M. & FLOTTMANN, T. 1996. Source of the Lachlan fold belt flysch linked to convective removal of the lithospheric mantle and rapid exhumation of the Delamerian-Ross fold belt. Geology, 24, 941-944. VANDENBERG, A.H.M. 1981. A complete Ordovician graptolitic sequence at Mountain Creek near Deddick, eastern Victoria. Geological Survey of Victoria, Melbourne. Report 1981/81. VANDENBERG, A.H.M. 1999. Timing of erogenic events in the Lachlan Orogen. Australian Journal of Earth Sciences, 46, 691-701. VANDENBERG, A.H.M. & STEWART, I.R. 1992. Ordovician terranes of southeastern Lachlan Fold Belt: stratigraphy, structure and palaeogeographic reconstruction. Tectonophysics, 214, 177-192. VANDENBERG, A.H.M., NOTT, R.J. & GLEN, R.A. 1992. Bendoc 1:100000 map geological report. Geological Survey of Victoria, Melbourne, Report 90. VANDENBERG, A.H.M., WILLMAN, C.E., HENDRICKX, M., BUSH, M.D. & SANDS, B.C. 1995. The geology and prospectivity of the 1993 Mount Wellington Airborne Survey area. Victorian Initiative for minerals and petroleum. Department of Agriculture, Energy and Minerals, Report 2.

VANDENBERG, A.H.M., WILLMAN, C.E. & MAKER, S.E.A. 2000. The Tasman Fold Belt System in Victoria. Geology and Mineralisation of Proterozoic to Carboniferous rocks. Geological Survey of Victoria, Melbourne, Special Publication. VAN DER HlLST, R.D., KENNETT, B.L.N. & SfflBUTANI,

T. 1998. Upper mantle structure beneath Australia from portable array deployments. In: BRAUN, J., DOOLEY, J., GOLEBY, B. & VAN DER HILST, R.D. & KLOOTWIJK, C. (eds) Structure and Evolution of the Australian Continent. American Geophysical Union, Geodynamics Monograph 26, 39-58. VASSALLO, X, GLEN, R.A. & DAVID, D. 2003. Structural re-evaluation of the Hill End Trough, Eastern Lachlan Orogen, NSW. Abstracts of the Geological Society of Australia, 72, 74. VEEVERS, JJ. 20000. Antarctic Beardmore-Ross and Mirny provenances saturate Palaeozoic-Mesozoic East Gondwanaland with 0.6-0.5 Ga zircons. In: VEEVERS, JJ. (ed.) Billion-year earth history of Australia and neighbours in Gondwanaland. GEMOC Press, Sydney, 110-130. VEEVERS, JJ. 20006. Impact on Australia and Antarctica of the collisional merging of Gondwanaland and Laurussia in Pangea. In: VEEVERS, JJ. (ed.) Billion-year earth history of Australia and neighbours in Gondwanaland. GEMOC Press, Sydney, 283-291. VEEVERS, JJ. 2000c (ed.) Billion-year earth history of Australia and neighbours in Gondwanaland. GEMOC Press, Sydney. VEEVERS, J J. 2000d. Permian-Triassic Pangean basins and fold belts along the Panthalassan margin of eastern Australia. In: VEEVERS, JJ. (eds) Billionyear earth history of Australia and neighbours in Gondwanaland. GEMOC Press, Sydney, 235-252. VEEVERS, JJ. 2000e. Time Scale. In: VEEVERS, JJ. (ed.) Billion-year earth history of Australia and neighbours in Gondwanaland. GEMOC Press, Sydney, 2-8. VEEVERS, JJ. & POWELL, C.M.A. 1984. Uluru and Adelaidean regimes: (d) Epi-Adelaidean: regional shear, (e) Cambrian: plate divergence on the east and north-west, followed by plate convergence on the east. In: VEEVERS, JJ. (ed.) Phanerozoic Earth history of Australia. Oxford University Press, Oxford, 278-289. VEEVERS, JJ., CONAGHAN, PJ. & POWELL, C. McA. 1994. Eastern Australia. In: VEEVERS, JJ. & POWELL, C. McA. (eds) Permian-Triassic Basins and Foldbelts Along the Panthalassan Margin of Gondwanaland. Memoir of the Geological Society of America, 184,11-171. VEEVERS, XX, WALTER, M.R. & SCHEIBNER, E. 1997. Neoproterozoic tectonics of Australia-Antarctica and the 560 Ma birth of the Pacific Ocean reflect the 400 m.y. Pangaean Supercycle. Journal of Geology, 105, 225-242. VERNON, R.H. & FLOOD, R.H. 1979. Microstructural evidence of time-relationships between metamorphism and deformation in the metasedimentary sequence of the northern Hill End Trough,

THE TASMANIDES OF EASTERN AUSTRALIA New South Wales, Australia. Tectonophysics, 58, 127-137. VERNON, R.H., WILLIAMS, V.A. & D'ARCY, W.F. 1983. Grain-size reduction and foliaton development in a deformed granitoid batholith. Tectonophysics, 92,123-145. VICKERS, M. 1999. Early Permian dextral transpression: evolution of a strike-slip basin at Cooplacurripa, southern New England Orogen. In: FLOOD, P.G. (ed.) Regional Geology, Tectonics, Metallogenesis, New England Orogen. University of New England, Armidale, 151-158. WARREN, R.G. & ELLIS, DJ. 1996. Mantle underplating, granite tectonics and metamorphic P-T-t paths. Geology, 24, 663-666. WARTENBERG, W, KORSCH, RJ. & SCHAEFER, A. 1999. Geometry of the Tamworth Belt in the New England Orogen beneath the Surat Basin, southern Queensland In: FLOOD, P.G. (ed.) Conference on Regional Geology, Tectonics and Metallogenesis, New England Orogen. University of New England, Armidale, 211-219. WATANABE, T, IWASAKI, M., ISHIGA, H. ET AL. 1988. Late Carboniferous orogeny in the southern New England Fold Belt, NSW, Australia. In: KLEEMAN, J.D. (ed.) New England Orogen Tectonics and Metallo genesis. University of New England, Armidale, 93-98. WATANABE, T, FANNING, C.M. & LEITCH, E.G. 1998. Neoproterozoic Attunga eclogite in the New England Fold Belt. Abstracts of the Geological Society of Australia, 49, 158. WATKINS, J.J. 19980. Geochemistry of the Carboniferous intrusions. Bathurst 1:250000 Geological Sheet SI/55-8. 2nd Edition. Explanatory Notes. Geological Survey of New South Wales, Sydney, 262-264. WATKINS, JJ. 19986. Geochemistry of the Gregra Group. Bathurst 1:250 000 Geological Sheet SI/558. 2nd Edition. Explanatory Notes. Geological Survey of New South Wales, Sydney, 185-186. WELLMAN, P. 1995. Tasman Orogenic System: a model for its subdivision and growth history based on gravity and magnetic anomalies. Economic Geology, 90, 1430-1442. WILLIAMS, E. 1978. Tasman Fold Belt System in Tasmania. Tectonophysics, 48,159-205. WILLIAMS, I.S. 2001. Response of detrital zircon and monazite, and their U-Pb isotopic systems, to regional metamorphism and host-rock partial melting, Cooma Complex, southeastern Australia. Australian Journal of Earth Sciences 48, 557-580. WILLIAMS, I.S. & CHAPPELL, B.W. 1998. Crustal evolution in southeastern Australia: a zircon viewpoint. The Bruce Chappell Symposium: granites, island arcs, the Mantle and ore deposits. Record of the Australian Geological Survey Organisation, 1998/33, 44. WILLIAMS, I.S., GOODGE, J.W., MYROW, P., BURKE, K. & KRAUS, J. 2002. Large scale sediment dispersal associated with the late Neoproterozoic assembly of Gondwana. Abstracts of the Geological Society of Australia, 67, 238.

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WILLMAN, C.E., VANDENBERG, A.H.M. & MORAND, V.J. 2002. Evolution of the southeastern Lachlan Fold Belt in Victoria. Australian Journal of Earth Sciences, 49, 271-289. WINGATE, M.T.D. & GIDDINGS, J.W. 2000. Age and palaeomagnetism of the Mundine Well dyke swarm Western Australia: implications for an Australia-Laurentia connection at 755 Ma. Precambrian Research, 100, 335-357. WINGATE, M.T.D., CAMPBELL, I.H., COMPSTON, W. & GIBSON, G.M. 1998. Ion microprobe U-Pb ages for Neoproterozoic basaltic magmatism in southcentral Australia, and implications for the breakup of Rodinia. Precambrian Research, 87, 135-159. WITHNALL, I.W. 1995. Pre-Devonian rocks of the southern Anakie Inlier, In: WITHNALL, I.W., BLAKE, P.R., CROUCH, S.B.S. ETAL. (eds) Geology of the southern part of the Anakie Inlier, central Queensland. Queensland Geology, 7, 20-66. WITHNALL, I.W, GOLDING, S.D., REES, I.D. & DOBOS, S.K. 1996. K-Ar dating of the Anakie Metamorphic Group: evidence for the extension of the Delamerian Orogeny into central Queensland. Australian Journal of Earth Sciences, 43, 567-572. WOOD, D. & LISTER, G.S. 2004. Age and implications of core complex formation in central east Queensland. Abstracts of the Geological Society of Australia, 73, 194. WOODWARD, N.B. 1995. Thrust systems in the Tamworth Zone, southern New England Orogen. Australian Journal of Earth Sciences, 42,107-117. WOODWARD, N.B., GRAY, D.R. & ELLIOTT, C.G. 1993. Repeated Palaeozoic thrusting and allochthoneity of Precambrian basement northern Tasmania. Australian Journal of Earth Sciences, 40,297-311. WYBORN, D. 1992. The tectonic significance of Ordovician magmatism in the eastern Lachlan Fold Belt. Tectonophysics, 214, 287-303. WYBORN, D. & CHAPPELL, B.W. 1986. The petrogenetic significance of chemically altered plutonic and volcanic units. Geological Magazine, 123, 619-628. WYBORN, D. & OWEN, M. 1986. Araluen New South Wales. 1:100 000 Geological map commentary. Bureau of Mineral Resources, Canberra. WYBORN, D., TURNER, B.S. & CHAPPELL, B.W. 1987. The Boggy Plain Supersuite: a distinctive belt of I-type igneous rocks of potential significance in the Lachlan Fold Belt. Australian Journal of Earth Sciences, 34, 21-43. YARROL PROJECT TEAM. 1997. New insights into the geology of the northern New England Orogen in the Rockhampton-Monto region, central costal Queensland: progress report on the Yarrol project. Queensland Government Mining Journal, 98,11-26. ZEN, E.-A. 1995. Crustal magma generation and lowpressure high-temperature regional metamorphism in an extensional environment; possible application to the Lachlan Belt, Australia. American Journal of Science, 295, 851-874. ZHOU, B. & WHITFORD, D.J. 1994. Geochemistry of the Mt. Wright Volcanics from the Wonominta

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Block, northwestern New South Wales. Australian Journal of Earth Sciences, 41,331-340. ZUCCHETTO, R.G., HENDERSON, R.A., DAVIS, B.K. & WYSOCZANSKI, R. 1999. Age constraints on defor-

mation of the eastern Hodgkinson Province, north Queensland: new perspectives on the evolution of the northern Tasman Orogenic Zone. Australian Journal of Earth Sciences, 46,105-114.

The Appalachian peri-Gondwanan realm: a palaeogeographical perspective from the south JAMES P. HIBBARD1, BRENT V. MILLER2, ROBERT J. TRACY3 & BRAD T. CARTER1 ^Department of Marine, Earth and Atmospheric Sciences, North Carolina State University, Raleigh, NC27695, USA (e-mail: [email protected]) ^Department of Geology & Geophysics, Texas A&M University, College Station, TX 77843-3115, USA ^Department of Geosciences, Virginia Polytechnic Institute and State University, Blacksburg, VA 24061, USA Abstract: The Appalachian peri-Gondwanan realm (APGR) is an extensive tract of exotic Neoproterozoic-early Palaeozoic crustal blocks that occupies the eastern flank of the orogen. Traditionally, southern APGR elements have been correlated with those of the northern Appalachians on the basis of gross geological similarities. Most palaeogeographical reconstructions of the APGR are based on data from the northern Appalachians; consequently in these reconstructions, southern APGR elements are viewed commonly either as being affiliated spatially with those of the north or ignored. However, emerging data from two southern Appalachian crustal blocks give new insights into the palaeogeography of the APGR. The Smith River allochthon may be a part of the APGR on the basis of recently obtained U-Pb monazite and staurolite ages that are apparently incompatible with a Laurentian origin. The allochthon and possibly adjacent terranes, appear to have followed a palaeogeographical track independent of other APGR elements. The Carolina zone is recognized as peri-Gondwanan in origin on the basis of its (i) gross geological evolution, (ii) fossil fauna and (iii) tectonic history. Mid-Palaeozoic regional kinematic patterns suggest that Carolina and its commonly held northern counterpart, the Avalon zone, travelled together on the same lithospheric plate, but their contrasting tectonic histories suggest that they formed along different margins of this plate. These interpretations lead to a new model for middle Palaeozoic interactions of the APGR with Laurentia.

The Appalachian peri-Gondwanan realm (APGR) (Fig. 1) is an extensive tract of exotic Neoproterozoic-early Palaeozoic crustal blocks that occupies the eastern flank of the orogen. The ultimate goal of orogenic analysis of such blocks is to determine significant portions of their palaeogeographical track, including the timing and nature of their accretion, their source craton and the time of their rifting from that source. Recently, there have been significant advances in Neoproterozoic-early Palaeozoic palaeogeographical reconstructions involving Appalachian elements of the APGR (e.g. MacNiocaill et al 1997; van Staal et al. 1998; Cawood et al 2001). These reconstructions tend to consider only the northern components of the APGR and, in particular, the Avalon zone (Fig. 1), mainly because the requisite data for palaeogeographical interpretation are available. Until recently, there have been only limited data available from the areally

extensive southern Appalachian peri-Gondwanan elements that are relevant to such reconstructions. However, the southern Appalachian Carolina zone (Fig. 1) is considered commonly to be equivalent to the Avalon zone (e.g. Williams & Hatcher 1983); consequently, where incorporated in palaeogeographical reconstructions, the Carolina zone is generally lumped, by default, with the Avalon zone (e.g. Murphy & Nance 2002). During the past few years, there has been a marked increase in data that have bearing on the Neoproterozoic-early Palaeozoic palaeogeography of the southern Appalachian Carolina zone and a second possible peri-Gondwanan element, the Smith River allochthon (SRA). This contribution compiles and summarizes emerging data from these two southern Appalachian crustal elements and explores the palaeogeographical implications of these data for the APGR. Two significant conclusions that

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 97-111. 0305-8719/$15.00 © The Geological Society of London 2005.

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Fig. 1. The Appalachian peri-Gondwanan realm and its major constituent elements. can be drawn from this synthesis are that (i) the SRA may be peri-Gondwanan in origin and, if so, it was likely independent of other APGR elements and (ii) the Carolina and Avalon zones may have formed above different boundaries of a common lithospheric plate. The paper concludes with a palaeogeographical model of the early-middle Palaeozoic interactions of the APGR with the eastern Laurentian margin.

Regional setting The APGR spans the southeastern flank of the Appalachian Orogen and includes the northern Appalachian Gander, Avalon and Meguma zones and the southern Appalachian Carolina zone, the subsurface Suwannee terrane and, tentatively, the SRA (Fig. 1). A variety of datasets demonstrates that these crustal elements are exotic to Laurentia and likely formed in a peri-Gondwanan setting (e.g. Wilson 1966; Secor et al 1983; Opdyke et al 1987; Horton et al 1989; Nance et al 1991; van Staal et al 1996; Schenk 1997; Hibbard et al 2002, 2003). The most common palaeogeographical reconstructions of the late Neoproterozoic depict the northern Appalachian APGR components as originating along the ocean-facing margin of west Gondwana (e.g. van Staal et al 1996; Murphy et al 1999; Nance etal 2002). These same models either ignore the

southern APGR or consider it to be contiguous with the northern APGR. The southern APGR elements of concern in this study - the SRA and the Carolina zone border on either native Laurentian rocks or rocks of unknown crustal affinity (Fig. 2). The SRA is a fault-bounded terrane within the western portion of the Piedmont zone, a collection of metamorphic-plutonic terranes of unknown affinity; i.e. available evidence fails to constrain most of these crustal blocks as being either uniquely native or exotic to Laurentia. The western Piedmont zone, including the Jefferson and Potomac terranes (Horton et al 1989) as well as the SRA (Fig. 2), is dominated by metaclastic rocks with subordinate mafic and ultramafic rocks and local eclogite; it has been interpreted commonly as representing the deeper structural levels of an early Palaeozoic accretionary complex (Horton et al 1989; Stewart et al 1997). North of the Virginia promontory (Fig. 1), the eastern Piedmont zone is represented by an Early to Middle Ordovician volcanic arc that was constructed on some form of continental crust (Pavlides et al 1994; Coler et al 2000). This arc may extend south of Virginia, but data bearing on the tectonic affinity of rocks in this region are sparse. The Piedmont zone is in tectonic contact with Laurentian rocks to the west along a series of faults, most of which have accommodated

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Fig. 2. Lithotectonic elements of the southern Appalachian orogen. Cat, Carolina terrane; Cht, Charlotte terrane; CPSZ, central Piedmont shear zone; GZ, Goochland zone; JT, Jefferson terrane; PT, Potomac terrane; SRA, Smith River allochthon. Native Laurentian rocks are represented by unpatterned areas within, and to the west of, the patterned elements.

complex, multiple movement histories ranging from pre-Middle Ordovician to late Palaeozoic (e.g. Adams et al 1995). In addition, Laurentian rocks form a structural inlier in the zone, the Sauratown Mountains window, immediately bordering the SRA to the southeast (Fig. 2). The Carolina zone forms the southeastern exposed flank of the orogen, as it is onlapped by Mesozoic-Cenozoic strata of the Atlantic Coastal Plain to the southeast. The zone comprises Neoproterozoic-early Palaeozoic volcanic arc terranes of peri-Gondwanan origin (e.g. Secor et al. 1983; Hibbard et al 2002). To the north, the zone is in contact with the Goochland zone, a crustal block containing c. 1 Ga orthogneiss of ambiguous provenance (e.g. Owens & Tucker 2003); however, the nature of this contact has been obscured by late

Palaeozoic tectonism (W. Burton pers. comm. 1999). To the east, the Carolina zone is faulted against the eastern Piedmont zone along a profound late Palaeozoic structure termed the central Piedmont shear zone. This shear zone reflects substantial post-accretion telescoping of Carolina over the Piedmont zone (Hibbard et al 1998; Vines et al 1998; West 1998); the locus of original suturing of Carolina to Laurentia is buried beneath the Carolina zone by the central Piedmont shear zone. Two large Neoproterozoic-Cambrian magmatic arc complexes, the Carolina and Charlotte terranes, form the bulk of the Carolina zone (Fig. 2) (Secor et al 1983; Hibbard et al 2002). The Carolina terrane is apparently stitched to the Charlotte terrane by a c. 550 Ma pluton (Barker et al 1998); consequently, any Palaeozoic history that is outlined

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for the Carolina terrane pertains directly to the Charlotte terrane and, thus, to the bulk of the zone.

Palaeogeographical aspects of the Smith River allochthon and the Carolina zone Prior to gauging the extent of the palaeogeographical knowledge of the SRA and the Carolina Zone, it is useful to review briefly the significant aspects of the palaeogeography of exotic terranes. Ideally, one should be able to identify four aspects of their palaeogeographical track, which are analogous to the four aspects of kinematic analysis outlined by Means (1976) (Fig. 3). These aspects include: (i) the distribution or configuration and defining character of the exotic terrane; (ii) the source of the exotic terrane; (iii) the true path, rather than just the shortest path, for the global traverse from source to present location; and, ultimately, iv) the dated path of the terrane, wherein the time at which terranes occupied particular positions along their paths is known. Clearly, it is a rare case in which one can attain the ultimate 'dated path' for an exotic terrane, but in many situations, data pertaining to significant portions of that path can be obtained, including the time of rifting from the source and the time of accretion into the orogen of its final destination. In the following overview of the SRA and the Carolina zone, the discussion of pertinent data will be organized according to the following

Fig. 3. Fundamental aspects of the palaeogeography of exotic terranes. Solid circle represents exotic terrane; open circle represents original position of exotic terrane. Adapted from the aspects of kinematic analysis (Means 1976).

aspects of exotic terrane palaeogeography: (i) distribution and defining character; (ii) source craton, and significant portions of the dated path, including (iii) time of separation from source and (iv) timing and nature of accretion to the Appalachian Orogen.

Smith River allochthon Distribution and defining character. The SRA is located in southwestern Virginia and northwestern North Carolina; it extends for approximately 250 km along-strike and it is less than 50 km wide (Fig. 4). The allochthon is bordered by the Ridgeway Fault, a thrust fault that defines most of its borders, except where it has been truncated by younger structures, such as the reverse-dextral Bowens Creek Fault to the northwest (Conley & Henika 1973) and the Mesozoic brittle Chatham Fault to the southeast. Modelling of potential field data indicates that the allochthon forms a thin, sheet-like mass above these faults (Conley 1985). The allochthon was recognized originally as a distinct fault-bounded crustal block on the basis of its lithostratigraphy and unique tectonothermal and plutonic history (Conley & Henika 1973). In contrast to adjacent Laurentian rocks, the SRA rocks are more pelitic and less quartzose generally and the allochthon contains a melange facies that is distinct from Laurentian melanges. This melange consists of angular to rounded blocks of biotitic quartzite, calc-silicate and amphibolite in a coarse biotite gneiss matrix (e.g. Conley 1985). The main distinguishing feature of the allochthon is its complex tectonothermal history that involves at least three events. The first event is unique to the allochthon and involves relatively high-T, low-P metamorphism that resulted in the growth of staurolite and sillimanite, which are associated with a highly transposed foliation; locally this event appears to have involved migmatization. Gates & Speer (1991) roughly estimated the peak P-T conditions of this metamorphism to be 616-655 °C at 0.61-0.67 GPa. The second event recorded in the allochthon involved the generation of recumbent isoclinal folds with a penetrative axial planar foliation and associated amphibolite facies metamorphism; rough estimates of peak conditions of this metamorphism in the allochthon range from 480-540 °C at 0.58 Gpa (Gates & Speer 1991). The second event is the main fabric-forming event in the region. The last event involved upright folding, multiple crenulation cleavages and retrograde chlorite metamorphism (Gates 1987). The allochthon is intruded by a suite of

Fig. 4. Geological setting of the Smith River allochthon in southwestern Virginia and north central North Carolina. BRA, Blue Ridge antielinorium; SMW, Sauratown Mountains Window

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plutons, termed the Martinsville Intrusive Suite, which includes granitoid and gabbros that have been interpreted as post-dating the second tectonothermal event in the SRA (e.g. Conley 1985). The most recent U-Pb zircon age of c. 445 Ma for granite in the suite (J. Wilson pers. comm. 2001) is consistent, within error, with a recent Th-U-Pb monazite age of 433 ± 12 Ma obtained for the contact aureole of gabbro in the suite (R. Tracy unpublished data). In light of these lithic, structural, metamorphic and plutonic characteristics, rocks of the SRA have been distinguished from bordering Laurentian rocks for a considerable time; however, the potential degree of this distinction was not realized until geochronological studies were launched. In a study focused on the use of staurolite as a chronometer, Lanzirotti & Hanson (1997) reported U-Pb staurolite ages for a pelitic schist from the allochthon of 536 ± 16 Ma for core fractions and 489 ± 28 Ma for rims. Both ages, c. 530 Ma and c. 485 Ma, have been reproduced by Th-U-Pb electron microprobe monazite ages from five pelite samples that span a distance of more than 100km in the allochthon (Hibbard et al 2003; R. Tracy, unpublished data). In addition, monazite cores have yielded ages older than c. 635 Ma and as old as 670 Ma (R. Tracy unpublished data). The staurolite and monazite ages are interpreted as reflecting two distinct tectonothermal events in the allochthon; the Early Cambrian event is viewed as representing the low-P high-T event that is restricted to the allochthon, whereas the Early Ordovician event is considered to represent the main fabric-forming event in the allochthon (Hibbard et al. 2003). The c. 635-670 Ma ages from the monazite cores, as well as the c. 530 Ma ages, are incompatible with the geological history of the eastern Laurentian margin. Ages of c. 635-670 Ma are unknown from the eastern Laurentian margin and lie distinctly between, and outside of, the age ranges of two rift-related magmatic pulses, one at c. 760-680 Ma followed by another at c. 620-550 Ma (e.g. Cawood et al. 2001; Tollo et al. 2004). Likewise, the c. 530 Ma tectonothermal event is difficult to reconcile with eastern Laurentian tectonics of the time; eastern Laurentia was just undergoing the rift-drift transition and would not be expected to be involved in any tectonothermal events (e.g. Hibbard & Samson 1995). Thus, on the basis of these 'un-Laurentian' ages, the SRA has been interpreted to be exotic with respect to Laurentia and most likely of peri-Gondwanan origin (Hibbard et al. 2003). Further studies, including detrital zircon geochronology, are

now in progress in order to test and refine the validity of this hypothesis. In the past, many age dates obtained from plutonic and metamorphic rocks in the Piedmont zone have also hinted at a Cambrian tectonothermal event (e.g. Tilton et al. 1970; Odom & Fullagar 1973; Sinha et al 1989). Many of these data were acquired during the period of rapid development of U-Pb dating techniques, when researchers utilized both mineral separates from multiple samples (in some cases, multiple samples included different rock units) and large multigrain mineral fractions without abrasion. Commonly, results were highly discordant and the resulting ages cannot be interpreted easily. In addition, many of these dates have been supplanted by more modern studies that indicate the rock bodies are significantly younger than the originally reported Cambrian ages (e.g. Aleinikoff et al 2002). Consequently, following a comprehensive review of these Cambrian ages, they are considered to be unreliable and geologically meaningless. However, recent evidence still hints that rocks equivalent to those of the SRA may occur elsewhere in the western Piedmont zone. Southwest of the SRA, in the Jefferson terrane of western North Carolina, Early Cambrian Th-U-Pb electron microprobe monazite ages have been reported recently from metaclastic rocks (D. Moecher, pers. comm. 2003; R. Tracy unpublished data). Northeast along-strike of the SRA, in the Potomac terrane, many workers have reported metaclastic breccias and melanges that are strikingly similar to those in the SRA (e.g. Pavlides 1981; Drake et al 1989; Horton et al 1989). Additionally, 40Ar/39Ar thermochronological data from the terrane hint at a preOrdovician metamorphic event (Drake et al 1999), possibly coeval with that in the SRA. In light of these sparse data, it is conceivable that other portions of the western Piedmont zone may also be of peri-Gondwanan origin. Source craton. The palaeogeographical source of the SRA is open to speculation; the Th-U-Pb monazite data suggest a Gondwanan source area. Many palaeogeographical reconstructions of the earliest Palaeozoic depict the palaeoPacific Gondwanan margin as being proximal to eastern Laurentia and undergoing active orogenesis (e.g. van Staal et al 1998; Cawood et al 2001) (Fig. 5). The early Palaeozoic Gondwanan active margin is now fragmented and forms portions of orogens on at least three continents, including the Pampean Orogen of South America, the Delamerian Orogen of Australia and the Ross Orogen of Antarctica.

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Fig. 5. Schematic palaeogeographical map of Laurentia and Gondwanaland in the Early Cambrian showing the possible source regions for the SRA, and Carolina and Avalon zones. Positions of Carolina and Avalon from Keppie & Ramos (1999). Modified from Cawood et al (2001). AM, Amazonia; AN, Antartica; AUS, Australia; C-SF, Congo-Sao Francisco; IND, India; K, Kalahari; LAUR, Laurentia; RP, Rio de la Plata; WA, West Africa.

Similar to the SRA, the metamorphic peak in the Pampean Orogen involved high-T, lowP metamorphism, with local migmatization that was attained at c. 530 Ma (e.g. Sims et al 1998; Lucassen & Becchio 2003). In contrast, tectonothermal activity in the other orogens, although similar in character to that in the Pampean Orogen, is younger markedly than the Pampean event. Activity in the Delamerian Orogen was concentrated in the span from c. 514-485 Ma (e.g. Foden et al. 1999; 2002), whereas low-pressure metamorphism in the Wilson terrane of the Ross Orogen appears to have peaked at c. 498 Ma (F. Tessensohn pers. comm. 2003). Thus, it is tempting to speculate that the SRA represents a crustal fragment that rifted from the Pampean margin,

traversed the lapetus Ocean and lodged against Laurentia. Time of accretion to Laurentia. Preliminary field observations from a study in progress indicate that the first tectonothermal event shared by both the SRA and adjacent Laurentian rocks is the second event in the SRA, involving recumbent folding, generation of the main foliation axial planar to these folds, and amphibolite facies metamorphism. The c. 485 Ma U-Pb staurolite and monazite ages from the allochthon have been interpreted as recording this event (Hibbard etal 2003). Thus, from preliminary work, it appears that the SRA accreted to Laurentia in the Early Ordovician.

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Time of rifting. The timing of rifting of the SRA from its peri-Gondwanan source cannot be determined directly at this time. However, the analysis above places indirect constraints on the departure of the SRA from the source area; it likely post-dates the Early Cambrian tectonothermal event, which is exotic to Laurentia, and pre-dates its Early Ordovician arrival at the Laurentian margin.

Carolina zone Distribution and defining character. The Carolina zone comprises Neoproterozoic-early Palaeozoic magmatic arc terranes that extend from central Virginia to Alabama. Terranes within the zone are considered to be periGondwanan in origin on the basis of (i) gross geological similarity with other Neoproterozoic-early Palaeozoic peri-Gondwanan terranes (e.g. Secor et al. 1983); (ii) the presence of an

exotic, Acado-Baltic fauna, that is similar to Gondwanan faunas (Samson et al. 1990); and (iii) the presence of tectonothermal events that pre-date formation of eastern Laurentia (e.g. Hibbard & Samson 1995; Dennis & Wright 1997). The geological evolution of the zone is complex and known only partially; most of the knowledge of the zone derives from work within the Carolina and Charlotte terranes and it is distilled into an abridged summary in Figure 6a. A fault-bounded 'old arc' phase is preserved as a magmatic complex with plutonic components dated at c. 670 Ma; this magmatic complex is related spatially to an ophiolitic unit that may represent the local basement (Hibbard et al. 2002). A younger Virgilina arc involves juvenile magmatism, probably formed in an open ocean setting, between c. 630 Ma and 610 Ma (Samson et al. 1995; Wortman et al. 2000). The two youngest arc complexes in the Carolina and Charlotte terranes formed between c. 575

Fig. 6. Interpretative columns depicting the Neoproterozoic-early Palaeozoic lithotectonic evolution of (a) the Carolina zone (E, local eclogite metamorphism; the c. 550 blob represents the pluton that stitches the Carolina and Charlotte arcs); (b) the Avalon zone (adapted from Nance et al. 2002).

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and 535 Ma. The Albemarle arc has been interpreted as overlying unconformably the Virgilina arc, although the contact has not been observed, and magmatism in the Albemarle arc clearly involved a component of older continental crust (e.g. Samson et al 1995). These arcs apparently terminated when they amalgamated at c. 550 Ma (Barker et al 1998); this collision involved deformation in both the Albemarle and Charlotte arcs (Hibbard & Samson 1995; Dennis & Wright 1997; Barker et al 1998) and it was most likely responsible for the generation of the eclogite in the Charlotte arc (Shervais et al 2003). Subsequently, the Albemarle arc was overlain unconformably by Middle Cambrian mudstone of the Asbill Pond Formation containing Acado-Baltic trilobites (Secor et al 1983; Samson et al 1990). The crustal fragment that contains the Carolina zone may extend beyond the southern Appalachians; commonly, the zone is considered to be the southern extension of the northern Appalachian Avalon zone (e.g. Williams & Hatcher 1983; Ingle et al 2003) mainly on the basis of roughly coeval arc magmatism in both zones. In counterpoint, others have noted significant stratigraphic disparities between the zones (Secor et al 1983; Hibbard et al 2002). The geological evolution of Avalon is similar grossly to that of Carolina, but their paths appear to diverge near the Neoproterozoic-Cambrian boundary (Fig. 6a, b). In particular, there appears to be a contrast in the nature of the temination of arc magmatism. Cessation of Avalonian arc magmatism appears to have involved ridge subduction followed by a transition into a strike-slip setting (Nance etal 2002). In contrast, Carolina zone magmatism appears to have been terminated by a collision of the Carolina and Charlotte arcs. In Avalon, the strike-slip tectonic setting was succeeded transitionally by a robust, early Palaeozoic platform sequence (e.g. Nance et al 2002), whereas, in the Carolina zone, Middle Cambrian mudstone rests unconformably on the Albemarle arc. In Carolina, the mudstone may represent the vestiges of a platform, but there is no documentation of a preceding strike-slip tectonic regime. Clearly, the contrasting Neoproterozoic-early Palaeozoic tectonic settings of the Avalon and Carolina zones can be reconciled with their mutual correlation by a scenario involving either different tectonic regimes along the strike of an extensive arc system or different tectonic regimes flanking different margins of a lithospheric plate. Alternatively, the two zones may not be correlative and may represent palaeogeographically distinct elements.

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Source craton. Investigations seeking the source area within Gondwana for the Carolina zone are limited and confined largely to geochronologic and isotopic comparisons. Nance & Murphy (1996) indicated that the limited available Sm-Nd data for the Carolina zone were compatible with the more robust dataset of the Avalon zone that includes generally positive 8Nd values and depleted mantle Nd model ages in the range of 0.8-1.1 Ga. They indicated that these isotopic data pointed to a source within the Amazonian craton, possibly similar to the Amazonian Tocantins province. More recent Nd data, as well as xenocrystic zircon data from Carolina plutons, are consistent with this interpretation (Ingle et al 2003). Time of accretion to Laurentia. The timing and nature of accretion of the Carolina zone to Laurentia is obscured because the tectonic contact between the the Carolina zone and the remainder of the orogen to the west, the central Piedmont shear zone, is a late Palaeozoic thrust fault that has buried the original suture beneath the Carolina zone (Hibbard et al 1998; Wortman et al 1998). Consequently, the modern contact of the zone with the orogen lacks a record of the nature and timing of the suture. However, a recent model involving Late Ordovician-Silurian sinistral transpressive docking of Carolina satisfies most of the evidence that bears on this problem (Hibbard 2000; Hibbard et al 2002). In contrast to previous proposed scenarios, this model demonstrates that the major pre-Middle Ordovician crustal blocks of the southern Appalachians were tectonically linked at this time; the model is also consistent with both a Late Ordovician timing of regional tectonism in the Carolina terrane and with palaeomagnetic data from two independent studies that indicate the Carolina zone was at Laurentian palaeolatitudes at that time (Vick et al 1987; Noel et al 1988). Time of rifting. Direct evidence for the timing of departure of the Carolina zone from Gondwana is unavailable, but it is constrained to the timespan between deposition of the Middle Cambrian Acado-Baltic fauna and Late Ordovician-Silurian docking to Laurentia. Intriguingly, gabbroic sills, dykes and stocks of the Stony Mountain suite (Hibbard et al 2002) intrude all of the units within the Albemarle arc, yet are involved in Late Ordovician deformation and metamorphism related to docking. Deductively, because the suite is the last magmatic event recorded in the arc prior to

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docking, it may well be related to rifting from Gondwana. Geochemical and geochronological studies are in progress in order to investigate this hypothesis.

Palaeogeographical implications The compilation above allows reconstruction of the palaeogeography of the southern APGR. It demonstrates that (i) the SRA is tentatively peri-Gondwanan and distinct from other APGR elements; (ii) other portions of the Piedmont zone may also be of peri-Gondwanan origin; and (iii) the Carolina and Avalon zones display contrasting tectonic histories but appear to share a common source region. At present, the data are insufficient to identify confidently either the specific source areas for the southern APGR or the details of earlier portions of their displacement paths, but with these data one can begin to construct a model for their arrival at the eastern Laurentian margin as well as their relationship to northern APGR components. Evaluation of these preliminary data suggests that the SRA had an origin along the Pampean margin of South America, well removed from the suggested sites of origin for other APGR components. Accentuating this difference, the SRA also appears to have arrived at the Laurentian margin in the Early Ordovician, well ahead of other APGR elements; all of the other APGR elements are believed to have arrived either during or after the Late Ordovician (e.g. van der Pluijm et al 1993; Hibbard 1994; 2000; van Staal et al. 1996; 1998). Thus, the SRA appears to be distinct palaeogeographically from other elements in the APGR. Considering that the Potomac and Jefferson terranes display tectonothermal and lithological similarities with the SRA, it is conceivable that these portions of the Piedmont zone may have a peri-Gondwanan origin and may have accreted to Laurentia synchronously with the SRA. The c. 485 Ma, Tremadocian, age of deformation and metamorphism interpreted to reflect accretion of the SRA, possibly accompanied by other Piedmont zone components, heralds the late Tremadocian-middle Arenig inception of tectonic instability on the Laurentian platform in the southern Appalachians (Drake et al 1989). Subsequent destruction of the passive margin and deposition of the Blount clastic wedge clearly pre-date Taconic events in the northern Appalachians; traditionally, these early southern phenomena have been termed collectively the Blountian event (Kay 1942) and viewed as a precocious phase of the Taconic

Orogeny (Rodgers 1953; Drake et al. 1989). Also, in contrast to the southern Appalachian event, the northern Taconic orogeny involved the accretion of a peri-Laurentian magmatic arc constructed possibly on a microcontinental block rifted from Laurentia (e.g. Waldron & van Staal 2001; Cawood et al 2001). In both regions, convergence appears to have involved the attempted subduction of the Laurentian margin (e.g. Stevens 1970; Drake et al 1989). If the analysis above is accurate, a logical conclusion to draw from these observations is that two distinct lithospheric plates resided outboard of the eastern Laurentian margin; the southern one containing peri-Gondwanan crustal elements (SRA, Piedmont zone [?]) that were obducted onto Laurentia immediately prior to the obduction of peri-Laurentian crustal components onto the northern margin (Fig. 7a). The hypothetical boundary between these plates is depicted to be a transform fault, mainly for the sake of simplicity; however, it could just as well have been a convergent boundary. Subsequent to eastward subduction of the Laurentian margin, it appears that the polarity of subduction reversed along the length of the Appalachians, leading to subduction beneath North America (van Staal et al 1998; Hibbard 2000). In the southern Appalachians, the Carolina zone appears to have arrived at the North American margin soon after this subduction polarity reversal; the nature of accretion has been interpreted to involve sinistral transpression (Hibbard 2000; Hibbard et al 2002). This style of accretion is consistent with the kinematic pattern of accretion of the Gander and Avalon zones in the northern Appalachians (e.g. Keppie 1992; Hibbard 1994; Piasecki 1995; Dube et al 1996). Thus, following subduction reversal, peri-Gondwanan terranes began to impinge on the eastern Laurentian margin in a sinistral oblique convergent manner. This kinematic pattern is consistent with the distribution of intense Late Ordovician to Silurian deformation and metamorphism in the orogen. The southern edge of the St. Lawrence promontory is the locus of Late Ordovician-Silurian blueschist metamorphism and ductile deformation in New Brunswick (van Staal et al. 1990) and southern Newfoundland was subjected to Silurian thrusting and amphibolite facies metamorphism (O'Brien et al 1991). Likewise, recent U-Pb ages from xenotime, titanite and monazite in southern New England (Sevigny & Hanson 1995; C. Dietsch, pers. comm. 2003; R. Tracy unpublished data,) indicate the presence of intense Silurian tectonism along the New York promontory. This pattern may well reflect

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Fig. 7. Schematic model for the arrival of the Appalachian peri-Gondwanan realm elements in (a) the Early Ordovician and (b) the Late Ordovician-Silurian. Bold arrow represents hypothetical plate vector for the peri-Gondwanan plate with respect to Laurentia.

sinistral oblique convergence of Laurentia with the peri-Gondwanan terranes, wherein the promontories act as restraining bends that accumulate intense shortening. In summary, both the presence of peri-Gondwanan crustal fragments, possibly from the same source region, along the entire eastern Laurentian margin and the consistent sinistral oblique kinematic regime from south to north implies that the outboard terranes were conveyed upon a

single lithospheric plate outboard of North America (Fig. 7b). Although the Carolina zone appears to arrive at the Laurentian margin on the same plate as the Avalon zone, differences between these two APGR elements as outlined above, suggest that they originated along different plate boundaries. In particular, Carolina zone arc magmatism appears to have terminated with arc collision, whereas Avalonian arc magmatism

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ceased with a transition from convergence to a strike-slip regime (e.g. Nance et al. 2002). Additionally, the Avalon zone exhibits a robust early Palaeozoic platformal sequence, which has no counterpart in the Carolina zone (Secor et al. 1983; Hibbard et al. 2002). Finally, it has been noted that although both zones contain Cambrian Acado-Baltic faunas, these faunas may well represent contrasting subprovinces (Nance & Murphy 1996). Palaeomagnetic data allow for a partial reconstruction of the palaeogeography of this peri-Gondwanan plate which contained the two largest elements in the APGR in the early Late Ordovician. Two independent palaeomagnetic studies indicate that the Carolina terrane was at a palaeolatitude of c. 22° S, i.e. at the same palaeolatitude as Laurentia, at c. 455 Ma, when the terrane experienced peak metamorphic conditions (Vick et al 1987; Noel et al 1988). In contrast, the Avalon zone appears to have been at a palaeolatitude of c. 41° S (Johnson & van der Voo 1990) at the time of eruption of the c. 460 Ma Dunn Pt volcanics in the Arisaig region of Nova Scotia (Hamilton & Murphy 2004). Thus, the Carolina and Avalon zones appear to have a minimal latitudinal separation on the order of 2000km in the earliest Late Ordovician. It is interesting to note that the modern direct distance between sites sampled for palaeomagnetic data is on the order of 1900 km, nearly the same as their minimum palaeolatitudinal separation. Hence, the Carolina and Avalon zones may have formed an extensive microcontinental archipelago that was situated athwart the length of the lapetus Ocean. Outboard APGR elements such as the Meguma zone of Nova Scotia and the Suwanee terrane in the subsurface of Florida were accreted later in the Palaeozoic. The first shared deformation between the Meguma zone and other rocks of the orogen appears to be the Early-Middle Devonian, or classic, Acadian (e.g. Keppie 1992), suggesting intitial interaction of the two crustal blocks at that time. Docking of the Suwannee terrane to Laurentia is generally considered to be late Palaeozoic (Dallmeyer et al. 1987; McBride & Nelson 1991). Concluding remarks Compilation of relevant data indicates that southern APGR elements shed new light on palaeogeographical interpretation of the entire realm; in particular, the tentatively periGondwanan SRA and possibly affiliated

portions of the Piedmont zone appear to have a source and palaeogeographical track distinct from other elements. In contrast, the Carolina zone appears to have been at the leading edge of a peri-Gondwanan plate that included the Gander and Avalon zones of the northern APGR, although distinctions between Carolina and Avalon may be related to positioning on this plate. This southern perspective on APGR palaeogeography only serves to raise new questions that need to be addressed in future studies; e.g. what are the specific source areas of the SRA, Carolina, Avalon and other APGR elements? Did they separate at approximately the same time from their Gondwanan source? Did they separate as a single plate, or piecemeal as many blocks? More data of the form such as detailed geological correlations, focused detrital zircon studies and geochronologic, palaeomagnetic and palaeontological studies are needed in order to begin to address these concerns, not only in the Appalachians, but also in the potential Gondwanan source regions. JH appreciates the invitation from Alan Vaughan and Phil Leat to present this work at the TAPMOG conference; he also acknowledges partial funding from the US National Science Foundation (EAR0106112, EAR 01124245) in support of this research. RT acknowledges partial funding from the US National Science Foundation (EAR-0124212). All authors express their gratitude to W. S. Henika for his generosity in sharing both his knowledge of the SRA and uncountable hours in the field. Thanks go to Dov Avigad, Brendan Murphy and Cees van Staal for insightful reviews that helped to sharpen the manuscript.

References ADAMS, M., STEWART, K., TRUPE, C. & WILLARD, R. 1995. Tectonic significance of high-pressure metamorphic rocks and dextral strike-slip faulting along the Taconic suture. In: HIBBARD, X, VAN STAAL, C. & CAWOOD, P. (eds) New Perspectives in the Appalachian-Caledonian Orogen. Geological Association of Canada Special Paper, 41, 21-42. ALEINIKOFF, J., HORTON, W., DRAKE, A. & FANNING, M. 2002. SHRIMP and conventional U-Pb ages of Ordovician granites and tonalites in the central Appalachian Piedmont: Implications for Paleozoic tectonic events. American Journal of Science, 302, 50-75. BARKER, C., SECOR, D.T., JR, PRAY, J. & WRIGHT, J. 1998. Age and deformation of the Longtown metagranite, South Carolina Piedmont: A possible constraint on the origin of the Carolina terrane. Journal of Geology, 106, 713-725. CAWOOD, P., MCCAUSLAND, P. & DUNNING, G. 2001. Opening lapetus: Constraints from Laurentian

APPALACHIAN PERI-GONDWANAN REALM margin in Newfoundland. Geological Society of America Bulletin, 113, 443-453. COLER, D., SAMSON, S. & HIBBARD, J. 2000. The Chopawamsic and Milton terranes: Correlative Ordovician arc terranes built on ancient continental crust, southern Appalachians. Journal of Geology, 108, 363-380. CONLEY, J.F. 1985. Geology of the Southwestern Virginia Piedmont. Virginia Division of Mineral Resources Publication 59. CONLEY, J.F. & HENIKA, W.S. 1973. Geology of the Snow Creek, Martinsville East, Price, and Spray quadrangles. Virginia Division of Mineral Resources Report of Investigations 33. DALLMEYER, D., CAEN-VACHETTE, M. & VILLENEUVE, M. 1987. Emplacement age of postmetamorphic granites in southern Guinea (West Africa) and the peninsular Florida subsurface: Implications for origins of southern Appalachian exotic terranes. Geological Society of America Bulletin, 99, 87-93. DENNIS, A.J. & WRIGHT, I.E. 1997. The Carolina terrane in northwestern South Carolina, USA: age of deformation and metamorphism in an exotic arc. Tectonics, 16, 460-473. DRAKE, A., SINHA, K., LAIRD, J. & GUY, R. 1989. The Taconic orogen. In: HATCHER, R., THOMAS, W. & VIELE, G. (eds) The Appalachian-Ouachita Orogen in the United States. Geological Society of America, Boulder, Colorado, The Geology of North America, F-2,101-177. DRAKE, A., SOUTHWORTH, S. & LEE, K. 1999. Geologic map of the Seneca quadrangle, Montgomery County, Maryland, and Fairfax and London Counties, Virginia. United States Geological Survey Geologic Quadrangle Map GQ-1802. DUBE, B., DUNNING, G, LAUZIERE, K. & RODDICK, J. 1996. New insights into the Appalachian Orogen from geology and geochronology along the Cape Ray fault zone, southwest Newfoundland. Geological Society of America Bulletin, 108, 101-116. FODEN, J., SANDIFORD, M., DOUGHERTY-PAGE, J. & WILLIAMS, I. 1999. Geochemistry and geochronology of the Rathjen Gneiss: implications for the early tectonic evolution of the Delamerian Orogen. Australian Journal of Earth Sciences, 46, 377-389. FODEN, J., ELBURG, M., TURNER, S., SANDIFORD, M., O'CALLAGHAN, J. & MITCHELL, S. 2002. Granite production in the Delamerian Orogen, South Australia. Journal of the Geological Society, London, 159, 557-575. GATES, A. 1987. Transpressional dome formation in the southwest Virginia Piedmont. American Journal of Science, 287, 927-949. GATES, A. & SPEER, A. 1991. Allochemical retrograde metamorphism in shear zones: an example in metapelites, Virginia, USA. Journal of Metamorphic Geology, 9, 581-604. HAMILTON, M. & MURPHY, J.B. 2004. Tectonic significance of a Llanvirn age for the Dunn Point volcanic rocks, Avalon terrane, Nova Scotia, Canada: implications for the evolution of the

109

lapetus and Rheic oceans. Tectonophysics, 379, 199-209. HIBBARD, J. 1994. Kinematics of Acadian deformation in the northern and Newfoundland Appalachians. Journal of Geology, 102, 215-228. HIBBARD, J. 2000. Docking Carolina: Mid Paleozoic accretion in the southern Appalachians. Geology, 28, 127-130. HIBBARD, J. & SAMSON, S. 1995. Orogenesis exotic to the lapetan cycle in the southern Appalachians. In: HIBBARD, I, VAN STAAL, C. & CAWOOD, P. (eds) Current Perspectives in the AppalachianCaledonian Orogen. Geological Association of Canada Special Paper 41,191-205. HIBBARD, I, SHELL, G, BRADLEY, P., SAMSON, S. & WORTMAN, G. 1998. The Hyco shear zone in North Carolina and southern Virginia: Implications for the Piedmont zone-Carolina zone boundary in the southern Appalachians. American Journal of Science, 298, 85-107. HIBBARD, J., STODDARD, E., SECOR, D. & DENNIS, A. 2002. The Carolina Zone: overview of Neoproterozoic to early Paleozoic peri-Gondwanan terranes along the eastern flank of the southern Appalachians. Earth Science Reviews, 57,299-339. HIBBARD, I, TRACY, R. & HENIKA, W. 2003. Smith River allochthon: a southern Appalachian periGondwanan terrane emplaced directly on Laurentia? Geology, 31, 215-218. HORTON, W., DRAKE, A. & RANKIN, D. 1989. Tectonostratigraphic terranes and their Paleozoic boundaries in the central and southern Appalachians. In: DALLMEYER, D. (ed.) Terranes in the CircumAtlantic Paleozoic Orogens. Geological Society of America Special Paper, 230, 213-245. INGLE, S., MULLER, P., HEATHERINGTON, A. & KOZUCH, M. 2003. Isotopic evidence for the magmatic and tectonic histories of the Carolina terrane: implications for stratigraphy and terrane affiliation. Tectonophysics, 371, 187-211. JOHNSON, R. & VAN DER Voo, R. 1990. Pre-folding magnetization reconfirmed for the Late Ordovician-Early Silurian Dunn Point volcanics, Nova Scotia. Tectonophysics, 178, 193-205. KAY, M. 1942. Development of the northern Allegheny synclinorium and adjoining regions. Geological Sociey of America Bulletin, 53, 1601-1657. KEPPIE, J.D. 1992. Structure of the Canadian Appalachians. Nova Scotia Department of Natural Resources, Mines and Energy Branch Bulletin 7. KEPPIE, J.D. & RAMOS, VA. 1999. Odyssey of terranes in the lapetus and Rheic Oceans during the Paleozoic. In: KEPPIE, J.D. & RAMOS, VA. (eds) Laurentia-Gondwana Connections before Pangea. Geological Society of America Special Paper, 336, 267-276. LANZIROTTI, A. & HANSON, G.N. 1997. An assessment of the utility of staurolite in U-Pb dating of metamorphism. Contributions to Mineralogy and Petrology, 129, 352-365. LUCASSEN, F. & BECCHIO, R. 2003. Timing of highgrade metamorphism: Early Paleozoic U-Pb

110

J. P. HIBBARD ETAL. formation ages of titanite indicate long-standing high-T conditions at the western margin of Gondwana (Argentina, 26-29°S). Journal of Metamorphic Geology, 21, 649-662.

MAC NlOCAILL, C, VAN DER PlAJIJM, B.A. & VAN DER

Voo, R. 1997. Ordovician paleogeography and the evolution of the lapetus ocean. Geology, 25, 159-162. MCBRIDE, J. & NELSON, D. 1991. Deep seismic reflection constraints on Palaeozoic crustal structure and definition of the Mono in the buried Southern Appalachian orogen. In: MEISSNER, R., BROWN, L., DURBAUM, H.-J., FRANKE, W., FUCHS, K. & SEIFERT, F. (eds) Continental Lithosphere: Deep Seismic Reflections. American Geophysical Union, Geodynamics Series, 22, 9-20. MEANS, W. 1976. Stress and Strain: Basic Concepts of Continuum Mechanics for Geologists. SpringerVerlag, New York. MURPHY, J.B. & NANCE, D. 2002. Sm-Nd isotopic systematics as tectonic tracers: an example from West Avalonia in the Canadian Appalahcians. Earth Science Reviews, 59, 77-100. MURPHY, J.B., KEPPIE, J.D., DOSTAL, J. & NANCE, D. 1999. Neoproterozoic-early Paleozoic evolution of Avalonia. In: RAMOS, V. & KEPPIE, D. (eds) Laurentia-Gondwana connections before Pangea. Geological Society of America Special Paper, 336, 253-266. NANCE, D. & MURPHY, B. 1996. Basement isotopic signatures and Neoproterozoic paleogeography of Avalonian-Cadomian and related terranes in the circum-North Atlantic. In: NANCE, D. & THOMPSON, M. (eds) Avalonian and Related PeriGondwanan Terranes of the Circum-North Atlantic. Geological Society of America Special Paper, 304, 333-346. NANCE, D., MURPHY, B., STRACHAN, R., D'LEMOS, R. & TAYLOR, G. 1991. Late Proterozoic tectonostratigraphic evolution of the Avalonian and Cadomian terranes, Precambrian Research, 53, 41-78. NANCE, D., MURPHY, J.B. & KEPPIE, J.D. 2002. A Cordilleran model for the evolution of Avalonia. Tectonophysics, 352, 11-31. NOEL, J., SPARIOSU, D. & DALLMEYER, D. 1988. Paleomagnetism and Ar/Ar ages from the Carolina slate belt, Albemarle, North Carolina: Implications for terrane amalgamation with North America. Geology, 16, 64-68. O'BRIEN, B., O'BRIEN, S. & DUNNING, G. 1991. Silurian cover, Late Precambrian-Early Ordovician basement, and the chronology of Silurian orogenesis in the Hermitage Flexure (Newfoundland Appalachians). American Journal of Science, 291, 760-799. ODOM, L. & FULLAGAR, P. 1973. Geochronologic and tectonic relationships between the Inner Piedmont, Brevard Zone, and Blue Ridge belts, North Carolina. American Journal of Science, 273-A, 133-149. OPDYKE, N, JONES, D., MACFADDEN, B., SMITH, D., MULLER, P. & SHUSTER, R. 1987. Florida as an exotic terrane: Paleomagnetic and geochrono-

logic investigation of lower Paleozoic rocks from the subsurface of Florida. Geology, 15, 900-903. OWENS, B. & TUCKER, R. 2003. Geochronology of the Mesoproterozoic State Farm Gneiss and associated Neoproterozoic granitoids, Goochland terrane, Virginia. Geological Society of America Bulletin, 115, 972-982. PAVLIDES, L. 1981. The Central Virginia volcanicplutonic belt: An island arc of Cambrian (?) age. United States Geological Survey Professional Paper 1231-A. PAVLIDES, L., ARTH, J. , SUITER, J., STERN, T. & CORTESINI, H. 1994. Early Paleozoic alkalic and calc-alkalic plutonism and associated contact metamorphism, central Virginia Piedmont. United States Geological Survey Professional Paper 1529. PIASECKI, M. 1995. Dunnage zone boundaries and some aspects of terrane development in Newfoundland. In: HIBBARD, J., VAN STAAL, C. & CAWOOD, P. (eds) Current Perspectives in the Appalachian-Caledonian Orogen. Geological Association of Canada Special Paper, 41, 323-347. RODGERS, J. 1953. Geologic map of East Tennessee with explanatory text. Tennessee Division of Geology Bulletin, 58, Part II. SAMSON, S., PALMER, A., ROBISON, R. & SECOR, D.T. 1990. Biogeographical significance of Cambrian trilobites from the Carolina slate belt. Geological Society of America Bulletin, 102, 1459-1470. SAMSON, S., HIBBARD, J. & WORTMAN, G. 1995. Nd isotopic evidence for juvenile crust in the Carolina terrane, southern Appalachians. Contributions to Mineralogy and Petrology, 121, 171-184. SCHENK, P. 1997. Sequence stratigraphy and provenance on Gondwana's margin: The Meguma Zone (Cambrian to Devonian) of Nova Scotia, Canada. Geological Society of America Bulletin, 109, 395-409. SECOR, D.T, SAMSON, S., SNOKE, A. & PALMER, A. 1983. Confirmation of the Carolina slate belt as an exotic terrane. Science, 221, 649-651. SEVIGNY, J. & HANSON, G. 1995. Late Taconian and pre-Acadian history of the New England Appalachians of southwestern Connecticut. Geological Society of America Bulletin, 107, 487^98. SHERVAIS, J., DENNIS, A., McGEE, J. & SECOR, D. 2003. Deep in the heart of Dixie: Pre-Alleghanian eclogite and HP granulite metamorphism in the Carolina terrane, South Carolina, USA. Journal of Metamorphic Geology, 21, 65-80. SIMS, J., IRELAND,T, CAMACHO, A. ETAL. 1998. U-Pb, Th-Pb, and Ar-Ar geochronology form the southern Sierras Pampeanas, Argentina: implications for the Palaeozoic tectonic evolution of the western Gondwana margin. In: PANKHURST, R. & RAPELA, C. (eds) The Proto Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 259-281. SINHA, A., HUND, E. & HOGAN, J. 1989. Paleozoic accretionary history of the North American plate

APPALACHIAN PERI-GONDWANAN REALM margin (central and southern Appalachians): constraints from the age, origin, and distribution of granitic rocks. In: HILLHOUSE, J. (ed.) Deep Structure and Past Kinematics of Accreted Terranes. American Geophysical Union Geophysical Monograph, 50, 219-238. STEVENS, R.K. 1970. Cambro-Ordovician flysch sedimentation and tectonics in west Newfoundland and their possible bearing on a proto-Atlantic ocean. In: LAJOIE, J. (ed.) Flysch Sedimentation in North America. Geological Association of Canada Special Paper 7,165-177. STEWART, K.G, ADAMS, M.G. & TRTJPE, C.H. 1997. Paleozoic structure, metamorphism, and tectonics of the Blue Ridge of western North Carolina. Carolina Geological Society 1997 Field Trip Guidebook. TILTON, G, DOE, B. & HOPSON, C. 1970. Zircon age measurements in the Maryland Piedmont, with special reference to Baltimore Gneiss problems. In: FISHER, G, PETTIJOHN, F, REED, J. & WEAVER, K. (eds) Studies of Appalachian Geology: Central and Southern. Interscience Publishers, New York, 429-^34. TOLLO, R., ALEINIKOFF, J., BARTHOLOMEW, M. & RANKIN, D. 2004. Neoproterozoic A-type granitoids of the central and southern Appalachians: intraplate magmatism associated with episodic rifting of the Rodinian supercontinent. Precambrian Research, 128, 3-38. VAN DER PLUIJM, B., JOHNSON, R. & VAN DER Voo, R. 1993. Paleogeography, accretionary history, and tectonic scenario: A working hypothesis for the Ordovician and Silurian evolution of the northern Appalachians. In: ROY, D. & SKEHAN, J. (eds) The Acadian Orogeny. Geological Society of America Special Paper, 275, 27-40. VAN STAAL, C., RAVENHURST, C., WINCHESTER, J., RODDICK, J. & LANGTON, J. 1990. Post-Taconic blueschist suture in the northern Appalachians of northern New Brunswick, Canada. Geology, 18, 1073-1077. VAN STAAL, C., SULLIVAN, R. & WHALEN, J. 1996. Provenance and tectonic history of the Gander Zone in the Caledonian/Appalachian orogen: Implications for the origin and assembly of Avalon. In: NANCE, D. & THOMPSON, M. (eds)

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Avalonian and Related Peri-Gondwanan Terranes of the Circum-North Atlantic. Geological Society of America Special Paper, 304, 347-367. VAN STAAL, C., DEWEY, I, MAC NIOCAILL, C. & McKERROW, W. 1998. The Cambrian-Silurian tectonic evolution of the northern Appalachians and British Caledonides: History of a complex, west and southwest Pacific-type segment of lapetus. In: BLUNDELL, D. & SCOTT, A. (eds) Lyell: The Past is the Key to the Present. Geological Society, London, Special Publications, 143,199-242. VICK, H., CHANNELL, J. & OPDYKE, N. 1987. Ordovician docking of the Carolina slate belt: Paleomagnetic data. Tectonics, 6, 573-583. VINES, J., HIBBARD, J. & SHELL, G. 1998. Structural geology of the High Rock granite: Implications for displacement along the Hyco shear zone, North Carolina. Southeastern Geology, 37, 163-176. WALDRON, J.W.F. & VAN STAAL, C.R. 2001. Taconic Orogeny and the accretion of the Dashwoods block: a peri-Laurentian microcontinent in the lapetus Ocean. Geology, 29, 811-815. WEST, T. 1998. Structural analysis of the CarolinaInner Piedmont terrane boundary: Implications for the age and kinematics of the central Piedmont suture, a terrane boundary that records Paleozoic Laurentia-Gondwana interactions. Tectonics, 17, 379-394. WILLIAMS, H. & HATCHER, R.D. 1983. Appalachian suspect terranes. In: HATCHER, R.D., JR, WILLIAMS, H. & ZIETZ, I. (eds) Contributions to the Tectonics and Geophysics of Mountain Chains. Geological Society of America Memoir, 158, 33-54. WILSON, IT. 1966. Did the Atlantic close and then reopen? Nature, 211, 676-681. WORTMAN, G, SAMSON, S. & HIBBARD, J. 1998. Precise U-Pb timing constraints on the kinematic development of the Hyco shear zone, southern Appalachians. American Journal Science, 298, 108-130. WORTMAN, G, SAMSON, S. & HIBBARD, J. 2000. Precise U-Pb zircon constraints on the earliest magmatic history of the Carolina terrane. Journal of Geology, 108, 321-338.

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Nd and Sr isotopic signatures of metasedimentary rocks around the South Pacific margin and implications for their provenance C. J. ADAMS1, R. J. PANKHURST2, R. MAAS3 & I. L. MILLAR4 ^Institute of Geological & Nuclear Sciences, PO Box 30368, Lower Hutt, New Zealand (e-mail: [email protected]) 2 NERC Isotope Geosciences Laboratory, British Geological Survey, Key worth, Nottingham, NG12 5GG, UK (e-mail: [email protected]) ^School of Earth Sciences, University of Melbourne, Victoria 3010, Australia (e-mail: [email protected]) ^British Antarctic Survey, High Cross, Madingley Rd, Cambridge CBS OET, UK (e-mail: [email protected]) Abstract: An Nd-Sr isotope database, including c. 200 new analyses, is presented for Palaeozoic and Mesozoic metasedimentary successions extending through southeastern Australia, New Zealand, West Antarctica and the Antarctic Peninsula to southern South America. Combining with U-Pb detrital zircon age data, this enables characterization of New Zealand terranes, especially within the Eastern Province, where there is a progression from westernmost terranes of both volcanic/volcaniclastic and accretionary origin with primitive sediment sources, to easternmost terranes with mature continental sediment inputs. In southern South America, West Antarctica and the Antarctica Peninsula, similar accretionary complexes have Nd model ages principally reflecting mixing of sedimentary material from multiple sources, both mature and juvenile. A mature Gondwana continental provenance dominates in sedimentary basins inboard of the active margin, especially in the Palaeozoic (Western Province, New Zealand, interior West Antarctica and NW Argentina), and contributes significantly to pre-Mesozoic sedimentary rocks of Patagonia east of the Andes. Along the Gondwanaland margin, Nd systematics for younger (late Palaeozoic to early Mesozoic) accretionary complex metasediments reflect younger source inputs, notably in the Scotia metamorphic complex. Many of the accretionary complex deposits must involve significant crustal reworking. There is no apparent South American equivalent of the primitive provenance of the westernmost accretionary terranes of New Zealand.

The margin of East Gondwanaland occurs close to the present South Pacific Ocean margin, as late Neoproterozoic, early Palaeozoic, late Palaeozoic or late Mesozoic metasedimentary successions (Fig. 1). In the New Zealand, southeast Australia and Ross Sea (Antarctica) sectors, these are defined mainly as tectonostratigraphic terranes, first, on the basis of their stratigraphy, structure and metamorphism and, secondly, by their tectonic association, as indicated by geochemical and radiogenic isotope attributes (in particular Nd, Sr, but also Hf). The latter characteristics have been explored further by detrital mineral age studies (mainly U-Pb zircon dating), which help to identify sediment sources and thus discriminate between autochthonous and allochthonous terranes. Although not as well established as those in New Zealand and Australia, similar tectono-

stratigraphic terranes have been hypothesized in West Antarctica, the Antarctic Peninsula and southern South America. This work establishes a comprehensive new Sr and Nd isotope database for the New Zealand region, in particular, and integrates it with commensurate data available for the adjacent sectors of the East Gondwanaland margin, as a basis for comparative terrane studies, j\ew Zealand, southeast Australia and Antarctica The East Gondwanaland margin in New Zealand comprises (1) an early Palaeozoic, Western Province, continental foreland, which is an extension of the Lachlan Fold Belt of southeast Australia, and is separated by a Median Tectonic Zone (Bradshaw 1993) or

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246,113-141. 0305-8719/$15.00 © The Geological Society of London 2005.

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Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN

Median Batholith (Mortimer et al 19990, b) from (2) a late Palaeozoic to Mesozoic, Eastern Province, a mobile belt which consists of several tectonostratigraphic terranes (Bishop et al. 1985; Figs 2, 3). The Western Province (Fig. 2) comprises the extensive Duller terrane, dominated by quartzose turbiditic greywackes of the Ordovician Greenland Group (Laird & Shelley 1975; Nathan 1976), and the smaller, more varied Takaka terrane, with Cambrian island-arc volcanic rocks overlain by Cambrian to Devonian successions including quartzite, black shale and carbonates (marble) (Cooper 1979; 1989; Roser et al. 1996; Rattenbury et al. 1998; Miinker & Cooper 1999). Buller terrane metasediments have been compared with those in the western Lachlan Fold Belt in Victoria (Cooper & Tulloch 1992). In Victoria (Fergusson & VandenBerg 2003) and southern New South Wales (Glen et al 1992), this comprises very extensive tracts of Palaeozoic (principally Cambrian to Silurian) low- to medium-grade metasediments divided into several provinces, and equally extensive granitoid terranes (Fig. 1). The most extensive metasediments are of a quartzose turbidite type, but there are important local volcanic units, and limestone and pelagic mudstone/chert environments are present (Birch 2003). The quartzite-black shale Ordovician environments of the Castlemaine Group, in the Bendigo Zone of Victoria, and the voluminous quartz turbidites to the west (Fergusson & VandenBerg 2003), are very similar to the eastern part of the Greenland Group. However, less fossiliferous Ordovician successions are now known in eastern Victoria. There, the Pinnak Sandstone, part of the Adaminaby Group (Fergusson & VandenBerg 2003), is similar to the Greenland Group. Further correlations have been made with the Robertson Bay Group in North Victoria Land and Swanson Formation in Marie Byrd Land, Antarctica (Bradshaw et al. 1983; Adams 1986; 1997; 2004; Fig. 1). Possible equivalents of the Takaka terrane occur in western Tasmania (Corbett et al. 1989; Cooper & Tulloch 1992),

115

the Grampians-Stavely zone of SE Australia (Stump et al 1986) and the Bowers terrane in North Victoria Land (Cooper 1979, 1989; Bradshaw et al 1983). The early Palaeozoic terranes in New Zealand, southeastern Australia and Antarctica were accreted to the Neoproterozoic-early Palaeozoic, Ross and Delamerian Fold Belts of the Transantarctic Mountains and South Australia, respectively (Fig. 1), a process completed during the mid-Devonian (e.g. Foster & Gray 2000; Fergusson 2003). Together, the Ross-Delamerian and Lachlan Fold Belts and other early Palaeozoic sections of the Tasmanides (Coney etal 1990; Fig. 1), probably all provided sediment sources for younger, late Palaeozoic to Mesozoic mobile belts of New Zealand and northeast Australia (Ireland 19926; Adams & Kelley 1998; Cawood et al 1999, 2002; Pickard et al 2000). The Eastern Province of New Zealand comprises several tectonostratigraphic terranes (Figs 2,3), principally of Permian to Cretaceous, low-grade metasedimentary rocks (Bishop et al 1985). The easternmost and most extensive of these (probably seen as far as the Chatham Islands, Fig. 1), is the Torlesse composite terrane, made up of relatively quartzose turbidite-dominated greywacke successions in the Rakaia and Pahau terranes (MacKinnon 1983), with the remaining, narrower terranes to the west having increasing degrees of volcanic and volcaniclastic detritus, (namely, Waipapa and Caples terranes, mainly acid-intermediate; Dun Mountain-Maitai, Murihiku and Brook Street terranes, mainly intermediate-basic). Significant volcanic edifices are only present in the Brook Street terrane. For the Torlesse composite terrane, petrographical, geochemical and detrital mineral age evidence suggests a continent-derived sediment supply, principally of Permian-Triassic granitoids, into an accretionary prism environment (MacKinnon 1983; Adams & Kelley 1998; Adams et al 1998; Roser & Korsch 1999; Pickard et al. 2000). A similar situation is recognized for Waipapa and Caples terranes, but with

Fig. 1. Pacific margin of Gondwanaland in a Late Triassic reconstruction (after Veevers & Powell 1994; Sutherland 1999) with 30° and 60° meridians and South Pole (SP) regions discussed in text. Dashed and dotted lines are the estimated late Precambrian and late Palaeozoic continental margins, respectively. Delamerian Fold Belt: cross-hatched; Lachlan Fold Belt: vertical ruling; New England Fold Belt: horizontal ruling. Some major granitoid terranes are noted: I-type as i, S-type as s. The New Zealand mobile belt (Eastern Province) terranes are shaded grey. Some specific locations discussed in text are: 1, Shoalwater and Wandilla terranes, Queensland; 2, Gympie terrane, Queensland; 3, Mallakoota, Victoria; 4, Bounty Is.; 5, Chatham Is.; 6, Campbell L; 7, Thurston L, West Antarctica; 8, Thiel Mountains; 9, Alexander L, Antarctic Peninsula; 10, Elephant L, South Shetland Is.; 11, extra-Andean Patagonia; 12, Diego de Almagro L; 13, Madre de Dios L; 14, Chonos Archipelago; 15, Sierras Pampeanas; 16, Puna.

116

C. J. ADAMS ETAL.

Fig. 2. New Zealand tectonostratigraphic terranes, their boundary faults (solid line) and probable extensions (dashed) as follows: b Duller, c Caples (includes Aspiring), p Pahau, r Rakaia, m Drook Street, Murihiku and Dun Mountain-Maitai (combined), t Takaka, w Waipapa (includes some possible South Island correlatives), mtz denotes the Median Tectonic Zone, or Median Batholith. Nd-Sr sampling locations are shown as dots.

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN

117

Fig. 3. Stratigraphic columns of New Zealand Eastern Province terranes, triangular flags at left of columns indicate tectonic intercalations of sediments of pelagic association (hemipelagites, cherts, limestones, tuffs etc.).

additional active volcanic influences (Sporli 1978; Turnbull 19790, b). In contrast, more probable back-arc or forearc environments are suggested for Murihiku and Dun MountainMaitai terranes (Campbell & Coombs 1966; Coombs et al 1976; Ballance & Campbell 1993), with some continent-derived sediment input in the latter. Finally, the Brook Street terrane represents an isolated and dissected, predominantly Permian volcanic island arc (Houghton 1981; Houghton & Landis 1989; Mortimer et al 1999&), but with platform sediments (Landis et al 1999). The relative positions of the Eastern Province terranes suggest that the Torlesse, Waipapa and, possibly, Caples terranes are 'suspect' and have distant sediment sources. Possible origins have been suggested, partly in the Lachlan Fold Belt and its continuations in New Zealand and Antarctica (Cawood et al 1999; 2002) and partly or completely in the New England Fold Belt of northeast Australia (Pickard et al 2000; Cawood et al 2002). The New Zealand Median Batholith continues northwest of New Zealand (Mortimer et al 1997) on to the Lord Howe Rise and Dampier Ridge (Fig. 1) and correlates with extensive Carboniferous-Triassic plutonic (and volcanic) terranes in the New England Fold Belt. Only the Brook Street terrane of New Zealand has a probable Australian correlate in the Gympie terrane of eastern Queensland (Fig.

1). In easternmost Queensland, the Wandilla and Shoalwater terranes are regarded as allochthonous metasedimentary terranes (Leitch et al 2003) and, although they are similar lithologically to Torlesse, Waipapa and Caples terrane rocks, they are probably entirely older (i.e. pre-Carboniferous). Eastern Province and Median Batholith rocks do not extend significantly on to the Campbell Plateau, southeast of South Island, New Zealand (Fig.l). However, Bradshaw et al (1997) regard a broad zone of Permian to Jurassic igneous rocks in eastern and western Marie Byrd Land, West Antarctica (Fig. 1) as an extension of the Median Tectonic Zone (Median Batholith), and some rare exposures of associated metasediments there might also represent Eastern Province rocks.

West Antarctica and southern South America Pre-Jurassic turbiditic metasedimentary rocks in the Antarctic Peninsula (Fig. 1) form generally isolated or semi-continuous outcrops in two main areas of the northern peninsula (the northernmost part and Alexander Island) and are considered to be of mostly Triassic age. Lowmedium-grade metasedimentary rocks of the Scotia metamorphic complex, exposed in the South Shetland and South Orkney islands, may

118

C. J. ADAMS ET AL.

also include equivalent pre-Jurassic elements, but a strong Cretaceous metamorphic overprint, e.g. in Elephant Island (Trouw etal 1990), masks resolution of younger juvenile accreted material. These sequences form part of the Western Domain defined by Vaughan & Storey (2000), considered as a suspect terrane accreted onto the Antarctic Peninsula margin, most probably in Cretaceous times. More isolated occurrences of deformed metasedimentary rocks occur throughout the other microcontinental blocks of West Antarctica (Fig. 1), most notably in the Ellsworth-Whitmore mountains block (Dalziel et al 1987), where they are intruded by Jurassic granites (Millar & Pankhurst 1987). Metasedimentary successions similar to those of the Antarctic Peninsula occur along the western archipelago of southern Chile (Fig. 1), where they are mostly of Permian or Triassic depositional age according to extremely rare fossil remains and detrital zircon studies (Ling et al 1987; Fang et al. 1998; Herve et al. 2003). According to the extended model of Vaughan & Storey (2000), these also form part of the marginal suspect terrane ('Western Domain'), but an exotic origin has only been argued in detail for the Madre de Dios area, which contains Permian limestones supposedly formed at lower palaeogeographical latitudes (Forsythe & Mpodozis 1979; Ling et al. 1987). Further inland, the Eastern Andes Metamorphic Complex (Herve 1993) consists of metaturbidites that are of undisputed Gondwana provenance (Herve et al. 2003), as are scattered Palaeozoic metasedimentary rocks from the extra-Andean regions of Patagonia (Pankhurst et al. 20010; 2003; Fig. 1). Previous studies and present work The New Zealand data presented here expand the Nd-isotope study of Eastern Province sedimentary rocks by Frost & Coombs (1989) and a few Western province analyses, included with primarily granitoid/volcanic data, of Pickett & Wasserburg (1989), Waight et al (1998) and Wombacher & Mirnker (2000). Frost & Coombs (1989) sampled Eastern Province terranes only in South Island, New Zealand, from (1) lower metamorphic grade rocks (zeolite to prehnitepumpellyite facies) mostly as sandstone, but also as mudstone lithologies, and (2) higher-grade (greenschist facies) equivalents, in the Haast Schist (Mortimer 19930) occurring as mica- and quartzo-feldspathic schists. Analyses of associated sandstone-mudstone horizons in Torlesse turbidites revealed probable provenance differences, suggesting a higher proportion of

Nd originating from younger sources in the coarser grain sizes - similar observations in modern turbidites have been related to sedimentary sorting processes (McLennan et al 1989). More generally, however, the uniform crustal residence ages, 1.0-1.4 Ga, irrespective of stratigraphic age, lithology or geographical position, were taken to indicate sediments well mixed during transport and dispersal in fluviatile and marine environments. However, they also highlighted the variability of Sm-Nd data within and amongst Caples, Dun MountainMaitai and Murihiku terrane samples, suggesting a local provenance of intrinsically more variable character, at an active continental margin receiving substantial juvenile (volcanic) input. This paper assembles Nd data for >350 samples (Tables 1 and 2; full data are available online at http://www.geolsoc.org.uk/SUP18225. A hardcopy can be obtained from the Society Library.) from the former Gondwanaland margin in New Zealand, southeast Australia, Antarctica and South America (more than half being new Nd-Sr analyses), in order (1) to place general isotopic constraints on the provenance of sediments in these areas, and (2) to investigate the status of the various terrane models that have been advanced for the margin. The Eastern and Western Province terranes of New Zealand were sampled in North Island and South Island, the adjacent Campbell Plateau and Chatham Rise, concentrating almost exclusively on (meta-)siltstone lithologies. These were taken from sandstone-siltstone-mudstone sample sets of graded beds in those turbiditic units used for Rb-Sr geochronological study (Adams & Graham 1996, 1997; Adams et al 1999, 2002; Adams & Maas 20040, b) and which are mostly from biostratigraphically wellconstrained horizons at low metamorphic grade (zeolite- to pumpellyite-actinolite mineral facies). Siltstones were used to minimize the possible isotopic effects of mineral sorting in sandstones and of (admittedly unlikely) inclusion of exotic pelagites and/or far-travelled airborne (?volcanic) dust in mudstone lithologies. The Rb-Sr isochron studies on these metasediments, including many localities sampled in the present work (Adams & Graham 1996, 1997; Adams et al 1999, 2002; Adams & Maas 20040, b) indicate a high degree of Srisotopic homogenization between different size fractions during metamorphism and, since the centroids of such isochron data are close to individual siltstone analyses, then the latter represent a good bulk sediment composition. Nd-Sr isotope data assembled for the

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN

119

Table 1. Nd- and Sr-isotope data for New Zealand sedimentary rocks and correlatives in West Antarctica and eastern Australia Line no.

Stratigraphic unit*

Location

WESTERN PROVINCE OF NEW ZEALAND AND CORRELATIVES BULLER TERRANE, NEW ZEALAND I Greenland Group Webb Stream 2 Greenland Group 14 Mile Bluff 3 Greenland Group Jackson Bay 4 Greenland Group (fossil locality) Waitahu River 5 Greenland Group Lake Kangaroo 6 Greenland Group Lake Kaniere 7 Greenland Group Lake Kaniere 8 Greenland Group (uncertain) Bounty Is. 9 Complex Point Formation Campbell Is. MARIE BYRD LAND, WEST ANTARCTICA Swanson Formation 10 Mackay Mrs., MBL 11 Swanson Formation La Gorce Pk., MBL LACHLANFOLD BELT, SOUTHEAST AUSTRALIA 12 Bendoc Group (Mallacoota Beds) Mallacoota, VIC 13 Castlemaine Supergroup Bacchus Marsh, VIC 14 Knowsley East Shale Kilmore, VIC 15 Bryo Gully Shale, black shale Kilmore, VIC 16 Belinda Shale Wallan, VIC 17 Deep Ck Siltstone Wallan, VIC 18 Kilmore Siltstone Wallan, VIC 19 Glenburn, VIC Humevale Formation 20 Lazarini Siltstone Howqua River, VIC 21 Mt Easton Shale, black shale Howqua River, VIC 22 Walhalla Group Howqua, VIC 23 Pinnak Sandstone, myl. phyllite Tallangatta Valley, VIC 24 Pinnak Sandstone, phyllite Club Terrace, VIC 25 Cobannah Group, siltstone near Briagolong, VIC 26 Cobbs Spur Andesite, volcaniclastic Licola, VIC 27 Wilsons Creek Shale Eildon, VIC 28 Murderer's Hill Siltstone, phyllite Heyfield, VIC 29 Castlemaine Supergroup, phyllite Bendigo, VIC 30 Bindaree Fm, lacustrine black shale Howqua River, VIC 31 St Arnaud Beds, phyllite Stawell, VIC 32 Knowsley East Shale Kilmore, VIC 33 Lano Gully Sandstone Lancefield, VIC 34 Darraweit Guim Mudstone Darraweit,VIC 35 Castlemaine Supergroup Bendigo,VIC 36 Yalmy Group, hornfels in granite Snowy River, VIC 37 Yalmy Group, hornfels in granite Snowy River, VIC 38 State Circle Shale Canberra, ACT 39 State Circle Shale Canberra, ACT 40 CSA Siltstone Cobar, NSW 41 CSA Siltstone Cobar, NSW 42 Girilambone Beds east of Cobar, NSW 43 Reefton Group Reefton TAKAKA TERRANE, NEW ZEALAND 44 Junction Formation Cobb Lake Baton Formation Baton River 45 46 47 (Pikikiruna Schist) Takaka Hill 48 Parapara Group Parapara Peak 49 Lake Peel Formation Northwest Nelson 50 Tasman Formation Northwest Nelson Tasman Formation 51 Northwest Nelson 52 Heath Creek Beds Northwest Nelson 53 Junction Formation Northwest Nelson 54 Junction Formation Northwest Nelson 55 Mount Benson Sandstone Northwest Nelson 56 Northwest Nelson Mount Benson Sandstone 57 Northwest Nelson Ruby Saddle Formation 58 Ruby Saddle Formation Northwest Nelson 59 Ruby Saddle Formation Northwest Nelson

Data ref.

Aget (Ma)

^Ndi

TDM# (Ma)

Sri

1 1 1 1 3 8 8 1 1

440 440 440 440 440 500 500 414 440

-7.5 -7.7 -10.0 -9.8 -7.2 -7.5 -10.2 -8.3 -9.9

1747 1757 1914 1906 1725 1788 1968 1786 1908

0.71754 0.71188 0.71865 0.72563 0.72102 0.71578 0.71656 0.70995 0.73022

1 1

440 440

-8.3 -10.6

1802 1954

0.71609 0.72141

1 2a 2a 2a 2a 2a 2a 2a 2a 2a 2a 2a 2a 2a 2a 2a 2a 2a 2a 2a 4 4 4 4 5 5 4 6 2b 2b 2b 1

450 475 508 485 437 435 420 406 435 455 390 470 470 430 500 395 415 475 375 495 508 490 437 470 440 440 435 435 410 410 470 410

-9.4 -8.7 -10.1 -8.4 -10.5 -9.5 -8.4 -9.9 -11.2 -8.5 -9.5 -8.4 -5.1 -9.6 -1.8 -10.2 -10.5 -8.9 -8.5 -11.6 -9.6 -8.4 -10.1 -10.7 -9.4 -12.4 -10.5 -9.5 -10.6 -11.9 -9.6 -10.9

1881 1850 1970 1841 1950 1881 1797 1884 1993 1824 1848 1831 1592 1882 1361 1900 1930 1865 1770 2056 1937 1845 1924 1980 1878 2076 1949 1879 1935 2024 1909 1953

0.71475 0.71444 0.71680 0.71457 0.71475 0.71919 0.72231 0.72540 0.71776 0.71637 0.72397 0.69372 0.65102 0.71251 0.70941 0.73054 0.72764 0.71691 0.72358 0.72231 0.71648 0.71438 0.71421 0.70961 0.71028 0.71400 0.71406 0.72242 0.72208 0.71900 0.72444 0.71097

1 1 1 3 1 9 9 9 9 9 9 9 9 9 9 9

500 405 405 440 260 500 505 505 515 515 515 510 510 505 505 505

-10.5 -9.2 -11.0 -10.8

1987 1842 1960 1971

0.70966 0.70475 0.67695 0.71076 0.70046 0.70788 0.70907 0.70852 0.70640 0.70962 0.71452 0.70063 0.68751 0.71331 0.69274 0.70019

2.7 2.2 -3.1 -2.7 -7.9 -9.5 -11.0 -3.0 -3.4 -4.8 -5.0 -4.1

807 1027 1470 1438 1828 1933 2034 1466 1495 1597 1611 1548

C. J. ADAMS ETAL.

120

Table 1. (continued} Line no.

Stratigraphic unit*

Location

TAKAKA TERRANE, NEW ZEALAND (continued) Northwest Nelson 60 Ruby Saddle Formation Northwest Nelson 61 Junction Formation Northwest Nelson 62 Anatoki Formation

Data ref.

Aget (Ma)

9 9 9

Ndi

TDM* (Ma)

Sri

505 515 505

-4.4 -9.5 0.2

1569 1936 1203

0.66860 0.70978 0.70633

E

EASTERN PROVINCE OF NEW ZEALAND: 'VOLCANISTIC TERRANES BROOK STREET TERRANE 63 Greenhills Group 64 Divide Formation 65 Consolation Formation 66 Takitimu Group 67 Takitimu Group 68 Takitimu Group 69 Takitimu Group 70 Takitimu Group (Consolation Formation) 71 Greenhills Group 72 Grampian Formation

Omaui Rd. Consolation Peak Melita Peak Wairaki River Elbow Creek Longwood Range Wairaki River Hollyford River Howell Point Flaxmoor

7 7 7 1 1 1 7 1 1 1

270 270 270 275 260 255 260 255 250 250

7.5 8.4 5.4 5.4 7.5 8.4 7.6 6.1 7.5 -0.7

335 230 562 559 334 224 318 470 315 1098

MURIHIKU TERRANE 73 Kuriwao Group 74 Murihiku Supergroup 75 Murihiku Supergroup 76 Murihiku Supergroup 77 Murihiku Supergroup 78 ?North Range Group 79 North Peak Fmtn 80 ?North Range Group 81 Taringatura Group 82 Taringatura Group 83 Rengarenga Group 84 Murihiku Supergroup 85 Huriwai Group 86 Purakauiti Formation

Mataura Is. Dipton Ohai Wairoa River Roaring Bay Oreti River Hokonui Hills Kaka Point Nugget Point Roaring Bay Kiritehere Beach Warnock Road Port Waikato Caberfeidh

1 1 1 1 1 7 7 7 7 7 1 1 1 7

255 225 240 215 215 245 245 240 235 215

592 899 823 733 959 476 477 908 935 1070 843 854 561 1015

0.70436 0.70475 0.70417 0.70480 0.70478

170 150 175

5.0 1.4 2.4 3.2 0.7 6.0 6.0 1.5 1.1 -0.6 1.9 1.5 4.5 -0.3

DUNMOUNTAIN-MAITAI TERRANE 87 Tramway Formation 88 Tramway Formation 89 Maitai Group (Little Ben) 90 Tramway Formation 91 Greville Formation 92 Greville Formation 93 ?Greville Formation 94 ? Greville Formation 95 Waiua Formation 96 Waiua Formation 97 Waiua Formation 98 Stephens Subgroup 99 Countess Formation

Roding River Key Summit Hollyford River, Southland Gyzeh Trig Roding River Oreti River East Eglinton River East Eglinton River Roding River Roding River, Nelson Mararoa River Wairoa River Countess Range

1 1 1 7 1 7 7 7 1 1 7 1 1

250 250 245 260 230 255 255 260 230 240 250 230 230

-0.2 -0.7 2.4 1.7 0.5 -2.7 2.9 6.5 3.9 3.9 -0.4 1.7 1.7

1057 1098 824 901 983 1263 789 433 676 681 1074 881 883

0.70512 0.70446 0.70149

CAPLES TERRANE 100 Omahuta Unit 101 Momus Sandstone 102 Kays Creek Formation 103 Bold Peak Formation 104 Caples Group 105 Caples Group 106 Tuapeka Group 107 Tuapeka Group 108 Caples Group 109 Caples Group 110 Caples Group (uncertain) 111 Caples Group (uncertain)

Paiokatutu Str. Mavora Lake Mavora Lakes Von River North Von River Stoney Creek Barnego Barnego Upper Mararoa R. Kays Creek Chrystalls Beach Quoin Pt.

1 1 1 1 7 1 7 7 7 7 1 1

190 220 250 210 245 220 235 235 235 245 220 220

4.0 2.0 -1.6 2.1 0.7 1.8 5.5 0.1 0.9 6.2 -1.5 0.0

639 845 1169 825 976 861 526 1023 949 459 1141 1016

0.70418 0.70401 0.70481 0.70348

ASPIRING TERRANE 112 (Haast Schist) 113 Haast Schist 114 Haast Schist

Matukituki Valley Nevis Bluff Nevis Bluff

1 7 7

250 240 240

-1.6 6.2 -1.1

1172 453 1127

0.70428

21Q

0.70299 0.70312 0.70279 0.70285 0.70322 0.70521

0.70472 0.70519 0.70359

0.70392

0.70422 0.70379 0.70390 0.70712

0.70393

0.70465 0.70490

121

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN Table 1. (continued) Line no.

Location

Stratigraphic unit*

Data ref.

Aget (Ma)

WAIPAPA TERRANE Waipapa Group 115 Waipapa Group 116 117 Waipapa Group 118 Waipapa Group Waipapa Group (uncertain) 119 (Haast Schist) 120 (Haast Schist) 121 122 (Haast Schist)

Stephenson Island Rangihoua Bay Puketona Benneydale Kapiti Island Blumine Island Fighting Bay Port Underwood

1 1 1 1 1 1 1 1

220 210 22Q 155 210 170 160 160

RAKAIA TERRANE 123 Torlesse Supergroup 124 Torlesse Supergroup 125 Torlesse Supergroup 126 Torlesse Supergroup 127 Torlesse Supergroup 128 Torlesse Supergroup 129 Torlesse Supergroup 130 Torlesse Supergroup 131 Torlesse Supergroup 132 Torlesse Supergroup 133 Torlesse Supergroup 134 Torlesse Supergroup 135 Torlesse Supergroup 136 Torlesse Supergroup 137 Torlesse Supergroup 138 (Haast Schist) 139 Haast Schist 140 Haast Schist 141 Haast Schist 142 Haast Schist 143 (Chatham Island Schist)

Selwyn River Kakahu River Pareora Gorge Hakataramea Balmacaan Str Benmore Dam Benmore Dam Pudding Hill Str. Awakino Skifield Pukerua Bay Ngauranga Gorge Ohau Skifield Arthurs Pass Otaki Gorge Red Rocks Sheepwash Creek Swinburn Home Hills Rd Pennyweight Hill Lake Hawea Matarakau Point

1 1 7 1 1 7 7 1 1 1 1 1 7 1 1 1 7 7 7 7 1

250 250 250 250 235 235 235 230 230 230 230 230

Waipawa River, Wairarapa Awatere River, Marlborough Clarence Valley Clarence Valley Hurupi Str. Wharekauhau

1 1 7 7 1 1

150 95 132 132

Byfield, QLD Cape Keppel, QLD

1 1

PAHAU TERRANE 144 Torlesse Supergroup 145 Torlesse Supergroup 146 Torlesse Supergroup 147 Torlesse Supergroup 148 Torlesse Supergroup 149 Torlesse Supergroup NEW ENGLAND OROGEN, AUSTRALIA 150 (Shoalwater Terrane) 151 Wandilla Group

222 220 220 240 240 240 240 240 220

no 11Q 300 300

TDM# (Ma)

Sri

3.9 1.2

776 658 916

-1.1 -2.9 -2.6

1068 1269 1194

0.70394 0.70431 0.70397 0.70567 0.70625 0.70530 0.70585 0.70605

£

Ndi

2.8

0.7

920

-2.5

1183

-3.2 -2.8 -3.4 -1.9 -3.0 -2.2 -4.0 -3.8 -3.0 -4.3 -4.5 -4.1 -4.6 -4.4 -4.7 -2.4 -3.7 -3.3 -3.3 -2.0 -4.1

1296 1271 1315 1197 1274 1207 1349 1335 1268 1372 1383 1353 1384 1371 1394 1230 1334 1303 1299 1198 1348

-3.8 -1.2 -1.1 -3.3 -3.9 -4.1

1278 1036 1047 1229 1254 1276

0.70761 0.70628

-8.4

1714 1043

0.70948 0.70547

0.4

0.70731 0.70613 0.70595 0.70785

0.70741 0.70685 0.70777 0.70693 0.70727 0.70752 0.70810 0.70608

0.70698

0.70848 0.70839

Data sources (recalculated if necessary): 1, this work; 2, this work (Uni Melbourne: 2a, O'Halloran, 1996 PhD thesis; 2b, R. Maas, unpublished); 3, Pickett & Wasserberg (1989); 4, Turner et al. (1993); 5, Maas et al. (1997); 6, McCulloch & Chappell (1982); 7, Frost & Coombs (1989); 8, Waight et al. (1998); 9, Wonbacher & Miinker (2000). *Stratigraphic unit: brackets indicate non-stratigraphic identifier, in the absence of a formal name. f Age (Ma); Stratigraphic age based on diagnostic fossils, otherwise (underlined) estimated age based upon minimum metamorphic ages ages from K-Ar and Rb-Sr data and maximum ages from U-Pb detrital zircon age data. Sr data were normalized to 86Sr/88Sr = 0.1194, and 87Sr/86Sr for NBS987 = 0.710235, where possible. Nd data were normalized to 146Nd/144Nd = 0.7219, and 143Nd/144Nd for La Jolla = 0.511864, where possible. TDM* is two-stage crustal model according to DePaolo et al. (1991).

Lachlan Fold Belt are, in large part, from O'Halloran (1996, and others are used with permission); the remainder are from a variety of sources (see Table 1). The Lachlan Fold Belt sediments were selected to cover the stratigraphic range from Cambrian to Devonian, marine to terrestrial facies and the three subprovinces of the fold belt (e.g. Foster & Gray 2000). In these cases, lithologies include both siltstone and mudstone. With the exception of two samples of Yalmy Group, collected as

hornfels xenoliths in a granite, and one sample of mylonitic Pinnak Sandstone, all samples are either unmetamorphosed, or are at lower greenschist facies (e.g. Offler et al 1998; Foster & Gray 2000), To compare with the New Zealand and Australian datasets of Table 1, equivalent data for proto-Pacific margin metasedimentary rocks of West Antarctica and southern South America are assembled in Table 2, although in this case the sample selection was often of a more

C.J.ADAMSCTAL.

122

Table 2. Neodymium and strontium isotope data for west Antarctic and South American sedimentary rocks Line no.

Stratigraphic unit/lithology*

ALEXANDER ISLAND (West Antarctica) 201 202 203 204 205 206 207 208 LeMay Fm 209 LeMay Fm 210 LeMay Fm 211 LeMay Fm 212 LeMay Fm 213 LeMay Fm 214 LeMay Fm 215 LeMay Fm 216 LeMay Fm 217 LeMay Fm 218 LeMay Fm 219 LeMay Fm 220 LeMay Fm ANTARCTIC PENINSULA: Trinity Peninsula 221 Trinity Peninsula Group 222 Greywacke Shale Group 223 Miers Bluff Fm 224 Miers Bluff Fm 225 Trinity Peninsula Group xenolith 226 Trinity Peninsula Group 227 Trinity Peninsula Group 228 Trinity Peninsula Group 229 Trinity Peninsula Group SOUTH SHETLAND ISLANDS 230 Gt-mp 231 Gt-mp 232 mp 233 mp 234 mp 235 phyllite 236 phyllite 237 phyllite 238 mp 239 mp 240 mp SOUTHERN CHILE (Western Domain) 241 schist 242 Chonos Metamorphic Complex 243 Chonos Metamorphic Complex 244 Chonos Metamorphic Complex 245 Chonos Metamorphic Complex 246 schist 247 Chonos Metamorphic Complex 248 Chonos Metamorphic Complex 249 Chonos Metamorphic Complex 250 Chonos Metamorphic Complex 251 phyllite 252 phyllite 253 Western Series

Ndi

TDM* (Ma)

Sri

150 150 150 189 189 189 120 128 128 128 128 128 128 128 128 128 128 234 234 234

-3.2 -4.5 -5.8 -5.0 -4.9 -4.6 -6.7 -2.7 -2.4 -2.9 -3.1 -2.0 -2.6 -2.5 -2.9 -3.0 -2.6 -5.2 -5.9 -3.9

1234 1330 1430 1392 1388 1361 1476 1175 1153 1189 1206 1119 1170 1165 1192 1197 1168 1438 1489 1339

0.7114 0.7110 0.7092 0.7088 0.7090 0.7092

1 1 1 1 1 1 1 1 1

240 240 240 240 240 220 220 220 220

-4.2 -3.7 -5.0 -5.8 -4.3 -3.7 -3.3 -4.3 -4.8

1367 1333 1427 1487 1380 1320 1286 1360 1403

0.7089 0.7078 0.7089 0.7088

Stinker Pt., Elephant I. Stinker Pt., Elephant I. Stinker Pt., Elephant I. Stinker Pt., Elephant I. Pta. Algas, Elephant I. C.Lindsay, Elephant I. C. Yelcho, Elephant I. Hut Pt.,Elephant I. Clarence I. Clarence I. Elephant I.

2 2 2 2 2 2 2 1 2 1 1

100 100 100 100 100 100 100 100 100 100 100

4.5 4.3 3.5 -5.7 1.6 3.7 -6.7 3.0 -6.2 -4.8 -5.7

523 541 619 1383 796 598 1461 670 1422 1319 1387

0.7047 0.7047 0.7048 0.7098 0.7061 0.7048 0.7046 0.7049 0.7090 0.7092 0.7070

Huinay I. Leniu I. Lemu I Patranca I Patranca F. Sangra I Ximena I Teresa I Teresa F. Puelma I Diego Ramirez I Diego Ramirez Piren

3 3 3 3 3 3 3 3 3 3 3 3 3

300 224 224 224 224 224 224 224 224 224 175 175 300

-6.5 -3.7 -6.2 -5.5 -5.8 -7.3 -5.1 -4.6 -6.9 -5.9 -5.4 4.2 -5.8

1583 1323 1505 1457 1475 1583 1429 1392 1560 1484 1412 604 1531

0.7114 0.7077 0.7105 0.7077 0.7066 0.7086 0.7195 0.7135 0.7066 0.7114 0.7074 0.7054 0.7046

Location

Data ref.

Age (Ma)

Arenite Ridge Central Douglas Range South Douglas Range N of Lemay Range N of Lemay Range N of Lemay Range Charcot Island LeMay Range LeMay Range LeMay Range LeMay Range LeMay Range LeMay Range LeMay Range LeMay Range LeMay Range LeMay Range Walton Mtns. Walton Mtns. Walton Mtns.

1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1

Group and Equivalents Lahille I Gibbs I. Livingston I. Livingston I. Trinity Peninsula Hope Bay Hope Bay Hope Bay Hope Bay

£

0.7070 0.7069 0.7069 0.7070 0.7069 0.7069 0.7070 0.7070 0.7070 0.7071 0.7071 0.7081

0.7078 0.7071 0.7099 0.7109

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN

123

Table 2. (continued) Line no.

Stratigraphic unit/lithology*

Location

SOUTHERN CHILE (Eastern Andes Metamorphic Complex) 254 Lago General Carrera unit 255 Lago General Carrera unit 256 Lago General Carrera unit 257 Lago General Carrera unit 258 Cochrane unit 259 Cochrane unit 260 Cochrane unit 261 Cochrane unit 262 Cochrane unit 263 Cochrane unit 264 Cochrane unit 265 Cochrane unit 266 Cochrane unit INBOARD WEST ANTARCTICA (Eastern Domain) 267 Cochrane unit 268 mp Mt Wrather 269 mp Mt Woollard 270 mp Hart Hills 271 mp Hart Hills 272 mp Mt Moore 273 mp Martin Hills 274 mp Stewart Hills 275 mp Stewart Hills 276 mp Morland Nunataks 277 mp Northern Nunataks 278 Polarstar Fm Ellsworth Mtns 279 Polarstar Fm Ellsworth Mtns 280 Patuxent Fm Patuxent 281 mudstone Milan Rocks 282 mudstone Milan Rocks 283 schist Mt Petras

Sri

Data ref.

Age (Ma)

^Ndi

TDM# (Ma)

3 3 3 4 4 4 4 4 4 4 4 4 4

320 320 320 350 350 350 350 350 350 350 350 350 350

-3.5 -3.9 -7.1 -2.2 -5.5 -4.8 -6.0 -4.6 -5.4 -6.0 -6.0 -6.5 -6.2

1372 0.7111 1398 0.7092 1639 0.7090 1289 1540 1488 1580 1472 1531 1576 1580 1613 1595

4 1 1 1 1 1 1 1 1

1 1 1

350 500 500 500 500 467 220 420 395 370 370 250 250 440 330 330 300

-5.1 -3.4 -4.1 -4.5 -6.8 -7.3 -8.2 -5.7 -6.0 -7.5 -7.8 -4.7 -7.2 -6.0 -10.0 -9.9 -3.7

1513 1492 1542 1570 1741 1751 1645 1607 1607 1698 1721 1417 1599 1642 1845 1834 1370

0.7079 0.7201 0.7093 0.7340 0.7069 0.7178 0.7189 0.7090 0.7094 0.7148 0.7208 0.7233 0.7078

0.7142 0.7144

SIERRAS PAMPEANAS (NW Argentina) 284 Phyllite 285 Phyllite 286 Schist 287 Schist 288 Schist

Malanzan Malanzan El Pilon El Pilon Los Tuneles

5 5 5 5 5

525 525 525 525 525

-7.0 -7.0 -7.7 -3.1 -6.9

1771 1769 1820 1483 1762

0.7107 0.7119 0.7158 0.7132 0.7125

ARGENTINE PATAGONIA 289 sst 290 metasst 291 metasst 292 gneiss 293 schist 294 metasst 295 metasst 296 metasst 297 Schist 298 Schist 299 Schist

Sierra Grande Nahuel Niyeu El Jaguelito Mina Gonzalito Esquel Catreleo Collon Cura Cushamen Dos Hermanos Dos Hermanos Dos Hermanos

6 6 6 6 6 6 6 6 7 7 7

500 500 500 460 460 460 440 460 580 580 580

-5.6 -5.4 -4.0 -4.7 -0.7 -6.1 -5.5 -5.9 -2.6 -2.3 -3.0

1654 1636 1534 1575 1135 1661 1600 1599 1487 1461 1516

0.6772 0.7128 0.7123 0.7105 0.7051 0.7130 0.7109 0.7121 0.7110 0.7096 0.7128

8 8 8 8 8

489 489 492 488 488

-10.0 -10.3 -10.5 -11.6 -10.9

1948 1966 1984 2054 2010

0.7123 0.7118 0.7129 0.7120 0.7094

PUNA (NW Argentina) 300 Tolar Chico Fm 301 Tolar Chico Fm 302 Tolar Chico Fm 303 Tolar Chico Fm 304 Tolar Chico Fm

C. J. ADAMS ET AL.

124 Table 2. (continued) Line no. 5305 306 307 308 309 310 311 312 313 314 315 316 317 318 319 320 321 322 323 324 325 326 327 328

Stratigraphic unit/lithology*

Tolar Chico Fm Tolar Chico Fm Tolillar Fm Tolillar Fm Tolillar Fm Tolillar Fm Tolillar Fm Tolillar Fm (ash) Tolillar Fm (ash) Tolillar Fm (ash) Tolillar Fm (ash) Tolillar Fm (ash) Tolillar Fm (ash) Tolillar Fm (ash) Diablo Fm Diablo Fm Diablo Fm Diablo Fm Diablo Fm Coquena/Falda Cienaga fms Coquena/Falda Cienaga fms Coquena/Falda Cienaga fms Coquena/Falda Cienaga fms Coquena/Falda Cienaga fms

Location

Data ref.

Age (Ma)

8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8 8

488 492 485 482 480 480 480 485 485 478 485 484 478 475 477 476 476 476 475 473 468 468 473 468

Ndi

TDM# (Ma)

Sri

-9.8 -9.4 -5.5 -5.4 -5.2 -5.5 -5.7 -6.2 -5.9 -5.2 -5.0 -6.9 -3.1 -1.4 -2.8 -1.4 -2.2 -2.1 -2.2 -6.9 -6.5 -6.7 -5.4 -5.7

1938 1913 1632 1625 1612 1628 1644 1686 1665 1610 1600 1735 1447 1317 1425 1319 1380 1373 1376 1723 1695 1706 1621 1639

0.7102 0.6985 0.7062 0.7081 0.7075 0.7078 0.7071 0.7087 0.7087 0.7087 0.7081 0.7057 0.7036 0.7088 0.7085 0.7078 0.7078 0.7079 0.7081 0.7105 0.7111 0.7114 0.7088 0.7096

E

*Where specified. Data sources (recalculated if necessary): 1, this work (NIGL); 2, Trouw et al. (1990); 3, Herve & Pankhurst, unpublished; 4, Augustsson & Bahlburg (2003); 5, Rapela et al. (1998); 6, Rapela & Pankhurst, unpublished; 7, Pankhurst et al (2003); 8, Zimmerman & Bahlburg (2003). Sr data were normalized to 86Sr/88Sr = 0.1194, and 87Sr/86Sr for NBS987 = 0.710235, where possible. Nd data were normalized to 146Nd/144Nd = 0.7219, and 143Nd/144Nd for La Jolla = 0.511864, where possible. TDM* is two-stage crustal model according to DePaolo et al. (1991).

random nature, often incidental to other field collecting, and without grain-size control (most of the samples are metapelitic). There is no published systematic Sm-Nd study of West Antarctic metasedimentary rocks and most of the results presented here are hitherto unpublished data from samples collected by several British Antarctic Survey field parties. Data for the Scotia metamorphic complex have been presented by Trouw et al (1990), from which the results for metapelitic samples have been selected. In southern South America, data have been published for recent studies of Nd systematics in the Ordovician sequences of the Puna, NW Argentina (Zimmermann & Bahlburg (2003) and with which are now included unpublished Rb-Sr analyses of the same samples, courtesy of U. Zimmermann). Five analyses of Cambro-Ordovician rocks in the Sierras Pampeanas have been taken from Rapela et al (1998) and Pankhurst et al (19980). Further south, in Patagonia, Augustsson & Bahlburg (2003) have published data for the eastern

Andes metamorphic complex along the ChileArgentina border. Many of the remaining data included here for the archipelago of Chilean Patagonia and the extra-Andean Argentine Patagonia are derived from collaborative projects between RJP and F. Herve (Chile) and C. W. Rapela (Argentina), respectively. These results were summarized initially by Pankhurst et al (1994), but using a different model age calculation and without presentation of the measured data.

Analytical methods The majority of the new data were obtained by conventional chemical separation and thermal ionization mass-spectrometry at NERC Isotope Geosciences Laboratory (NIGL), UK. Isotope dilution was used for Sm and Nd with a mixed 149 Sm-150Nd spike and X-ray fluorescence for Rb and Sr concentrations. New Zealand samples were analysed using a Finnigan MAT 262 mass spectrometer, during 2001-2002, but

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN

those from West Antarctica and South America over a much longer period of time, including some older data determined on a VG 354 instrument. Internal precision of 87Sr/86Sr and 143 Nd/144Nd analyses was routinely better than 10 ppm (s.e), and Nd data were adjusted to a common basis using Johnson Matthey Nd as an internal standard, correcting to a long-term average 143Nd/144Nd ratio of 0.511123. The adjusted values obtained for La Jolla and basalt BHVO-1 are 0.511864 ± 0.000003 and 0.512979 ± 0.000004, respectively (errors at 95% confidence limits). Measured 87Sr/86Sr ratios were adjusted to a value of 0.710235 for strontium standard NBS 987. Other new analyses were carried out at La Trobe University, Melbourne, as described in Maas et al (1997) on 50 mg of sample powder spiked with mixed 149Sm-150Nd and 85Rb-84Sr tracers. As at NIGL, Sm and Nd were purified using the method of Richard et al. (1976). Total chemistry blanks were <200 pg for all elements. Sr and Nd isotopic analyses were made in static multicollection mode on a Finnigan MAT262 mass-spectrometer. Internal precisions for Sr and Nd isotope ratios were similar to those quoted above; external (2a) precisions based on replicate analyses of rock standards and sediment samples are estimated as 0.01% (87Sr/86Sr), 0.004% (143Nd/144Nd), 0.5% (87Rb/86Sr) and 0.2% (147Sm/144Nd). The final results were adjusted to the same standard values as listed above. The 87Sr/86Sr ratios calculated at the time of sediment deposition are designated by the abbreviation Sri. Isotopic abundances used in calculation of present-day and initial E^d values at the time of sediment deposition (abbreviated here to £Ndi) and Nd model ages are as follows: 147Sm/144Nd (CHUR) - 0.1966, 143Nd/144Nd (CHUR) 0.512638, 147Sm/144Nd (DM) = 0.222, i43Nd/i44Nd (DM) = 0.513114, 87Rb/86Sr (Bulk Earth) - 0.0839,87Sr/86Sr (Bulk Earth) - 0.7045. Since the metasediments are thought to be of continental provenance, Nd model ages were calculated according to De Paolo et al. (1991), whereby derivation of the bulk sedimentary material at the time of deposition is from a notional reservoir with time-varying continental Sm/Nd, and the age calculated is for extraction of the parent reservoir material from depleted mantle. Depositional ages were estimated, where possible, from palaeontological constraints using the geological time-scale of Young & Laurie (1998), otherwise from published U-Pb detrital zircon (i.e. maximum) ages and Rb-Sr and K-Ar metamorphic (i.e. minimum) ages. All decay constants used are those recommended by Steiger & Jager (1977).

125

Results from New Zealand Western Province terranes and their correlatives The Buller and Takaka terrane data (Table 1), representing the early Palaeozoic Gondwanaland margin, show their evolved character clearly, with ENdi mostly in the range -12.4 to -7.2, compared to the less evolved, more mantle-like data of seven Eastern Province terranes (-4.7 to +8.4) of the New Zealand Permian-Cretaceous mobile belt. The Buller terrane samples are all from Greenland Group localities [1-7], which are assigned a minimum, Late Ordovician (440 Ma) stratigraphic age on the basis of K-Ar and Rb-Sr metamorphic ages (Adams et al. 1975), although a single fossil locality near Reef ton is Early Ordovician. (NB: throughout the following text, numbers in square brackets denote data line numbers in Tables 1 and 2.) Also included are data from unfossiliferous, but probable Ordovician correlatives on the Campbell Plateau, namely, quartzose semischists [8] at seafloor outcrops near Bounty Island (Cullen 1975) and quartzo-feldspathic schists [9] on Campbell Island. In Marie Byrd Land, West Antarctica, other correlatives are unfossiliferous lower greenschist facies metasediments [10, 11] of the Swanson Formation. K-Ar and Rb-Sr ages of these rocks (Adams 1986; Adams et al. 1979), provide the minimum stratigraphic age estimates in Table 1. The Bendoc Group sample [12], at Mallacoota, easternmost Victoria, is from a locality where associated cherts indicate a Late Ordovician age (Vandenberg et al. 2000). In the Sri-ENdi diagram (Fig. 4), Buller terrane data show a well-defined field with a large range in Sri and fairly restricted range of eNdi values. Setting aside a Devonian sample ([13], see below), they define a compact envelope, Sri 0.710-0.730, 8Ndi -10.6 to -7.2), but with Bounty and Campbell Island data falling closer to Buller, rather than Takaka, terrane compositions (below). The Lachlan Fold Belt sample data (eNdi -12.4 to -8.4; Sri 0.709-0.731) occupy a similar field to Buller terrane samples (Fig. 4), but with e Ndi values below -11. Excluded from this set are two clearly anomalous samples: [27], with £ Ndi = -1.8, is a volcaniclastic sediment from the Cobbs Spur Andesite, part of the Cambrian mafic-intermediate volcanic basement of much of the orogen. The other [24], with eNdi = -5.1, is a siliceous shale from the Early-Middle Ordovician Pinnak Sandstone in the Mallacoota Zone. The reason for its high ENdi is unclear. Palaeocurrent data and sandstone petrology exclude major provenance variations in the

126

CJ. ADAMS ETAL.

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN

region south of the Ordovician Macquarie Arc, before the Late Ordovician (Fergusson & Tye 1999). This is consistent with the isotopic uniformity of all other Ordovician samples (ENdi -12.4 to -8.4) and confirms petrographic (Fergusson & Tye 1999) and Sr isotopic (Gray & Webb 1995) evidence suggesting that contributions from Cambrian and Ordovician volcanic sources have been minimal. In addition, the Rb-Sr isotopic systems in both Pinnak Sandstone samples [24, 25] must be disturbed anomalously as they have impossibly low Sri values, i.e. <0.699. It has been argued that Cambro-Ordovician turbidites in the Lachlan Fold Belt are derived from the Delamerian and Ross orogens (e.g. Cas 1983); more recently, this has been confirmed by studies of petrology (Pound et al 1994; Fergusson & Tye 1999), Sr isotopes (Gray & Webb 1995), 40Ar/39Ar geochronology (Turner et al 1996) and U-Pb zircon inheritance (Williams et al 1994; Ireland et al 1998; Veevers 2000; Fergusson & Fanning 2002). Turner et al (1993) added Nd isotopic evidence, demonstrating substantial overlap of eNdi in the Lachlan Fold Belt and age-corrected ENdi of the late Proterozoic and Cambrian sediments in the Delamerian Fold Belt. They suggested that Lachlan Fold Belt flysch could represent recycled, weathered Delamerian metasediment, but with some input from less evolved, mantle-derived Nd sources such as the late Delamerian intrusive rocks. The data presented here support this conclusion. The older Takaka terrane samples [44-46] have Sri values >0.7097, clearly belonging with other Western Province data; their eNdi values of -11.0 to -9.2 coincide with the lower end of the data for the Buller terrane and Lachlan Fold Belt. Close to the eastern margin of the Takaka terrane at Parapara Peak, a tectonic enclave of amphibolite-facies metasediments may represent a Permian-Triassic cover sequence to the terrane (Campbell et al 1998; Rattenbury et al 1998). The E Ndi value, +2.7 [48], is completely outside the range of the Western Province data and is closer to Permian or Triassic data from the Eastern Province, e.g. Murihiku or Caples terranes. This might indicate that the Parapara Peak rocks originated in the Eastern Province (below), but have been emplaced tectonically

127

into the Western Province, Takaka terrane. Its Sri value, 0.7005, is anomalous and unrealistically low. This suggests that Sr isotope exchange has occurred between sediment detritus and a less radiogenic extraneous source, accompanied by a significant increase in Rb-Sr ratio (for example, Sr loss from calcareous fossil fragments destroyed during metamorphism). A siltstone from Devonian sediments (Baton Formation) (Bradshaw 2000) has Sri 0.7048 [45] and a second sample [46] is even more anomalous. Since metasediments at this locality also frequently contain calcareous fossils, then an explanation similar to the Parapara Peak example above seems possible. Sr-isotope resetting anomalies are present in many Cambrian samples [49-62] from the Takaka terrane, analysed by Wombacher & Munker (2000), which they attribute to isotopic resetting of terrestrial sediment with an extraneous (possibly carbonate) component. This is clearly a problem throughout the Takaka terrane, since, unlike all others, limestones are important in Cambrian and Ordovician units and calcareous fossils are commonly found in siliciclastic units throughout. These authors note, however, that the Nd isotope systematics are unaffected and, although partly overlapping Buller terrane values, their E Ndi values, mostly in the range, -11 to +0.2, reflect the persistent volcanic/volcaniclastic nature of the Cambrian successions more strongly. Within the Buller terrane, Devonian quartzites and siltstone/mudstone of the Reefton Group form local outliers upon Greenland Group (Bradshaw 1995; 2000). In this case, the £ value, -10.9 [13], is just outside the range of older Buller terrane data (above) and its Sri value, 0.7110, places it within the Takaka terrane dataset (Fig. 4). However, this siltstone is from a locality with occasional calcareous fossils and the anomalies encountered with samples [45-46] and [48] might also arise.

Results for New Zealand Eastern Province terranes The Nd-Sr isotopic data for the seven Eastern Province terranes comprise a connected

Fig. 4. Initial epsilon Nd and initial 87Sr/86Sr ratio data calculated at the estimated time of deposition for metasediments of New Zealand Western and Eastern Province terranes. Data for each terrane are grouped within solid line envelopes. Also shown are examples of U-Pb detrital zircon age patterns for greywacke/ sandstone samples (mostly of Triassic age) from selected terranes (Ireland 1992«; Pickard et al. 2000; Adams et al 2002; Adams unpublished data). Those at the top are from Eastern Province volcaniclastic terranes, those at the right are from Eastern Province accretionary terranes, those at the bottom are Western Province.

128

C. J. ADAMS ET AL.

Fig. 5. Initial Nd isotopic (epsilon Nd) values for metasediments of New Zealand Eastern Province terranes, displayed from west to east across the diagram. Data box widths indicate lithology (sandstones wide, siltstones medium, mudstones narrow). Data of the present study (black) are aligned on the vertical dashed gridlines and grouped within fine solid line envelopes, those of Frost & Coombs (1989) are set to the right.

sequence of individual terrane arrays (Fig. 4), each having some degree of overlap (and substantial for the Caples terrane). This overlap is related, in part, to uncertainties in the assignment of particular localities into Caples, Waipapa and Rakaia terranes discussed below. With this reservation, the sequence follows a progression from primitive (Brook Street terrane) to evolved (Pahau terrane) rock characteristics. The Nd-Sm data in this general trend approach Buller/Takaka compositions, but without reaching them. With the Caples terrane exception (see below), the sequence follows the present terrane positions from west to east, but significantly, the most evolved Rakaia and Pahau terrane data are closest to those from the Duller terrane, from which they are geographically the most distant. This feature clearly highlights the 'suspect' nature of at least the Rakaia and Pahau terranes. The new ENdi data fall within the earlier data of Frost & Coombs (1989), but are grouped more closely (Fig. 5). This feature reflects their choice of either coarser (sandstone) or, to a lesser extent, finer (mudstone) lithologies, in contrast to the almost exclusively medium

grain-size (siltstone) samples of the present study. The new results are closer to their sandstone data for Brook Street, Murihiku and Dun Mountain-Maitai terranes, but mostly to their mudstone/siltstone data for Rakaia and Pahau terranes. From their sandstone-mudstone pairs in the latter, Frost & Coombs (1989) concluded that Nd is concentrated in younger sources in the coarser grain-size sediments, a pattern confirmed by the new siltstone data which, within each terrane, consistently have ENdi values lower than, or at least comparable with, the Frost & Coombs (1989) sandstone datasets.

Brook Street terrane The Brook Street terrane is dominated by redeposited volcaniclastic sedimentary rocks around several Permian basic-intermediate volcanic centres. Most aspects of the petrochemistry and Sr-isotope geochemistry strongly suggest an origin from mantle-derived (oceanic arc or back-arc) volcanic source rocks (Mortimer et al. 19990, b\ Adams et al 2002). The new isotopic data from the Takitimu Group [63-70] and correlatives [71] support this

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN

conclusion. However, in the Nelson area (from where the terrane derives its name), the Brook Street terrane is reduced in size and the equivalent late Permian volcaniclastic metasediments (and rare limestones) of the Grampian Formation (Johnston 1981) have far more evolved Nd and Sr isotopic values, ENdi = -0.7 and Sri = 0.7052 [72], than any Southland counterparts, 8Ndi = +5.4 to +8.4 [63-71]. This raises some doubt about the correlation (detailed petrographical and geochemical comparisons are unavailable) and the Grampian Formation is better compared isotopically, with late Permian analogues in the Dun Mountain-Maitai [87, 88] or, perhaps, Murihiku [73] terranes. With this uncertainty, the Grampian Formation data are not included in the Brook Street terrane data field in Figures 4 and 5.

Murihiku terrane The Murihiku terrane comprises voluminous Permian, Triassic and Jurassic volcaniclastic sedimentary successions (Murihiku Supergroup, Campbell & Coombs 1966), generally reflecting intermediate compositions, in two broad synclines in North Island and South Island (Kawhia and Southland, respectively). There are surprisingly few associated volcanic or intrusive rocks. Frost & Coombs (1989) noted an unusual degree of variability in Nd data for Murihiku terrane, with eNdi values in the range -0.6 to +6.0, within which our new data [73-77, 83-85] fall. However, as for the Brook Street terrane, the eNdi-Sri data (Fig. 4) support geochemical evidence for sediment provenance from a back-arc (or island-arc) environment, with little, or no, terrigenous detritus.

Dun Mountain-Maitai terrane The Dun Mountain-Maitai terrane is the most varied of the Eastern Province terranes, comprising Early Permian ophiolites (Dun Mountain Ophiolite Complex, Coombs et al 1976), Late Permian re-deposited calcturbidites and sandstones (Landis 1974; 1980) and Early-Middle Triassic turbidites (Maitai Group, Cawood 1986; 1987). However, there is persistent evidence for volcaniclastic detrital inputs throughout the Maitai Group, and tuffs (but not flows) are frequent in the Triassic part. This explains the considerable variation, noted by Frost & Coombs (1989), in £ Ndi values -2.7 to +6.5, from the Maitai Group. The new data [87-89, 91, 95, 96, 98, 99] (Figs 4, 5) show less variation, with e Ndi -0.7 to +3.9, with a corre-

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sponding Sri range of 0.7038-0.7071 (excluding one anomalously low sample [89]). The Permian rocks [87, 88] are the more evolved, ENdi = -0.7 to -0.2 and match a trend, seen throughout the terrane, of higher, i.e. more evolved, Sri values in the Permian than the Triassic (Adams et al. 2002). This trend is mirrored by detrital zircon data, which indicate that Maitai Group sediments received greater proportions of continent-derived detrital zircons in the Permian than the Triassic (Adams et al 2002, and unpublished data).

Caples terrane In the type area of Caples Group in Southland (Turnbull 19790, b) several formations are defined, but elsewhere the extensive, unfossiliferous, turbidite sequences cannot be subdivided formally. A Permian to Triassic age range is probable and the Momus Sandstone Formation in northern Southland is definitely Triassic (CJA, unpublished U-Pb detrital zircon ages). With two exceptions discussed below, the new Caples terrane samples [100-103,105] yield eNdi -1.6 to +4.0, and Sri 0.7035-0.7048, forming a more compact group than the Frost & Coombs (1989) data with £ Ndi +0.1 to +6.2. Caples terrane rocks extend into East Otago, but on the coast between Dunedin and Balclutha they are considered atypical and, on geochemical characteristics, have a 'Torlesse affinity' (Mortimer 19930). Two samples [110, 111] from this area are at the more evolved end of the Caples range, with 8Ndi values of -1.5 to +0.0, and Sri 0.70470.7049. These values are similar to comparable Waipapa terrane data (below) - an alternative assignment suggested, on the basis of Rb-Sr age and isotope characteristics alone, by Adams & Graham (1997). Correlation of Caples terrane rocks into the North Island, to Northland (Black 1994), is also supported by new data for sample [100] from Puketi Forest.

Aspiring terrane The Aspiring terrane forms a distinctive (greenschist facies) zone of uncertain protolith within the Haast Schist (Craw 1984; Norris & Craw 1987). Falling between schists of true Caples and Torlesse (Rakaia) terrane turbidite-dominated protoliths, the combination of mica-schists, metabasites and metacherts suggests a pelagic association. Frost & Coombs (1989) reported Nd isotope data for greenschist [113], possibly a metabasite, £ Ndi - +6.2, and metapelagite [114], £ Ndi - -1-1 at Nevis Bluff, Otago, in the higher grade part of the terrane. New data from lower

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greenschist fades rocks, in the Matukituki River valley [112], with ENdi -1.6, and Sri 0.7043, are similar to the metapelagite value. Together, the £ Ndi values fall between Caples (sensu stricto) and Torlesse (Rakaia) terrane data and are comparable with those of Waipapa terrane. Waipapa terrane The Waipapa terrane (Sporli 1978) is defined in the northern part of North Island, where several sedimentary associations of Permian, Triassic and Jurassic, moderately volcaniclastic (acidintermediate) rocks are recognized (Black 1994). These are commonly unfossiliferous, terrigenous clastic rocks of fine sandstone and siltstone lithologies, but occasionally include important horizons of fossiliferous pelagic mudstones, limestones and spilites. Petrography, geochemical and Rb-Sr age and Sri characteristics suggest terrane continuation into Marlborough and Wellington regions (Mortimer 1993/?; Adams et al 1999) and, more speculatively, into East Otago (Adams & Graham 1997). Despite their broad range, ENdi values -2.9 to +3.9 and Sri values 0.7039-0.7063 [115-122] for the Waipapa terrane form a group spanning the gap between Caples and Rakaia terranes, and show generally decreasing ENdi with stratigraphic age. However, the age of the least evolved sample [116], from Northland, is uncertain, since successions nearby have Permian fossils (but possibly reworked), but the local sandstones contain Late Triassic and Early Jurassic detrital zircons (Pickard et al. 2000). Similarly, the position of the most evolved sample [119] from Kapiti Island is also ambiguous, since petrographically it resembles the Torlesse Supergroup of the Rakaia terrane, but has unusually low Sri and metamorphic age characteristics, more typical of Waipapa terrane rocks (Adams & Graham 1996). Rakaia and Pahau terranes (Torlesse composite terrane) Distinctive quartzose greywacke-dominated Torlesse Supergroup metasediments dominate the Torlesse composite terrane, and both Permian-Triassic Rakaia and JurassicCretaceous Pahau terrane rocks show clear petrographical, geochemical and Rb-Sr age and Sri patterns, indicating deposition from continental sources of I-type granitoid type (MacKinnon 1983; Roser & Korsch 1988; Adams & Graham 1996; Adams & Maas 2004). Their detrital U-Pb zircon and 40Ar/39Ar

muscovite age populations are the most complex of all the Eastern Province terranes, with significant Palaeo-, Meso- and Neoproterozoic, and early and late Palaeozoic contributions (Ireland 19920; Adams & Kelley 1998; Adams et al 1998; Pickard et al 2000). The Ndi and Sri data [123, 124,126,127,130-134,136-138] confirm these as the most evolved group of Eastern Province rocks, with eNdi values of -1.9 to -4.7 and Sri of 0.7060-0.7081, again with a trend towards lower ENdi and higher Sri from Permian to Triassic. There is, however, a large (c. 25%) overlap between Rakaia and Waipapa terrane data, particularly Permian data of the former [123, 124, 126] with Jurassic data of the latter [118, 120-122]. Data for the 'Older Torlesse' (mostly Triassic) in Frost & Coombs (1989) have £Ndi -4.6 to -2.0, very similar to the new analyses (Fig. 5), despite a broader lithological range (sandstone to mudstone). Ndi and Sri data for the Pahau terrane [144, 145, 148, 149] follow the general Rakaia trend (above) and, although only an approximate subset of the broader Rakaia terrane data field, these youngest, Early Cretaceous, sedimentary rocks are amongst the most evolved of all the Eastern Province (Fig. 4). Data groups in the Eastern Province terranes The Nd-Sr isotopic data follow a simple trend (Fig. 4) that starts with a Brook Street terrane end-member, coincident with the conventional mantle array and reflecting solely juvenile inputs, and then progressively demonstrates increasing contributions of continent-derived detritus. Matching detrital zircon patterns for greywackes and sandstones (a set of mostly Triassic examples is shown in Fig. 4) reflect this evolution, showing increasing proportions of older age components as £Ndi progresses towards more negative values. Given the potential for widely differing Sr-Nd isotopic signatures associated with the various zircon age components, it is interesting to note (1) the relatively well-confined isotopic signatures for each terrane and (2) a consistent evolution sequence, with some overlap, from one terrane to its geographical neighbour (excepting the Caples terrane, discussed below). The former feature suggests a high degree of sediment homogenization during collection from an extensive hinterland, in large-scale transport on major river and delta systems, and delivery by submarine canyons to large sedimentary basins (on the scale of hundreds of kilometres). The

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latter feature requires a clear geographical separation of terranes (on a similar scale, assuming the absence of physical barriers) sufficient to sustain their different sediment provenances and tectonic associations. On a larger scale (perhaps a thousand kilometres) they show a gradual change in their relative importance of these factors. Thus, the Brook Street depocentre, although adjacent to an active continental margin, occurs within an offshore volcanic island arc isolated from continental influence. On the other hand, at the same margin and at the same time, the Torlesse depocentre was controlled by continent-derived detritus and isolated from contemporaneous volcanism. The anomalous position of the Caples terrane dataset in the general sequence of Figure 4 can be resolved if the sequence is reconsidered as two overlapping arrays, comprising: (1) a Brook Street, Murihiku and Dun Mountain-Maitai terrane group, dominated by back-arc, forearc and/or island-arc volcanic environments, but including a platform sedimentary successsion and (2) Caples, Waipapa, Rakaia and Pahau terrane group, dominated by mid-fan turbidite environments that, with more pelagic tectonic intercalations, form accretionary complexes. The 'volcanic/volcaniclastic' group shows only a modest evolutionary trend culminating with the Dun Mountain-Maitai terrane showing some continental influence in its detrital age/isotopic signatures. The isotopic overlap of the Dun Mountain-Maitai and Caples terranes raises the possibility that these represent a transitional tectonic and depostional environment, with generally similar sediment provenances. A detailed comparison is hindered by the lack of more comprehensive detrital zircon data and isotopic characteristics for the Caples terrane rocks, i.e. Caples Group, in their type area.

Results from West Antarctica and southern South America The data presented in Table 2 represent metasediment samples with a wide range of ages and tectonic environments. The data for continental West Antarctica (excluding the Antarctic Peninsula, this equates to the Eastern Domain of Vaughan & Storey 2000) are for Cambrian to Permian metasedimentary rocks from the Thiel Mountains [268], the EllsworthWhitmore mountains crustal block [269-277], the Ellsworth Mountains [278, 279], the Transantarctic Mountains [280] and Marie Byrd Land [281-283]. Despite a wide range of

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lithologies, stratigraphic ages and tectonic situations, the majority have £Ndi -4.1 to -7.8; only the Cambro-Ordovician Thiel Mountains sample (Pankhurst etal 1989) and the ?Carboniferous samples from Ickes Mountains from western Marie Byrd Land (Pankhurst et al 19985) falling outside this range at -3.4 and -10, respectively. Sri values of 0.7078-0.7233, discounting those poorly constrained due to high Rb-Sr ratios, also indicate mature continental crust as the principal provenance for this group. In the northern Antarctic Peninsula, preJurassic metasedimentary samples [221-229] correspond to the Trinity Peninsula Group of Hyden & Tanner (1981) and presumed equivalents in the Miers Bluff Formation, South Shetland Islands. The latter [223, 224] have yielded a Triassic Rb-Sr age taken as approximating deposition (Willan et al 1994), although recent microfossil evidence suggests that Cretaceous components may be interleaved in the Miers Bluff outcrop (Pimpirev et al 2003). The LeMay Formation of Alexander Island [208-220], which falls within the Western Domain (?terrane) of Vaughan & Storey (2000), also has evidence for both Triassic and Cretaceous deposition, in this case dated by the Rb-Sr whole-rock isochron method (Pankhurst et al unpublished data). Together with a few Jurassic samples from elsewhere in Alexander Island [201-206], all these samples have a relatively narrow range of relatively mature ENdi at the time of sedimentation, -6.7 to -2.0. These depositional ages of these rocks cover essentially the same range as those of the Torlesse and they are, in general, isotopically comparable; there is a tendency to rather more negative £Ndi and more radiogenic Sri (up to 0.7114) seen in the most mature examples. The oldest samples from southern South America have maximum possible stratigraphic ages constrained largely by unpublished U-Pb detrital zircon studies. These are Eocambrian to Early Ordovician metapelites and phyllites from Patagonia east of the Andes [289-291, 297-299] and the Sierras Pampeanas of NW Argentina [284-288]. The Patagonian samples are largely of uncertain origin, no underlying basement having been identified; the Pampean rocks were formed and deformed during the mid-Cambrian Pampean orogeny, probably as a result of the collision of a parautochthonous terrane with the margin of the Rio de La Plata craton (Rapela et al 1998). The Precambrian samples from southern Argentine Patagonia have eNdi values of -3.0 to -2.3 and Sri 0.7096-0.7128, whereas most Cambro-Ordovician metapelites from

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northern Patagonia have more negative eNdi values of -5.6 to -4.0, but with similar Sri. The Esquel schist sample [293] appears to be isotopically more juvenile (eNdi = 0.7, Sri = 0.7051) but this may be due partly to overestimation of its age based on detrital zircons. The Sierra Grande meta-sandstone [289] has disturbed Sr isotopes, probably due to contact metamorphism from a nearby Cambrian granite. The Pampean schists and phyllites reach even more extreme values of -7.7, with Sri 0.7107-0.7158, possibly suggesting an older average source. All of these metasediments would be considered as having a primary provenance within cratonic Gondwana. Extensive Ordovician deposits from the Puna (north of the Sierras Pampeanas, NW Argentina) [300-328], show a wide range of ENdi from -11.6 to -1.4. Zimmermann & Bahlburg (2003) interpreted the less negative values from the Tolillar and Diablo formations [317-323] as reflecting a higher proportion of volcanogenic input associated with the broadly contemporaneous Famatinian magmatic cycle (Pankhurst et al 19980). In contrast, the Lower Ordovician Tolar Chico Formation is characterized by nonvolcanic provenance and uniformly low £ Ndi averaging c. -10. Intermediate ENdi values in the later sediments of the Coquena and Falda Cienaga formations could be due to mixing between these two sources or provenance from younger crustal basement. The Sri values of the Puna samples (0.7036-0.7088 for the volcanogenic rocks, mostly >0.709 for the remainder) would be consistent with either interpretation. The low-grade metasedimentary rocks of southern Chile occur in two belts, partly separated by the Mesozoic-Cenozoic Patagonian batholith. Parts of the Eastern Andes Metamorphic Complex (Herve 1993) could have depositional ages as old as Devonian according to their detrital zircon content, whereas other parts must be as young as Permian (Herve et al 2003). Augustsson (2003) argues that this complex is essentially Carboniferous in age, which is also consistent with the sparse fossil evidence. Nd data for samples near the town of Cochrane have been published by Augustsson & Bahlburg (2003). Assuming an age of 350 Ma [257-267], 8Ndi for these are quite uniform, -6.6 to -4.6, with a further sample from the Lago General Carrera unit to the north giving a higher value of -2.2. If the age is taken as 320 Ma, £Ndi become more positive by about 0.3, but the crustal model ages (see below) are scarcely affected. The samples from the current study are from this latter unit [254-256] and give eNdi values of-7.1 to -3.5 and Sri 0.7090-0.7111,

assuming an age of 320 Ma. In general, these eastern metasedimentary rocks thus show a predominance of continental crustal input comparable to that inferred for the early Palaeozoic Gondwana low-grade metasedimentary rocks mentioned above. The western belt of metasediments crops out mainly in offshore islands. The majority of the samples here are from the Chonos accretionary complex, which is considered to have a Late Triassic depositional age (Fang et al. 1998; Herve et al 2003), which mostly have ENdi -7.3 to -5.1, with only two samples falling outside this range at -3.7 and -4.6; Sri values are 0.7066-0.7195. These isotope compositions are indistinguishable from those of the eastern belt, despite significantly younger stratigraphic ages. Two mylonitized samples from the Jurassic Diego de Almagro complex (see Herve & Fanning 2003) have contrasting Sm and Nd concentrations, £Ndi of -5.4 and +4.2 and Sri of 0.7054-0.7074, assuming an Early Jurassic age; the isotopically more primitive sample is probably derived from a metabasic rock of oceanic origin. The Huinay schist ([241], age unknown, but assumed to be late Carboniferous on the basis of unpublished Rb-Sr data) is the most evolved sample in this group, with £Ndi -6.5 and Sri 0.7114. Samples from the South Shetland Islands, mostly taken from Trouw et al (1990) have £Ndi values as high as +4.5 (calculated at 100 Ma, the time of major metamorphic resetting - assuming an early Mesozoic age makes calculated £Ndi values only slightly more positive). This indicates juvenile protoliths of oceanic type, consistent with their high 147Sm-144Nd ratios (>0.15) and low Sri of <0.7050. In both senses they show an affinity with samples from the volcanic and volcaniclastic terranes of New Zealand, although the even higher £Ndi of the latter imply an even more LREE-depleted source. However, a few South Shetland Islands samples [233, 238-240] are indistinguishable isotopically from those described above as having a Gondwana continental provenance, with £Ndi -6.7 to -4.8 and Sri 0.7070 to 0.7098.

Crustal residence ages and their implications Model ages based on Nd isotope composition depend on assumptions about the fractionation of Sm-Nd ratios in major crustal processes. In the simplest case, the only significant fractionation event is considered to be the extraction of primary crustal material from long-term LIL (large ion lithophile element)-depleted upper

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN

mantle (e.g. DePaolo 1981), so that the age calculated should give an estimate for this crust formation event. However, this ignores fractionation processes within the crust by processes such as metamorphism, anatexis and weathering, all of which are likely to have occurred in the pre-history of sediment source areas. The inadequacy of the simple TDM model in such cases is highlighted by the unrealistic and often highly negative model ages that result (indicated by blank entries in Tables 1 and 2). The next stage of model age calculation allows for an average crustal Sm-Nd ratio during the period of crustal residence prior to the last possible fractionation event - in this case the erosion and deposition of the sedimentary rocks. More complex models might allow for subsequent metamorphism, but this is not feasible in the general case due to lack of quantitative control on such individual fractionations. Therefore, two-stage Nd model ages (TDM#) have been calculated as defined by DePaolo et al (1991), very slightly modified to allow for a different mass-spectrometric normalization procedure, and these are used in the following discussion. These yield an estimate for the age of derivation of the crustal protolith from average upper mantle, allowing for continuous LIL-depletion of the latter (and consequently LIL-enrichment in the crustal protoliths) over geological time. The calculation depends essentially on two parameters - the eNdi value and the age of deposition. It should be emphasized that the values of model ages depend very critically on the exact model parameters employed, especially for the 147 Sm/144Nd ratio during the first crustal stage. The model of DePaolo et al. (1991) allows for continuously increasing 147Sm/144Nd in newly formed crust over geological time. In the model age range presented here this ratio is mostly close to 0.130, so that ages calculated using this formulation are older significantly than those which use a constant, average crustal 147 Sm/144Nd ratio of <0.120 (e.g. Liew & Hofmann 1988). The difference between the two models is most significant for the Early Palaeozoic sediments, whose DePaolo et al. (1991) model ages are up to 250 Ma older than those of Liew & Hoffman (1988). Despite this sensitivity to model parameters, in practice the calculated TDM# age allows a more understandable discrimination between the isotopic characteristics of the wide range of sample data and ages presented here in the database; the results are shown in the histograms of Figure 6. In the case of New Zealand, there is a marked distinction between

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the Western and Eastern Province data. More than 90% of the Buller terrane (Western Province) samples have model ages of 1750-2000 Ma, whereas those for the Eastern Province samples (Central Arc, as well as the Caples/Aspiring/Waipapa and Torlesse terranes) do not exceed 1400 Ma. This must reflect very mature continental crust in the provenance areas of the Buller terrane; if it had all formed in a single event, the model would predict ultimate derivation from an early Mesoproterozoic or late Palaeoproterozoic craton. However, subsequent transformation by processes such as incorporation into younger S-type granites cannot be resolved and, in the case of multiple or mixed sources, it is quite possible that some components could be derived from exposed sources as old as Archaean. Supporting information for this type of provenance is given by detrital zircon age patterns (Fig. 4), with a very minor Archaean component, rather more Mesoproterozoic and a preponderance of Neoproterozoic and early Palaeozoic zircons. This is a typical multicomponent Gondwana pattern (Ireland et al 1998) - the Nd model age calculations here suggest that many of the younger zircons are likely to have been derived from anatectic or high-grade metamorphic rocks with older protolith ages. The model ages for the Takaka terrane are not uniformly so old, generally extending down to c. 1400 Ma and, in the case of two samples, as low as 1027 Ma and 807 Ma. As noted above, this may be due to some more juvenile input or minor disturbance of the Sm-Nd systems during the metamorphic events that disturbed the Rb-Sr systems severely. In contrast, samples from the volcanic/ volcaniclastic (Central Arc) terranes of the Eastern Province show model ages extending down to 300 Ma, which is very little older than the depositional age of the sedimentary rocks. This is in accord with lithological evidence that a major component was derived from contemporaneous igneous rocks and the abundance of detrital zircons in the 200-300 Ma age range (Fig. 4), although the scatter of TDM# ages up to 1070 Ma (1262 Ma in one case) is in agreement with the Nd-isotope interpretation of mixing with some material of older crustal provenance. The western terranes of the Eastern Province accretionary complex-type (Caples, Aspiring and Waipapa) have a wide range of slightly older TDM# ages (c. 450-1300 Ma) that could reflect a similar mixing of detritus from juvenile and mature sources, with a higher proportion of the latter. If mature crust were the preponderant end-member, it would appear to be of early

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Fig. 6. Histograms of Nd model ages (DePaolo et al. 1991) for data from terranes in New Zealand (Table 1), West Antarctica and southern South America (Table 2). The New Zealand data are subdivided as follows: Western Province as combined Buller and Takaka terranes; Eastern Province as a central arc terranes group (Brook Street, Murihiku and Dun Mountain-Maitai terranes) and an accretionary group (C-A-W, Caples, Aspiring and Waipapa terranes). West Antartica: TPG = Trinity Peninsula Group. For South America, the 'Western Domain' includes all samples from the Chonos Archipelago and Isla Diego de Almagro, but is mostly from the Chonos metamorphic complex. EAMC = Eastern Andes Metamorphic Complex. The ages are estimates of the 'average' time that the parent material spent within the continental crust, although in many cases this may represent mixing of detritus from rocks with very different formation ages.

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN

Mesoproterozoic age, but this cannot be discerned from the detrital zircon age data (where available), which only shows a slightly older range of ages than in the volcanic/volcaniclastic terranes (Fig. 4). Finally, samples from the Torlesse composite terrane have Nd model ages restricted largely to the 1150-1400 Ma interval. Although this might imply a fairly uniform early Neoproterozoic source, the ages of zircon incorporated into these rocks barely extends to this age range and shows a great preponderance of Permo-Triassic detritus (Fig. 4). The implication is that many of these zircons, which generally have zoned igneous cores, could be derived from S-type granites generated by melting of older continental crust. The results from West Antarctica lack the strongly crustal signature of the New Zealand Western Province data. The oldest model ages, for the deposits of the Ellsworth Mountains, the Ellsworth-Whitmore crustal bock and Marie Byrd Land, range from 1370 Ma to 1840 Ma; this compares well with the data for the Buller terrane of New Zealand, albeit suggesting a somewhat greater contribution of detritus from younger arcs mixed with the cratonic provenance. A zircon inheritance pattern for Ellsworth Mountains quartzite (Millar et al. unpublished data) is very similar to that for the Greenland Group, although it actually has a more prominent Archaean peak. The South Shetland Islands sedimentary rocks, for which there are no U-Pb zircon data, have a bimodal distribution of model ages, some juvenile (Palaeozoic) and some with a Gondwanaland provenance. The model ages of Triassic deposits of the Trinity Peninsula Group and of the LeMay and Miers Bluff formations are indistinguishable from each other (1120-1490 Ma) and also correspond closely to those of the Torlesse composite terranes. Zircon inheritance patterns for these groups (Millar et al. 2001; 2002; unpublished data) also show the same type of distribution as the Torlesse, with a major Permian input, a moderate early Palaeozoic component and a scattering of older ages. A high degree of crustal re-working in the source regions is indicated, as is also shown by Hf-isotope data for the detrital zircons (Millar et al. 2003). In this case, possible sources may be identified in the Antarctic Peninsula itself, where sparse Palaeozoic and Permo-Triassic granitoids are characterized by negative E Ndi and relatively high Sri (Millar et al 2002; Wever et al. 1994). In southern South America, the metasediments that are presumed to have a predominantly Gondwana provenance (from the Puna and Sierras Pampeanas) have a considerable

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spread of Proterozoic model ages (c. 13202050 Ma), which encompasses those of the New Zealand Western Province and the interior regions of West Antarctica. Most of these samples are associated with early Palaeozoic magmatic arcs, so that the spread is explained most easily by incorporation of contemporaneous I-type sources. Some of the oldest metasedimentary rocks contain a high proportion of Precambrian (especially c. 600 Ma and c. 1100 Ma) detrital zircons (Pankhurst et al. 20015). The metasediments from Argentine Patagonia, which include both lower and upper Palaeozoic deposits, have a range of model ages of 1130-1660 Ma, towards the younger end of the Gondwana provenance range and, in many of these, a major input of early Palaeozoic detrital zircons is also apparent (Pankhurst et al 20010; 2003; Rapela et al unpublished data). Model ages for the Eastern Andes Metamorphic Complex are essentially indistinguishable from these (1290-1640 Ma), suggesting a similar balance of inputs from the same source regions. Surprisingly, the Nd model ages of the Chonos Formation ([242-250]: 1320-1580 Ma) also fall within the same range, despite the fact that they are younger deposits, presumably more distal with respect to Gondwanaland continental crust, and contain a large proportion of Permian to Triassic zircons (Herve et al 2003). This again suggests that the penecontemporaneous arcs that contributed to the detritus - in this case, as yet only identified locally in the North Patagonian Massif (Pankhurst et al 1992) were generated largely by reworking of older Gondwanaland continental crust. However, there is minimal overlap between the model ages of any of the South American Gondwana margin metasediments and those of the New Zealand Central Arc or Caples-AspiringWaipapa terranes, so that a common provenance for these is very unlikely. There is more consistent overlap of TDM# ages between the western accretionary complexes of Chile, the LeMay Formation of Alexander Island and the Torlesse terranes of New Zealand, with a tendency to slightly younger ages through this sequence.

Summary and conclusions Examination of the large Nd-Sr isotope database that has been assembled yields a more detailed characterization of the terrane structure of New Zealand than previously, allowing comparison with metasedimentary rocks of accretionary complexes along the margin of Gondwana as far as southern South America.

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Although there are severe limitations on the interpretation of Nd model ages, principally due to the mixing of sedimentary material from both mature and juvenile sources in the depositional environment, some broad generalizations may be recognized, and these are further constrained where more specific data on the ages of detrital zircons are also available. Material of mature Gondwana continental provenance is predominant in sedimentary basins located inboard of the active margin, especially during the Palaeozoic (Western Province of New Zealand, interior West Antarctica and NW Argentina), and is also a major contributor in the preMesozoic sedimentary deposits of Argentine Patagonia. In contrast, the western part of the Eastern province terranes of New Zealand (the Central Arc and the westernmost accretionary complexes), which are of late Palaeozoic to early Mesozoic depositional age, have Nd isotope systematics that reflect much younger input, with contemporaneous arc material often dominating the provenance. The Scotia metamorphic complex of the western archipelago of the Antarctic Peninsula is comparable in this to the Central Arc terranes, although its depositional age is not well constrained. The easternmost New Zealand terranes (Torlesse) and the majority of the late Palaeozoic to early Mesozoic accretionary complexes of the Antarctic Peninsula and South America generally have Mesoproterozoic TDM# model ages; these signify thorough mixing between continental Gondwanaland sources and contemporaneous arc material. Discrepancies with the ubiquitous influx of Permian-Triassic detritus (signified by detrital zircon age studies) show that many of the arc source rocks must have been generated through significant reworking of older crust. The younger provenance of the Torlesse composite terranes of New Zealand appears to be similar to that of contemporaneous metasedimentary rocks on the South American proto-Pacific margin. More detailed investigation, for example using Hf-isotope study of detrital zircons, should refine these conclusions and should assist in identifying the most likely source regions for the constituent sedimentary successions of tectonostratigraphic terranes. CJA thanks the Alexander von Humboldt Foundation (Germany) and Transantarctic Expedition Association (NZ-UK) for grants enabling presentation of this work at TAPMOG (Cambridge) and ISAES (Potsdam) Symposia in September 2003, and subsequent preparation of the present work. RJP acknowledges the support of a Leverhulme Trust Emeritus Fellowship. RM thanks Gerry O'Halloran (ESSO Australia) for permission to use unpublished data.

The authors thank Carita Augustsson and Francisco Herve for helpful reviews of this paper.

References ADAMS, CJ. 1986. Geochronological studies of the Swanson Formation of Marie Byrd Land, West Antarctica, and correlation with Northern Victoria Land, East Antarctica, and South Island, New Zealand. New Zealand Journal of Geology and Geophysics, 29, 345-359. ADAMS, CJ. 1997. Initial strontium isotopic signatures of Late Precambrian-early Paleozoic metasediments from northern Victoria Land terranes, East Antarctica. In: Rica, C. A. (ed.) The Antarctic Region: Geological Evolution and Processes. Terra Antartica Publication, Siena, 227-236. ADAMS, CJ. 2004. Rb-Sr age and strontium isotope characteristics of the Greenland Group, Duller terrane, New Zealand, and correlations at the East Gondwanaland margin. New Zealand Journal of Geology and Geophysics, 47,189-200. ADAMS, CJ. & GRAHAM, IJ. 1996. Metamorphic and tectonic geochronology of the Torlesse terrane, Wellington, New Zealand. New Zealand Journal Geology and Geophysics, 39, 157-180. ADAMS, CJ. & GRAHAM, IJ. 1997. Age of metamorphism of Otago Schist in eastern Otago and determination of protoliths from initial strontium isotope characteristics. New Zealand Journal of Geology and Geophysics, 40, 275-286. ADAMS, CJ. & KELLEY, S. 1998. Provenance of PermoTriassic and Ordovician metagreywacke terranes in New Zealand: evidence from 40Ar/39Ar dating of detrital micas. Geological Society of America Bulletin, 110, 422-432. ADAMS, CJ. & MAAS, R. 20040. Rb-Sr age and strontium isotopic characterisation of the Torlesse Supergroup in Canterbury, New Zealand, and implications for the status of the Rakaia terrane. New Zealand Journal of Geology and Geophysics, 47, 201-217. ADAMS, CJ. & MAAS, R. 20046. Age/isotopic characterisation of the Waipapa Group in Northland and Auckland, New Zealand, and implications for the status of the Waipapa Terrane. New Zealand Journal of Geology and Geophysics, 47, 173-187. ADAMS, C.J.D., HARPER, C.T & LAIRD, M.G 1975. K-Ar ages of low grade metasediments of the Greenland and Waiuta Groups in Westland and Buller, New Zealand. New Zealand Journal of Geology and Geophysics, 18, 39^8. ADAMS, C J., MORRIS, PA. & BEGGS, J.M. 1979. Age and correlation of volcanic rocks of Campbell Island and metamorphic basement of the Campbell Plateau, South-west Pacific. New Zealand Journal of Geology and Geophysics, 22, 679-681. ADAMS, CJ., BARLEY, M.E., FLETCHER, I.R. & PICKARD, A.L. 1998. Evidence from U-Pb zircon and 40Ar/39Ar muscovite detrital mineral ages in metasandstones for movement of the Torlesse suspect terrane around the eastern margin of Gondwanaland. Terra Nova, 10,183-189.

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN ADAMS, C.J., GRAHAM, IJ. & JOHNSTON, M.R. 1999. Age and isotopic characterisation of geological terranes in Marlborough Schist, Nelson/Marlborough. New Zealand Journal of Geology and Geophysics, 42, 33-55. ADAMS, C.J., BARLEY, M.E., MAAS, R. & DOYLE, M.G. 2002. Provenance of Permian-Triassic volcaniclastic sedimentary terranes in New Zealand: evidence from their radiogenic isotope characteristics and detrital mineral age patterns. New Zealand Journal of Geology & Geophysics, 45, 221-242. AUGUSTSSON, C. 2003. Provenance of Late Palaeozoic sediments in the southern Patagonian Andes: age estimates, sources, and depositional setting. PhD thesis, Westfalische Wilhelms-Universitat, Minister. AUGUSTSSON, C. & BAHLBURG, H. 2003. Active or passive margin? Geochemical and Nd isotope constraints of metasediments in the backstop of a pre-Andean accretionary wedge in southernmost Chile (46°30'-48°30' S). In: McCANN, T (ed.) Tracing Tectonic Deformation using the Sedimentary Record. Geological Society, London, Special Publications, 208, 253-268. BALLANCE, P.P. & CAMPBELL, J.D. 1993. The Murihiku arc-related basin of New Zealand (TriassicJurassic). In: BALLANCE, PF. (ed.) South Pacific Sedimentary Basins. Sedimentary Basins of the World, 2. Elsevier, Amsterdam, 21-33. BIRCH, W.D. (ed.) 2003. Geology of Victoria. Geological Society of Australia, Special Publications, 23. BISHOP, D.G, BRADSHAW, J.D. & LANDIS, C.A. 1985. Provisional terrane map of South Island, New Zealand. In: HOWELL, D.G. (ed.) Tectonostratigraphic Terranes. Circum-Pacific Council for Energy and Mineral Resources, Earth Science Series No. 1, Houston, Texas, 515-521. BLACK, P.M. 1994. The 'Waipapa terrane', North Island, New Zealand: subdivision and correlation. Geoscience Reports of Shizuoka University, 20, 55-62. BRADSHAW, J.D. 1993. A review of the Median Tectonic Zone: terrane boundaries and terrane amalgamation near the Median Tectonic Line. New Zealand Journal of Geology and Geophysics, 36, 117-125. BRADSHAW, J.D., ANDREWS, P.B. & FIELD, B.D. 1983. Swanson Formation and related rocks of Marie Byrd Land and a comparison with the Robertson Bay Group of Northern Victoria Land. In: OLIVER, R.L., JAMES, PR. & JAGO, J.B. (eds) Antarctic Earth Science. Australian Academy of Science, Canberra, 176-189. BRADSHAW, J.D., PANKHURST, R.J., WEAVER, S.D., STOREY, B.C., MUIR, RJ. & IRELAND, T.R. 1997. New Zealand superterranes recognized in Marie Byrd Land and Thurston Island. In: RICCI, C.A. (ed.) The Antarctic Region: Geological Evolution and Processes. Terra Antartica Publication, Siena, 429-436. BRADSHAW, M.A. 1995. Stratigraphy and structure of the Lower Devonian rocks of the Waitahu and

137

Orlando Outliers, near Reefton, New Zealand, and their relationship to the Inangahua Outlier. New Zealand Journal of Geology and Geophysics, 38, 81-92. BRADSHAW, M.A. 2000. Base of the Devonian Baton Formation and the question of a pre-Baton tectonic event in the Takaka terrane, New Zealand. New Zealand Journal of Geology and Geophysics, 43, 601-610. CAMPBELL, H.J., SMALE, D., GRAPES, R., HOKE, L., GIBSON, G.M. & LANDIS, C.A. 1998. Parapara Group: Permian-Triassic rocks in the Western Province, New Zealand. New Zealand Journal Geology and Geophysics, 41, 281-296. CAMPBELL, J.D. & COOMBS, D.S. 1966. Murihiku Supergroup (Triassic-Jurassic) of Southland and South Otago. New Zealand Journal of Geology and Geophysics, 9, 393-398. CAS, R.A.F. 1983. A review of the paleogeographic and tectonic development of the Paleozoic Lachlan Fold Belt of SE Australia. Geological Society of Australia, Special Publications, 10. CAWOOD, PA. 1986. Stratigraphic and structural relations of the southern Dun Mountain Ophiolite Belt, and enclosing strata, northwestern Southland, New Zealand. New Zealand Journal of Geology and Geophysics, 29,179-203. CAWOOD, P.A. 1987. Stratigraphic and structural relations of strata enclosing the Dun Mountain Ophiolite Belt in the Arthurton-Clinton region, Southland, New Zealand. New Zealand Journal of Geology and Geophysics, 30, 19-36. CAWOOD, PA., NEMCHIN, A. A., LEVERENZ, A., SAEED, A. & BALANCE, PF. 1999. U/Pb dating of detrital zircons: implications for the provenance record of Gondwana margin terranes. Geological Society of America Bulletin, 111, 1107-1119. CAWOOD, P.A., LANDIS, C.A., NEMCHIN, A.A. & HADA, S. 2002. Permian fragmentation, accretion and subsequent translation of a low-latitude Tethyan seamount to the high-latitude east Gondwana margin: evidence from detrital zircon age data. Geological Magazine, 139, 131-144. CONEY, P.J., EDWARDS, A., HINE, R., MORRISON, F. & WINDRUM, D. 1990. The regional tectonics of the Tasman erogenic system, eastern Australia. Journal of Structural Geology, 125, 19-43. COOMBS, D.S., LANDIS, C.A., NORRIS, R.J., SINTON, J.M., BORNS, DJ. & CRAW, D. 1976. The Dun Mountain Ophiolite Belt, New Zealand, its tectonic setting, constitution, and origin, with special reference to the southern portion. American Journal of Science, 276, 561-603. COOPER, R.A. 1979. Lower Paleozoic rocks of the New Zealand region. Journal of the Royal Society of New Zealand, 9, 29-84. COOPER, R.A. 1989. Early Paleozoic terranes of New Zealand. Journal of the Royal Society of New Zealand, 19, 73-112. COOPER, R.A. & TULLOCH, A.J. 1992. Early Paleozoic terranes in New Zealand and their relationship to the Lachlan Fold Belt. Tectonophysics, 214, 129-144. CORBETT, K.D., SOLOMON, M., MCCLENAGHAN, M.P.

138

C. J. ADAMS ETAL.

ET AL. 1989. Cambrian Mt Read Volcanics and associated mineral deposits. In: BURRETT, C.F. & MARTIN, E.L. (eds) Geology and Mineral Resources of Tasmania. Geological Society of Australia, Special Publications, 15, 84-153. CRAW, D. 1984. Lithologic variations in Otago schist, Mt. Aspiring area, northwest Otago, New Zealand. New Zealand Journal of Geology and Geophysics, 27, 151-166. CULLEN, D.G. 1975. Autochthonous rocks of the Bounty Islands region, southwest Pacific Ocean. New Zealand Journal of Geology and Geophysics, 18, 767-785. DALZIEL, I.W.D., GARRETT, S.W., GRUNOW, A.M., STOREY, B.C., PANKHURST, R.J. & VENNUM, W.R. 1987. The Ellsworth-Whitmore crustal block: its role in the tectonic evolution of West Antarctica. In: MACKENZIE, G. (ed.) Gondwana Six: Structure, Tectonics and Geophysics. American Geophysical Union, Monographs, 40,173-182. DEPAOLO, D.J. 1981. Neodymium isotopes in the Colorado Front Range and crust-mantle evolution in the Proterozoic. Nature, 291, 193-196. DEPAOLO, D.J., LINN, A.M. & SCHUBERT, G. 1991. The continental crustal age distribution; methods of determining mantle separation ages from Sm-Nd isotopic data and application to the Southwestern United States. Journal of Geophysical Research, 96, 2071-2088. FANG, Z., BOUCOT, A., COVACEVICH, V. & HERVE, F 1998. Discovery of Late Triassic fossils in the Chonos Metamorphic Complex, southern Chile. Revista Geologica de Chile, 26, 165-174. FERGUSSON, C.L. 2003. Ordovician-Silurian accretion tectonics of the Lachlan Fold Belt, southeastern Australia. Australian Journal of Earth Sciences, 50, 475^90. FERGUSSON, C.L. & TYE, S.C. 1999. Provenance of early Paleozoic sandstones, southeastern Australia, part 1: vertical changes through the Bengal fantype deposit. Sedimentary Geology, 125,135-151. FERGUSSON, C.L. & FANNING, C.M. 2002. Late Ordovician stratigraphy, zircon provenance and tectonics, Lachlan Fold Belt, southeastern Australia. Australian Journal of Earth Sciences, 49, 423^36. FERGUSSON, C.L. & VANDENBERG, A.H.M. 2003. Ordovician - the development of craton-derived deep-sea turbidite successions. In: BIRCH, W.D. (ed.) Geology of Victoria. Geological Society of Australia, Special Publications, 23, 96-115. FORSYTHE, R.D. & MPODOZIS, C. 1979. El archipielago Madre de Dios, Patagonia Occidental, Magallanes: rasgos generates de la estratigrafia y estructura del basamento pre-Jurasico Superior. Revista Geologica de Chile, 1, 13-29. FOSTER, D.A. & GRAY, D.R. 2000. Evolution and structure of the Lachlan Fold Belt (Orogen) of eastern Australia. Annual Reviews of Earth and Planetary Sciences, 28, 47-80. FROST, C.D. & COOMBS, D.S. 1989. Nd isotopic character of New Zealand sediments: implications for terrane concepts and crustal evolution. American Journal of Science, 289, 744-770. GLEN, R.A., DALLMEYER, R.D. & BLACK, L.P. 1992.

Isotopic dating basin inversion - the Paleozoic Cobar Basin, Lachlan Orogen, Australia. Tectonophysics, 214, 249-268. GRAY, C.M. & WEBB, J.A. 1995. Provenance of Paleozoic turbidites in the Lachlan Orogenic Belt: strontium isotopic evidence. Australian Journal of Earth Sciences, 42, 95-105. HERVE, F 1993. Paleozoic metamorphic complexes in the Andes of Aysen, southern Chile (West of Occidentalia). In: ORTEGA-GUTIERREZ, F., CENTENO-GARCIA, E., MORAN-ZENTENO, D.J. & GOMEZ-CABALLERO, A. (eds) Proceedings, First Circum-Pacific and Circum-Atlantic Terrane Conference, Guanajuato. Institute de Geologia, Universidad Autonoma de Mexico, Mexico City, 64-65. HERVE, F. & FANNING, C.M. 2003. Early Cretaceous subduction of continental crust at the Diego de Almagro archipelago, southern Chile. Episodes, 26, 285-289. HERVE, F, FANNING, C.M. & PANKHURST, R.J. 2003. Detrital zircon age patterns and provenance of the metamorphic complexes of Southern Chile. Journal of South American Earth Sciences, 16, 107-123. HOUGHTON, B.F. 1981. Lithostratigraphy of the Takitimu Group, central Takitimu Mountains, western Southland, New Zealand. New Zealand Journal of Geology and Geophysics, 24, 333-348. HOUGHTON, B.F. & LANDIS, C.A. 1989. Sedimentation and volcanism in a Permian arc-related basin, southern New Zealand. Bulletin of Volcanology, 51, 433-450. HYDEN, G. & TANNER, P.W.G. 1981. Late Palaeozoic-Early Mesozoic fore-arc basin sedimentary rocks at the Pacific margin in western Antarctica. Geologische Rundschau, 70, 529-541. IRELAND, T.R. 19920. Crustal evolution of New Zealand: evidence from age distributions of zircons in Western Province paragneisses and Torlesse greywacke. Geochimica et Cosmochimica Acta, 56, 911-920. IRELAND, T.R. 1992/?. Correlation and provenance of sedimentary units around the southwestern Pacific margin. Geological Society of New Zealand, 1992 Annual Conference Programme and Abstracts. Geological Society of New Zealand Miscellaneous Publications 63A, 84. IRELAND, T.R., FLOTTMANN, T, FANNING, C.M., GIBSON, G.M. & PREISS, WV. 1998. Development of the early Palaeozoic Pacific margin of Gondwana from detrital-zircon ages across the Delamerian orogen. Geology, 26, 243-246. JOHNSTON, M.R. 1981. Sheet O27AC - Dun Mountain. Geological Map of New Zealand 1:50000. Department of Scientific Research, Wellington. LAIRD, M.G. & SHELLEY, D. 1975. Sedimentation and early tectonic history of the Greenland Group, Reefton, New Zealand. New Zealand Journal of Geology and Geophysics, 17, 839-854. LANDIS, C.A. 1974. Paleocurrents and compositions of lower Bryneira strata (Permian) at Barrington Peak, northwest Otago. New Zealand Journal of Geology and Geophysics, 17, 799-806. LANDIS, C.A. 1980. Little Ben Sandstone, Maitai

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN Group (Permian): nature and extent in the Hollyford-Eglinton region, South Island, New Zealand. New Zealand Journal of Geology and Geophysics, 23, 551-567. LANDIS, C.A., CAMPBELL, H.J., ASLUND, T, CAWOOD, PA. ET AL. 1999. Permian-Jurassic strata at Productus Creek, Southland, New Zealand: implications for terrane dynamics at the eastern Gondwanaland margin. New Zealand Journal of Geology and Geophysics, 42, 255-278. LEITCH, E.C., FERGUSSON, C.L. & HENDERSON, R.A. 2003. Arc to pluton switching in a Late Paleozoic subduction complex, Wandilla and Shoalweater terranes, New England Fold Belt, eastern Australia. Australian Journal of Earth Sciences, 50, 919-929. LIEW, T.C. & HOFMANN, A.W. 1988. Precambrian crustal components, plutonic associations, plate environment of the Hercynian Fold Belt of central Europe: indications from a Nd and Sr isotopic study. Contributions to Mineralogy and Petrology, 98, 129-138. LING, H.Y., FORSYTHE, R.D. & DOUGLAS, C.R. 1987. Late Paleozoic microfaunas from southernmost Chile and their relation to Gondwanaland forearc development. Geology, 13, 357-360. MAAS, R., NICHOLLS, LA. & LEGG, C. 1997. Igneous and metamorphic enclaves in the S-type Deddick Granodiorite, Lachlan Fold Belt, SE Australia: petrographic, geochemical and Nd-Sr isotopic evidence for crustal melting and magma mixing. Journal of Petrology, 38, 815-841. MACKINNON, T.C. 1983. Origin of the Torlesse terrane and coeval rocks, South Island, New Zealand. Geological of Society of American Bulletin, 94, 967-985. MCCULLOCH, M.T. & CHAPPELL, B.W. 1982. Nd isotopic characteristics of S- and I-type granites. Earth and Planetary Science Letters, 58, 51-64. MCLENNAN, S.M., MCCULLOCH, M.T, TAYLOR, S.R. & MAYNARD, IB. 1989. Effects of sedimentary sorting on Nd-isotopes in deep-sea turbidites. Nature, 337, 547-549. MILLAR, I.L. & PANKHURST, R.J. 1987. Rb-Sr geochronology of the region between the Antarctic Peninsula and the Transantarctic Mountains: Haag Nunataks and Mesozoic Granitoids. In: MACKENZIE, G. (ed.) Gondwana Six: Structure, Tectonics and Geophysics. American Geophysical Union, Monographs, 40, 151-160. MILLAR, I.L., PANKHURST, R.J., VAUGHAN, A.P.M. & FANNING, C.M. 2001. Exotic crustal fragments on the Pacific margin of Gondwana. Third South American Symposium of Isotope Geology, Pucon, Chile, 21-24 October. CD-ROM, SERNAGEOMIN, Santiago, 589-592. MILLAR, I.L., PANKHURST, R.J. & FANNING, C.M. 2002. Basement Chronology of the Antarctic Peninsula: recurrent magmatism and anatexis in the Palaeozoic Gondwana Margin. Journal of the Geological Society, London, 159, 145-157. MILLAR, I.L., VAUGHAN A.P.M., HORSTWOOD, M.S.A., FLOWERDEW, M., PANKHURST, R.J., PASHLEY, V.H. & FANNING C.M. 2003. Hf isotopes in detrital zircons from accretionary complex rocks of

139

Alexander Island, Antarctic Peninsula. Geophysical Research Abstracts, 5, 09257. MORTIMER, N. 19930. Geology of the Otago Schist and Adjacent Rocks. Scale 1:500000. Institute of Geological & Nuclear Sciences geological map 7. Institute of Geological and Nuclear Sciences, Lower Hutt. MORTIMER, N. 1993b. Metamorphic zones, terranes and Cenozoic faults in Marlborough Schist, New Zealand. New Zealand Journal of Geology and Geophysics, 36, 357-368. MORTIMER, N., TULLOCH, A.J. & IRELAND, T.R. 1997. Basement geology of Taranaki and Wanganui Basins, New Zealand. New Zealand Journal of Geology and Geophysics, 40, 223-268. MORTIMER, N., TULLOCH, A.I, SPARK, R.N., WALKER, N. W., LADLEY, E., ALLIBONE, A. & KIMBROUGH, D.L. 19990. Overview of the Median Batholith, New Zealand: a new interpretation of the geology of the Median Tectonic Zone and adjacent rocks. Journal of African Earth Sciences, 29, 257-268. MORTIMER, N., TULLOCH, A.J., GANS, P., CALVERT, A. & WALKER, N. 19996. Geology and thermochronometry of the east edge of the Median Tectonic Batholith (Median Tectonic Zone): a new perspective on Permian to Cretaceous crustal growth of New Zealand. The Island Arc, 8, 404-425. MUNKER, C. & COOPER, R.A. 1999. The Cambrian arc complex of the Takaka terrane, New Zealand: integrated stratigraphical, paleontological and geochemical approach. New Zealand Journal of Geology and Geophysics, 42, 415-455. NATHAN, S. 1976. Geochemistry of the Greenland Group (early Ordovician), New Zealand. New Zealand Journal of Geology and Geophysics, 19, 683-706. NORRIS, R.J. & CRAW, D. 1987. Aspiring terrane: an oceanic assemblage from New Zealand and its implications for Mesozoic accretion in the southwest Pacific. In: LEITCH, E.G. & SCHEIBNER, E. (eds) Terrane Accretion and Orogenic Belts. American Geophysical Union, Geodynamics Series, 19,169-177. OFFLER, R., MCKNIGHT, S. & MORAND, V. 1998. Tectonothermal history of the western Lachlan Fold Belt, Australia - insights from white mica studies. Journal of Metamorphic Geology, 16, 531-540. O'HALLORAN, G.J. 1996. The sedimentological, volcanological and stratigraphic evolution of the Mt. Howitt Province, east-central Victoria: implications for tectonic settings in the Late-DevonianEarly Carboniferous Lachlan Fold Belt. PhD thesis, Monash University, Melbourne. PANKHURST, R.J., STOREY, B.C., MILLAR, I.L., MACDONALD, D.I.M. & VENNUM, W.R. 1989. Cambrian-Ordovician magmatism in the Thiel Mountains, Transantarctic Mountains, and implications for the Beardmore Orogeny. Geology, 16, 246-249. PANKHURST, R.J., RAPELA, C.W., CAMINOS, R., LLAMBIAS, E.J. & PARICA, C. 1992. A revised age for the granites of the central Somuncura

140

C. J. ADAMS ET AL.

batholith, North Patagonian Massif. Journal of South American Earth Sciences, 5, 321-325. PANKHURST, R.J., HERVE, F. & RAPELA, C.W. 1994. Sm-Nd evidence for the Grenvillian provenance of the metasedimentary basement of Southern Chile and West Antarctica. In: CAMPOS, E. & CECIONI, A. (eds) Actas del VII Congreso Geologico Chileno, Vol. II, Universidad del Norte Chile, Departamento de Geociencias, Facultad de Ciencias. Antofagasta, Chile, 1414-1418. PANKHURST, R.J., RAPELA, C.W., SAAVEDRA, X, BALDO, E., DAHLQUIST, X, PASCUA, I. & FANNING, C.M. 19980. The Famatinian magmatic arc in the southern Sierras Pampeanas. In: PANKHURST, R.X & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 343-367. PANKHURST, R.X, WEAVER, S.D., BRADSHAW, J.D., STOREY, B.C. & IRELAND, T.R. 1998&. Geochronology and geochemistry of pre-Jurassic superterranes in Marie Byrd Land, Antarctica. Journal of Geophysical Research, 103, 2529-2547. PANKHURST, R.X, RAPELA, C.W & FANNING, C.M. 2001fl. The Mina Gonzalito gneiss: Early Ordovician metamorphism in northern Patagonia. Third South American Symposium of Isotope Geology, Pucon, Chile, 21-24 October. CD-ROM, SERNAGEOMIN, Santiago, 604-607. PANKHURST, R.X, RAPELA, C.W. & FANNING, C.M. 2001&. Age and origin of coeval TTG, I- and Stype granites in the Famatinian belt of NW Argentina. Transactions of the Royal Society of Edinburgh: Earth Sciences, 91,151-168. PANKHURST, R.X, RAPELA, C.W, LOSKE, WP, FANNING, C.M. & MARQUEZ, M. 2003. Chronological study of the pre-Permian basement rocks of southern Patagonia. Journal of South American Earth Sciences, 16, 27-44. PICKARD, A.L., ADAMS, C.X & BARLEY, M.E. 2000. Australian provenances for Upper Permian to Cretaceous rocks forming accretionary complexes on the New Zealand sector of the Gondwanaland margin. Australian Journal of Earth Sciences, 47, 987-1007. PICKETT, D.A. & WASSERBURG, G.X 1989. Neodymium and strontium isotopic characteristics of New Zealand granitoids and related rocks. Contributions to Mineralogy and Petrology, 103, 131-142. PIMPIREV, CHR., DIMOV, D. & IVANOV, M. 2003. The Miers Bluff Formation, Livingston Island, South Shetland Islands - part of the Upper JurassicCretaceous depositional history of the Antarctic Peninsula. 10th International Symposium of Antarctic Earth Sciences (Potsdam), Terra Nostra Schriften der Alfred-Wegener Stiftung, Berlin, 256-257. POUND, K., GRAY, D.R. & CAS, R.A.F 1994. Provenance, sedimentologic and stratigraphic relationships of Ordovician sandstones of the Bendigo-Ballarat zone, Victoria: implications for the tectonosedimentary evolution of the Lachlan Fold Belt. Abstracts of the Geological Society of Australia, 37, 353-354.

RAPELA, C.W, PANKHURST, R.X, CASQUET, C., BALDO, E., SAAVEDRA, X, GALINDO, C. & FANNING, C.M. 1998. The Pampean orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de Cordoba. In: PANKHURST, R.X & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 181-217. RATTENBURY, M.S., COOPER, R.A. & JOHNSTON, M.R. (compilers) 1998. Geology of the Nelson Area. Scale 1:250000. Institute of Geological & Nuclear Sciences 1:250 000 geological map 9. Institute of Geological & Nuclear Sciences, Lower Hutt. RICHARD, P., SHIMIZU, N. & ALLEGRE, C.X 1976. 143 Nd/144Nd, a natural tracer: an application to oceanic basalts. Earth and Planetary Science Letters, 31, 269-278. ROSER, B.P. & KORSCH, R.X 1988. Provenance signatures of sandstone-mudstone suites determined using discriminant function analysis of major element data. Chemical Geology, 67,119-139. ROSER, B.P. & KORSCH, R.X 1999. Geochemical characterisation, evolution and source of a Mesozoic accretionary wedge, the Torlesse terrane, New Zealand. Geological Magazine, 136, 493-512. ROSER, B.P, COOPER, R.A., NATHAN, S. & TULLOCH, A.X 1996. Reconnaissance sandstone geochemistry, provenance, and tectonic setting of the lower Paleozoic terranes of the West Coast and Nelson, New Zealand. New Zealand Journal of Geology and Geophysics, 39,1-16. SPORLI, K.B. 1978. Mesozoic tectonics, North Island, New Zealand. Geological Society of America Bulletin, 89, 412-425. STEIGER, R.H. & JAGER, E. 1977. Subcommission on geochronology: convention on the use of decay constants in geo- and cosmochronology. Earth and Planetary Science Letters, 36, 359-362. STUMP, E., WHITE, A.XR. & BORG, S.G. 1986. Reconstruction of Australia and Antarctica: evidence from granites and recent mapping. Earth and Planetary Science Letters, 79, 348-360. SUTHERLAND, R. 1999. Basement geology and tectonic development of the greater New Zealand region: an interpretation from regional magnetic data. Tectonophysics, 308, 314-362. TROUW, R.A.X, PANKHURST, R.X & KAWASHITA, K. 1990. New radiometric age data from Elephant Island, South Shetland Islands. Zentralblatt fur Geologie und Paldontologie, Teil I, 1990 (1/2), 105-118. TURNBULL,I.M. 19790. Stratigraphy and sedimentology of the Caples terrane of the Thomson Mountains, northern Southland, New Zealand. New Zealand Journal of Geology and Geophysics, 22, 555-574. TURNBULL, I.M. 19796. Petrography of the Caples terrane of the Thomson Mountains, northern Southland, New Zealand. New Zealand Journal of Geology and Geophysics, 22, 709-727. TURNER, S., FODEN, X, SANDIFORD, M. & BRUCE, D. 1993. Sm-Nd isotopic evidence for the provenance of sediments from the Adelaide Fold Belt and southeastern Australia with implications

Nd AND Sr ISOTOPES ON THE PACIFIC MARGIN for episodic crustal addition. Geochimica et Cosmochimica Acta, 57, 1837-1856. TURNER, S.P., KELLEY, S.P., VANDENBERG, A.H.M., FODEN, J.D., SANDIFORD, M. & FLOTTMANN, T. 1996. Source of the Lachlan fold belt flysch linked to convective removal of the lithospheric mantle and rapid exhumation of the Delamerian-Ross Fold Belt. Geology, 24, 941-944. VANDENBERG, A.H.M., WILLMAN, C.E., MAHER, S. ET AL. 2000. The Tasman Fold Belt System in Victoria. Geological Survey of Victoria, Melbourne, Special Publication. VAUGHAN, A.P.M. & STOREY, B.C. 2000. The eastern Palmer Land shear zone: a new terrane accretion model for the Mesozoic development of the Antarctic Peninsula. Journal of the Geological Society, London, 157, 1243-1256. VEEVERS, J.J. 2000. Antarctic Beardmore-Ross and Mirny provenances saturate Palaeozoic-Mesozoic East Gondwanaland with 0.6-0.5 Ga zircons. In: VEEVERS, J.J. (ed.) Billion-year Earth History of Australia and Neighbours in Gondwanaland. GEMOC Press, Sydney, 110-130. VEEVERS, J.J. & POWELL, C. McA. (eds) 1994. PermianTriassic basins andfoldbelts along the Panthalassan margin of Gondwanaland. Geological Society of America, Memoirs, 184. WAIGHT, T.E., WEAVER, S.D., MUIR, R.J., MAAS, R. & EBY, G.N. 1998. The Hohonu Batholith of North Westland, New Zealand: granitoid compositions controlled by source H2O contents and generated

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during tectonic transition. Contributions to Mineralogy and Petrology, 130, 225-239. WEVER, H.E., MILLAR, I.L. & PANKHURST, R.J. 1994. Geochronology and radiogenic isotope geology of Mesozoic rocks from eastern Palmer Land, Antarctic Peninsula; crustal anatexis in arcrelated granitoid genesis. Journal of South American Earth Sciences, 7, 69-83. WILLAN, R.C.R., PANKHURST, R.J. & HERVE, F. 1994. A probable Early Triassic age for the Miers Bluff Formation, Livingston Island, South Shetland Islands. Antarctic Science, 6, 401-408. WILLIAMS, I.S., CHAPPELL, B.W., CROOK, K.A.W. & NICHOLL, R.S. 1994. In search of the provenance of the early Palaeozoic flysch in the Lachlan Fold Belt, southeastern Australia. Abstracts of the Geological Society of Australia, 37, 464. WOMBACHER, F. & MuNKER, C. 2000. Pb, Nd, and Sr isotopes and REE systematics of Cambrian sediments from New Zealand: implications for the reconstruction of the early Paleozoic Gondwana margin along Australia and Antarctica. Journal of Geology, 108, 663-686. YOUNG, G.C. & LAURIE, J.R. (eds) 1998. An Australian Phanerozoic Timescale. Oxford University Press, Melbourne. ZIMMERMANN, U. & BAHLBURG, H. 2003. Provenance analysis and tectonic setting of the Ordovician clastic deposits in the southern Puna Basin, NW Argentina. Sedimentology, 50,1-26.

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Episodicity of Mesozoic terrane accretion along the Pacific margin of Gondwana: implications for superplume-plate interactions A. P. M. VAUGHAN & R. A. LIVERMORE British Antarctic Survey, High Cross, Madingley Rd, Cambridge CBS GET, UK (e-mail: [email protected]) Abstract: A review of evidence for deformation and terrane accretion on the Late Triassic-Early Jurassic margins of Pangaea and the mid-Cretaceous margins of the palaeoPacific ocean shows that deformation was global and synchronous with probable superplume events. Late Triassic-Early Jurassic deformation appears to be concentrated in the period 202-197 Ma and was coeval with eruption of the Central Atlantic Magmatic Province, onset of Pangaea break-up, a period of extended normal magnetic polarity and a major mass extinction event, all possible expressions of a superplume event. MidCretaceous deformation occurred in two brief periods, the first from approximately 116 Ma to 110 Ma in the west palaeo-Pacific and the second from roughly 105 Ma to 99 Ma in the east palaeo-Pacific, with both events possibly represented in northeast Siberia. This deformation was coeval with eruption of major oceanic plateaux, core-complex formation and rifting of New Zealand from Gondwana, the Cretaceous normal polarity epoch, and a major radiation of flowering plants and several animal groups, all linked with the midCretaceous superplume event. A simple unifying mechanism is presented suggesting that large continental or oceanic plates, when impacted by a superplume, tend to break-up/reorganize, associated with gravitational spreading away from a broad, thermally generated topographic high and with a resulting short-lived pulse of plate-marginal deformation and terrane accretion.

Long-term episodicity of geological activity, i.e. variations of intensity or style with time but not necessarily to any particular periodicity, has been proposed for many processes, including aspects of magmatism (Prokoph et al. 2004) and tectonism (Condie 1998), from the Early Archaean to the present era. Generally, the origin of episodicity has been attributed to mantle thermal behaviour. This is either as gradual changes associated with the growth and fragmentation of supercontinents, the so-called Wilson Cycle, as a result of processes such as thermal shielding (Condie 2002), or more catastrophic, non-linear events, referred to as superplumes (Larson 19915; Condie 1998; Abbott & Isley 2002) or mantle avalanches (Solheim & Peltier 1994; Butler & Peltier 1997; Machetel 2003). This paper reviews the evidence for the timing and nature of two episodes of Mesozoic terrane accretion and deformation (summarized in Tables 1 and 2). It is proposed that these were superplume-driven and a unifying mechanism and conceptual model is presented for the interaction between large plates and superplume arrival from depth, focusing on deformation on the Gondwana/Pangaea margin during the Late Triassic-Early Jurassic and deformation on the margin of the palaeo-Pacific basin during the mid-Cretaceous. Both mid-Cretaceous and Late

Triassic-Early Jurassic deformation events were temporally associated with a short-term global increase in the rate of ophiolite obduction (Vaughan & Scarrow 2003) and were associated with major, concentrated magmatic episodes (e.g. Larson 1997; Riley et al 2004), changes in geomagnetic field behaviour (Biggin & Thomas 2003) and large-scale environmental perturbation (Hesselbo et al 2002; Larson & Erba 1999). Time and timing is obviously very important, particularly when relating event chronologies determined from the stratigraphic record with those determined radiometrically. For consistency, all chrononstratigraphic series and stage range and boundary ages are taken from the International Commission on Stratigraphy time-scale published by Gradstein et al (2004). The two enigmatic orogenic events under consideration here are well-represented in rocks on the Pacific margin of Gondwana during the Mesozoic. The Late Triassic-Early Jurassic event is called the Peninsula Orogeny in West Antarctica (Miller 1983) and the Rangitata I Orogeny in New Zealand (Bradshaw et al 1981); the mid-Cretaceous event is called the Palmer Land event in the Antarctic Peninsula region (Kellogg & Rowley 1989; redated by Vaughan et al 20025) and the Rangitata II Orogeny in New Zealand (Bradshaw et al

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246,143-178. 0305-8719/$15.00 © The Geological Society of London 2005.

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Table 1. Triassic-Jurassic events Location Antarctica South Orkney Islands

Antarctic Peninsula

Pensacola Mountains

South America Patagonia

North Chilean Coastal Cordillera

Central and North America SW Mexico

Timing

Event Deformation and metamorphism of Scotia complex rocks Peninsula Orogeny affecting Alexander Island and Trinity Peninsula accretionary rocks D3 phase of deformation in Patuxent Range

Post-Triassic and pre-Early to Interval between youngest Middle Jurassic (200-176 Ma) deformed rocks and oldest unconformable sediments Post-Triassic and pre-Middle Interval between youngest Jurassic deformed rocks and oldest unconformable sediments Post-Early Triassic and Interval between end of pre-183 Ma Permo-Triassic Gondwanian deformation and intrusion of lamprophyre dyke

Deformation and 234-198 Ma metamorphism of Madre de Dios and Chonos accretionary complex rocks Transpressional deformation 200-174 Ma of Bolfin Complex amphibolites

Post-early Permian and pre-Middle Jurassic

Strike-slip and folding of Palaeozoic sediments

Late Triassic

Gobi Basin, Mongolia

Deformation of Late Triassic sediments

209-206 Ma

Liaoning, North China

Oblique ductile thrusting of paragneiss

c. 219 Ma

Nevada, USA

Nevada, USA

Canadian Cordillera

SE British Columbia

E Central Alaska

Southern Yukon Territory

SE Alaska Eurasia Taimyr Peninsula

Combined zircon fission track and U-Pb method dating

U-Pb zircon method dating of syn-deformational igneous protoliths (200-191 Ma) and post deformational granite (174 Ma)

Interval between youngest deformed rocks and oldest unconformable sediments, and geometric correlation of structures Thrusting of submarine fan Post-Norian and pre-Aalenian Interval between youngest deposits, Guerrero and Sierra (203-175 Ma) deformed rocks and oldest Madre terranes unconformable sediments Interval between youngest Overthrusting of marine Post-mid-Triassic and pre-Late Triassic-Early deformed rocks and clastic rocks by ophiolitic Jurassic melange post-deformational felsic dykes (222-192 Ma) Ar/Ar dating of deformation Activity on 201-193 Ma Luning-Fencemaker Thrust structures Belt Imbrication of Cache Creek 210-173 Ma Interval between eruption of basaltic lavas (210 Ma) and terrane intrusion of cross-cutting pluton (173 Ma) 215-197 Ma Interval between intrusion of Deformation of Quesnel terrane subsequently deformed diorite (215 Ma) and intrusion of cross-cutting plutons (197-181 Ma) Imbrication of arc, subduction Late Triassic to Early Jurassic Interval between intrusion of subsequently deformed granite complex and continental (212-188 Ma) (212 Ma) and Ar/Ar cooling rocks ages in sheared rocks (188-185 Ma) Yukon-Tanana terrane ductile Late Triassic to Early Jurassic Interval between youngest deformed rocks and oldest thrust deformation unconformable sediments Ar/Ar dating of deformation Blueschist metamorphism of c. 195 Ma structures McHugh Complex Thrusting and folding of sedimentary rocks, Oaxaca terrane

Mesa Central, Mexico

Dating technique/criteria

Interval between youngest deformed rocks and oldest unconformable sediments Interval between youngest deformed rocks and oldest unconformable sediments Ar/Ar dating of deformation fabrics

EPISODICITY OF TERRANE ACCRETION

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Table 1. (continued) Location Sichuan, west-central China Longmen Shan, west China Fujian, SE China

Peninsular Malaysia

Vietnam/Laos

Kunlun Mountains, SW China Central Iran

Ukraine

Northern Turkey

New Zealand South Island

Event

Timing

Deformation on Songpan-Garz fold-thrust belt Foreland basin deposition Amphibolite facies metamorphism of arc rocks and terrane collision Block collision

c. 197 Ma

Terrane collision and metamorphism

225-170 Ma

Dating technique/criteria U-Pb and Rb-Sr geochronology

Depositional ages of sediments 227-206 Ma Late Triassic-Middle Jurassic Interval between youngest deformed rocks and oldest unconformable sediments Late Triassic-Early Jurassic Interval between youngest deformed rocks and oldest unconformable sediments Interval between youngest Closure of Permo-Triassic rift Middle Triassic-Latest basins along Song-Da zone Triassic deformed rocks and oldest unconformable sediments Deformation of Mazar 215-190 Ma Syn- and post-deformational accretionary prism and plutons terrane collision Deformation of sediments and Late Triassic Interval between youngest widespread unconformity deformed rocks and oldest development unconformable sediments Deformation in Donbas fold Latest Triassic Interval between youngest belt and unconformity deformed rocks and oldest development unconformable sediments Cimmeride deformation 215-197 Ma Blueschist Ar/Ar ages and age of oldest unconformable sediments

1981). The global extent of deformation associated with these events is illustrated below. The Late Triassic-Early Jurassic event is developed widely in accretionary complex rocks that now sit outboard of rocks characteristic of continental Gondwana, and recently has been identified tentatively in Lower Palaeozoic rocks of the Pensacola Mountains (Curtis 2002). The deformation episode, which is found in Gondwanamargin rocks from New Zealand, Antarctica and South America, is particularly significant because it appears to have been active just prior to, or during, the earliest stages of Gondwana break-up. Suggested causes for the mountain-building event have varied widely, from Late Permian to mid-Jurassic dextral transpression and terrane collision (Martin et al. 1999), earliest Jurassic to early mid-Jurassic sinistral transpression (Scheuber & Gonzalez 1999), thermal uplift associated with the breakup of Gondwana (Dalziel 1982), a terraneaccretionary complex collision event (Bradshaw et al 1981), or episodic accretionary complex activity (Thomson & Herve 2002). Aptian to Albian mid-Cretaceous tectonism widely affected rocks around the margin of the Pacific basin (Vaughan 1995) (Fig. 2), and is welldeveloped in all terranes of the Antarctic

Metamorphic cooling ages and age of oldest unconformable sediments

Peninsula where it is reasonably well dated to have occurred between c. 107 Ma and c. 103 Ma (Vaughan et al 20020, 2002ft). Distribution of Late Triassic-Early Jurassic deformation in Pangaea

Antarctica The Peninsula Orogeny (Miller 1983; reassessed by Storey et al 1987) of West Antarctica was recognized to be an orogenic event that postdated the Late Palaeozoic, Gondwanian Orogeny (Curtis 2001), affecting Late Palaeozoic and Mesozoic accretionary complex rocks such as the LeMay Group of Alexander Island (Al: Fig. 1) (Burn 1984; Tranter 1987) and Trinity Peninsula Group of Graham Land (TP: Fig. 1) (Aitkenhead 1975; Smellie 1981). Deformed accretionary complex rocks of the Western Domain (Vaughan & Storey 2000), LeMay Group in eastern Alexander Island are overlain by Bathonian (168-165 Ma) sediments of the Fossil Bluff Group (Doubleday et al 1993). Dated LeMay Group rocks range in age from mid-Cretaceous in the west (Holdsworth & Nell 1992) to Sinemurian (197-190 Ma) in the east (Thomson & Tranter 1986) and possibly

A. P. M. VAUGHAN & R. A. LIVERMORE

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Table 2. Mid-Cretaceous events Location Antarctica Elephant Island Antarctic Peninsula

Marie Byrd Land

East Antarctica

South America South Georgia

Cordillera Darwin

Western South America

Event

Blueschist Ar/Ar ages

107-103 Ma

U-Pb and Ar/Ar dating of deformation structures and syn-deformational dyke

Post c. 146 Ma and pre-105-94 Ma extensional mylonite formation Possibly mid-Cretaceous

Dyke emplacment timing and structural relationships

Thrusting of ophiolitic rocks

Aptian-Albian (125-99 Ma)

Amphibolite facies metamorphism and back-arc basin closure 'Mirano diastrophic phase'

110-90 Ma

Interval between youngest deformed rocks and oldest unconformable sediments K-Ar dating of metamorphism

Sinistral transpression of Jurassic volcanic rocks in Lanterman Range

Shoaling of basins and unconformity development

Northern Chile

Inversion of Atacama Fault zone

Western Peru

Mochica deformation event

Western Columbia

Uplift and erosion of mid-Albian sandstones

Northern Venezuela

Transpressional deformation on the Caribbean-South

W Guatemala Honduras

Southwestern Mexico

North America Southern California Central California

Dating technique/criteria

110-90 Ma

Strike-slip shear on plate boundary Palmer Land event and general uplift affecting arc terranes and Gondwana-margin rocks Dextral shear of crystalline basement rocks

Western South America

Central America and the Carribean Hispaniola

Timing

Regional correlation

Mid-Cretaceous (112-105 Ma) Interval between youngest deformed rocks and oldest unconformable sediments Interval between youngest Post c. Ill Ma and pre-Cenomanian deformed rocks and oldest unconformable sediments; dating of tuff bed Postc. Ill Ma Interval between youngest deformed rocks and oldest unconformable sediments Late Albian (c. 100 Ma) Interval between youngest deformed rocks and oldest unconformable sediments Late Albian (c. 100 Ma) Interval between youngest deformed rocks and oldest unconformable sediments White mica and amphibole c. 96 Ma Ar/Ar ages from high pressure American plate boundary metamorphic rocks

Interval between youngest deformed rocks and oldest unconformable sediments Ar/Ar deformation ages from Collision of the Chortis Block 125-113 Ma with western Mexico phengitic micas Uplift and erosion of a Post-mid-Albian and pre-Late Interval between youngest limestone sequence Cretaceous (105-99 Ma) deformed rocks and oldest unconformable sediments Collision of the Guerrero 112-101 Ma Stitching plutons showing greenschist facies terrane metamorphism coeval with deformation

Thrust emplacement of the Hispaniola peridotite belt

Aptian-Albian (125-99 Ma)

Thrusting of the western Peninsular Ranges batholith Thrusting in Sierra Nevada

Probably 115-108 Ma and pre-99 Ma 110-92 Ma with peaks at c. 105 Ma and c. 101 Ma Mid-Cretaceous (c. 105-80 Ma)

Northwestern North America Thrusting of Coast Belt

Ar/Ar dating of deformation structures Ar/Ar dating of deformation structures Interval between youngest deformed rocks and dated cross-cutting plutons

EPISODICITY OF TERRANE ACCRETION

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Table 2. (continued) Location

Event

Timing

Southwest British Columbia

Thrusting of arc rocks

105-97 Ma

SE Alaska

Thrusting of ultramafic and arc rocks on Duke Island

Syn- or Post-c. 108 Ma and pre-c. 95 Ma

Central and northern Alaska

Widespread thrusting of arc terranes

c. 108 Ma

Deformation of arc rocks in Anadyr-Koryak region Ophiolitic thrusting in Sikhote-Alin mountain range Underthrusting and accretion of ophiolitic arc rocks

c. 124-117 Ma and 112-109 Ma c. 110 Ma

East and Southeast Asia Northeast Siberia Eastern Siberia Northern Japan

Central Japan Shikoku, Japan

Metamorphism of South Kitakami and palaeo-Ryoke Sanbagawa metamorphism

Southeastern China

Thrusting of a dismembered ophiolite suite, Fujian

Western China Central Vietnam

Rapid exhumation of arc rocks Metamorphism and mylonite formation Thrusting of ophiolitic rocks during block collision

Western Philippines

Post-c. 125 Ma and pre-c. 110 Ma but possibly to 100 Ma

c.l 16 Ma

120-110 Ma with a peak at 116 Ma 118-107 Ma

115-90 Ma 120-90 Ma with peaks at c. 115 Ma and 106-103 Ma c. 100 Ma

Java, Kalimantan and Sulawesi Uplift and recrystallization of 120-115 Ma high-pressure rocks during block collision Australia East and northeast Australia Southeastern Australia

New Zealand South Island

North Island

Dating technique/criteria Interval between youngest deformed rocks and dated cross-cutting plutons Interval between intrusion of subsequently deformed plutons (111-108 Ma) and regional mid-Cretaceous thrusting Ar/Ar dating of blueschist and amphibolite-facies metamorphism structures Ar/Ar dating of deformation fabrics Ar/Ar dating of deformation fabrics Interval between youngest deformed rocks and oldest unconformable sediments and Ar/Ar dated blueschist metamorphism Ar/Ar dating of metamorphic belts fabrics Ar/Ar dating of deformation fabrics Age of youngest deformed plutons and Ar/Ar dating of deformation fabrics Fission track dating Ar/Ar dating of deformation fabrics Interval between youngest deformed rocks and oldest unconformable sediments Interval between youngest deformed rocks and oldest unconformable sediments and regional correlation

Widespread erosional Post-Late Aptian-pre-Early unconformity development in Albian (c. 112 Ma) major basins Folding of igneous rocks in 120-95 Ma the New England fold belt

Interval between youngest deformed rocks and oldest unconformable sediments Interval between youngest deformed rocks and oldest dated cross-cutting plutons

Contractional deformation of 116-101 Ma and probably arc terranes pre c. 110 Ma

Syn-deformational pluton emplacement and Interval between youngest deformed rocks and oldest unconformable sediments Fission track dating and age of cross-cutting dykes

Deformation and metamorphism of accretionary complex rocks

older (Kelly et al 2001), but, in the absence of detailed field data, deformation can be placed only at pre-Bathonian (c. 168 Ma). PermianTriassic turbiditic rocks of the Trinity Peninsula Group show polyphase deformation at

100-85 Ma and possibly pre-c. 99 Ma

anchizonal and epizonal metamorphic grade (Smellie et al 1996) prior to deposition of the Early to Middle Jurassic non-marine Botany Bay Group (Cantrill 2000). D2 deformation, which probably corresponds to the Peninsula

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A. P. M. VAUGHAN & R. A. LIVERMORE

Fig. 1. Distribution of areas preserving Late Triassic-Early Jurassic deformation and magmatism on 200 Ma reconstruction of the Pangaea supercontinent. Areas of large igneous provinces or inferred oceanic plateaux outlined with dashed line; kimberlite magmatism marked with K; plutonism marked with a star. AI, Alexander Island; CAMP, Central Atlantic Magmatic Province; CC, Coastal Cordillera; CI, Central Iran; C/M, Chonos/Madre de Dios accretionary complexes; CT, Cache Creek terrane; CTP, Cache Creek terrane plateau; DF, Donbas fold belt; FP, Fujian Province; GB, Gobi Basin; GF, Gastre Fault system; CLIP, Gondwana Large Igneous Province; IA, Izmir-Ankara suture; KM, Kunlun Mountains; LP, Liaoning Province; LS, Longmen Shan basin; MC, Mesa Central; MH, McHugh Complex; NN, north-central Nevada; NS, North China-South China block collision zone; NUP, Niliifer Unit plateau; NO, Novaya Zemlya; NZ, New Zealand; OT, Oaxaca terrane; PE, Peninsular Malaysia; PM, Pensacola Mountains; QT, Quesnel terrane; SC, Scotia Complex; SD, Song-Da zone; SG, Songpan-Garz; TA, Taimyr Peninsula; TP, Trinity Peninsula; TZ, Teslin tectonic zone; WN, western Nevada; YT, Yukon-Tanana terrane.

Orogeny, constitutes an episode of marginparallel sinistral transpressional shear (A. P. M. Vaughan unpublished data). Deformation and metamorphism of accretionary complex metapelites and metagreywackes in the Scotia Complex (SC: Fig. 1) of the South Orkney Islands occurred in Early Jurassic times (200-176 Ma) (Trouw et al 1997) and these rocks are overlain by the Early to Middle Jurassic Powell Island conglomerate (Cantrill 2000). In the Patuxent Range of the Pensacola Mountains, East Antarctica (PM: Fig. 1), an early Jurassic phase of deformation pre-dates the intrusion of 183 Ma lamprophyre dykes and

involved an inferred vertical axis rotation of pre-existing, end-Cambrian Ross and PermoTriassic Gondwanian Orogeny structures with the localized development of a spaced foliation and mesoscale folding (Curtis 2002).

South America The Duke of York flysch of the Madre de Dios accretionary complex and the accretionary, Chonos metamorphic complex of Patagonia (C/M: Fig. 1) both show a Late Triassic-Early Jurassic metamorphic event (Thomson & Herve 2002), between 234 Ma and 198 Ma, affecting

EPISODICITY OF TERRANE ACCRETION sediments with detrital age patterns indicating Middle Triassic to Early Jurassic deposition (Herve & Fanning 2001; Herve et al 2003). In the Bolfin Complex of the north Chilean Coastal Cordillera (CC: Fig. 1), amphibolitic gneisses and amphibolites were deformed by ductile transpressional shear zones that were active during or post-emplacement of gneiss igneous protoliths at c. 200-191 Ma, and which are cross-cut by granite emplaced at c. 174 Ma (Scheuber & Gonzalez 1999). Across South America, multiphase reactivation of preexisting basement fault systems, such as the Gastre Fault system (GF: Fig. 1) (Rapela et al 1992), occurred during plate reorganization in Triassic-Jurassic times (Jacques 20030, b).

149

posed deformed arc and subduction complex rocks with parautochthonous continental rocks (Hansen & Dusel-Bacon 1998). Yukon-Tanana terrane deformation of this episode in the Teslin tectonic zone of southern Yukon Territory (TZ: Fig. 1) is Late Triassic to Early Jurassic in age (Stevens & Erdmer 1996). In the McHugh Complex subduction melange of southeast Alaska (MH: Fig. 1), blueschists record a pulse of accretion at c. 195 Ma (Kusky & Bradley 1999).

Eurasia

In the Late Triassic and Early Jurassic the southern Eurasian margin was affected by a major erogenic event, the Indosinian (Golonka 2004). At this time several microplates were Central and North America sutured to the Eurasian margin, closing the In the central Oaxaca terrane of Mexico (OT: Palaeotethys Ocean (Golonka 2004). Late Fig. 1), post-early Permian and pre-Middle Triassic strike-slip and folding affected rocks Jurassic thrusting and folding affected sedi- from the Taimyr Peninsula in northern Russia mentary rocks (Centeno-Garcia & Keppie (TA: Fig. 1) (Inger et al 1999) and extends to 1999) and post-Norian and pre-Aalenian Novaya Zemlya (NO: Fig. 1) and the south (203-175 Ma) thrusting affected Late Triassic Barents Sea (Torsvik & Andersen 2002). Late submarine fan deposits at the contact between Triassic deformation of sediments in the Gobi the Guerrero and Sierra Madre terranes in the Basin of Mongolia (GB: Fig. 1) is constrained Mesa Central of Mexico (MC: Fig. 1) (Silva- by radiometric dating to have occurred between Romo et al 2000). An episode of Late Triassic 209 Ma and 206 Ma (Johnson 2004). Many sedideformation is widespread in western Nevada mentary basins in central Asia record Late (Thomson & Herve 2002). For example, Permo- Triassic deformation (Otto 1997). Ductile shear Triassic marine clastic rocks of the Diablo and zones in Liaoning, North China (LP: Fig. 1) Candelaria formations of western Nevada (WN: show top-to-the-southeast oblique thrusting Fig. 1) were overthrust by ophiolitic melange dated at c. 219 Ma by the Ar/Ar method (Zhang prior to emplacement of rhyolitic and dacitic et al 2002). Chen et al (2003) described seven dykes dated at 222-192 Ma (Thomson et al Mesozoic compressional phases in an overall 1995). Deformation fabrics associated with the tensional, probably subduction-related regime Luning-Fencemaker Thrust Belt of north- from southeast China, the earliest of which is central Nevada (NN: Fig. 1) formed between Late Triassic in age and coincided with the 201 Ma and 193 Ma with deformation probably climax of collision between the North and South continuing to early Toarcian (Wyld et al 2001). China blocks (NS: Fig. 1) (Chang 1996). A synThe Cache Creek terrane of the Canadian kinematic granite pluton, dated at c. 197 Ma by Cordillera (CT: Fig. 1) was imbricated by the U-Pb zircon and Rb-Sr methods, times latecollision with the Quesnel terrane following stage Indosinian deformation on the Songpaneruption of basaltic lavas at c. 210 Ma and prior Ganzi fold belt, Sichuan, west-central China to granodiorite emplacement at c. 173 Ma (SG: Fig. 1) (Roger et al 2004). Deposition in (Harris et al 2003; Lapierre et al 2003). The the flexural foredeep to this fold belt, the Quesnel terrane in southeastern British Longmen Shan foreland basin (LS: Fig. 1), Columbia (QT: Fig. 1) preserves evidence of spanned the time period from 227 Ma to 206 Ma this deformation where c. 215 Ma Late Triassic (Yong et al 2003). High greenschist to amphidiorite is deformed prior to intrusion by unde- bolite grade metamorphism occurred during formed 197-181 Ma diorite and unconformable Late Triassic-Middle Jurassic terrane collision deposition of Early Jurassic sediments (Acton et in the Fujian Province of southeast China (FP: al 2002). In the Yukon-Tanana terrane of east- Fig. 1) (Lu et al 1994); similar-aged collision on Central Alaska (YT: Fig. 1) Late Triassic to the Mianlue suture zone, part of the collision Early Jurassic (212 Ma to 188-185 Ma) north- zone between the North China and South China west-directed, apparently margin- parallel blocks, in central China is associated with major contraction and imbrication resulted in juxta- gold mineralization (NS: Fig. 1) (Vielreicher

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et al 2003). Late Triassic-Early Jurassic deformation is also weakly developed in Peninsular Malaysia, related to collision of an amalgamated Sibumasu/Indochina/South China with North China along the Qinling-Dabei Suture Zone (Meng & Zhang 2000) and to closure of Permo-Triassic rift basins along the Song-Da zone in Laos and Vietnam (SD: Fig. 1) (Metcalfe 2000). The Mazar accretionary prism in the Kunlun Mountains of southwest China (KM: Fig. 1) records Late Triassic-Early Jurassic docking between the Gondwanian Karakoram-Qiangtang block and the Eurasian Kunlun block immediately prior to or during emplacement of 215-190 Ma arc plutons (Xiao et al 2002). In the Karakorum, Carnian-Norian carbonate sediments folded by Eo-Cimmerian deformation of Late Triassic age are overlain unconformably by Pliensbachian-Toarcian redbed deposits (Gaetani 1997). Zircon fission track cooling ages from the northern edge of the Tibetan Plateau indicate uplift related to collision of the Qiantang block and Kunlun blocks (KM: Fig. 1) in the period 220-200 Ma, and this is reflected by a corresponding period of non-deposition or erosion in the nearby Tarim Basin (Jolivet et al 2001). Between the Caspian Sea and the Tien Shan in Central Asia, accretion of central Iran to the southern margin of Eurasia resulted in an intense, Late Triassic episode of deformation with development of a widespread unconformity (Thomas et al 1999). The eastern part of the Donbas fold belt of the Ukraine (DF: Fig. 1) records intense Late Triassic-aged (Cimmerian) deformation with widespread unconformity development between Triassic and Jurassic sediments (Stovba & Stephenson 1999). Blueschists in metabasite rocks along the Izmir-Ankara suture in northern Turkey (IA: Fig. 1) record a late Triassic pulse of Palaeo-Tethys subduction from 215 Ma to 205 Ma during the Cimmeride Orogeny (Okay et al 2002). Cimmeride deformation in Turkey was complete by Sinemurian times (c. 197 Ma) (Okay et al 2002).

New Zealand Bradshaw et al (1981) described a Late Triassic deformation episode, the Rangitata I orogeny, which marked the accretion of the older Rakaia subterrane of the Torlesse terrane to a trenchslope basin suite represented by the Caples terrane. White mica samples from deep levels of the Otago and Marlborough schists of South Island, New Zealand (NZ: Fig. 1), yielded Ar/Ar method ages that give a peak age for

Otago Schist metamorphism of 180-170 Ma (Little et al. 1999). This deformation is interpreted to have resulted from collision between the Torlesse terrane and Caples terrane (Mortimer 1993) and corresponds in time to the Rangitata I orogeny (Bradshaw et al 1981). Rb-Sr method isochron ages of Brook Street, Dun Mountain-Maitai and Murihiku terrane metasediments of South Island show cooling age patterns that indicate low-grade metamorphism in latest Triassic and earliest Jurassic times (Adams et al 2002). This is possibly coincident with conglomerate deposition from 200-180 Ma that recorded docking between the Median Tectonic Zone and Brook Street terranes (Tulloch et al 1999). Torlesse terrane metasediments also show Late Triassic to earliest Jurassic ages (225-200 Ma) of burial metamorphism for the Rakaia subterrane, with postmetamorphic uplift and cooling continuing through to Middle Jurassic times (c. 170 Ma) (Adams & Graham 1996).

Other Late Triassic-Early Jurassic events Magmatism Eruption of continental flood magmatism provinces (areas of lava in excess of 100 000 km2) and the formation of large igneous provinces (LIP) is characteristic of the Late Triassic-Early Cretaceous period. Three continental LIPs, two basaltic and one rhyolitic, were erupted in the period c. 205-160 Ma (Marzoli et al 1999; Riley & Knight 2001; Riley et al 2001) with conspicuous peaks of magmatism at 200 Ma and 183 Ma. The Central Atlantic Magmatic Province (CAMP: Fig. 1) is one of the largest continental igneous provinces erupted during the Phanerozoic (Tanner etal 2004). The province erupted largely from c. 205 Ma to 199 Ma (Marzoli et al 1999; Nomade et al 2002) with early olivine dolerites at c. 209 Ma in the Nova Scotia region (Pe-Piper & Reynolds 2000). The Gondwana Large Igneous Province (CLIP: Fig. 1) (Riley et al 2001; Storey & Kyle 1997) includes southern African Karoo basalts, Antarctic Ferrar gabbro and Kirkpatrick basalts, and the silicic magmatic province of South America and the Antarctic Peninsula, which constitutes one of the largest silicic provinces on Earth (Pankhurst et al 2000). Basaltic eruption appears to have been concentrated in a brief period at around 183-182 Ma (Riley & Knight 2001; Riley et al 2004), with rhyolitic magmatism in three pulses over a more extended period between 188 Ma and 153 Ma

EPISODICITY OF TERRANE ACCRETION (Pankhurst et al. 2000). Extensive bimodal basalt-rhyolite magmatism in the period 186-170 Ma, hosting economically important precious and base-metal mineral deposits, is also recorded from British Columbia (MacDonald et al 1996; Evenchick & McNicoll 2002). Absence of an oceanic record for this period makes identification of submarine flood magmatic provinces difficult; however, at least three potential oceanic plateau candidates exist: (1) the extensive Wrangellia terrane of northwest North America may represent an accreted basaltic oceanic plateau that erupted at c. 227 Ma (Kerr et al. 2000); (2) the Cache Creek terrane of the Canadian Cordillera may represent an accreted oceanic plateau that erupted at c. 210 Ma (CTP: Fig. 1) (Lapierre et al 2003); and (3) an allochthonous tectonic unit of mafic volcanic rocks, the Nilufer Unit of northwest Turkey, may represent a Palaeo-Tethyan, Triassic oceanic plateau accreted to the Laurasian margin during the Late Triassic (NUP: Fig. 1) (Okay et al. 2002). In addition to extrusive magmatism, major plutonic complexes of Late Triassic to Middle Jurassic age are common on the Pacific and Tethyan margins of Pangaea. For example, in the Antarctic Peninsula, plutonism was active in the Late Triassic-earliest Jurassic (220-200 Ma) (Wever et al. 1995; Millar et al. 2002) and in the Middle Jurassic (170-160 Ma) (Pankhurst et al. 2000). Within-plate granites of this age are also recorded from the Ellsworth-Whitmore Mountains elsewhere in West Antarctica (Storey et al. 1988). Plutonism in southeastern Peru was active in the intervals 225-190 Ma and 180-170 Ma (Clark et al. 1990) and in Ecuador from 227 Ma to 200 Ma (Noble et al 1997). Igneous rocks of Late Triassic to Early Jurassic age (228-175 Ma) are widespread in the western Cordillera of North America and have potassic or shoshonitic compositions unlike igneous rocks before or after (Mortimer 1986,1987). For example, in the east-central Sierra Nevada of California an intense pulse of K-rich plutonism has been identified at 180-165 Ma (Coleman et al 2003) and in northwest Nevada high-K calcalkaline plutons intruded between c. 196 Ma and c. 190 Ma (Quinn et al 1997). In Sichuan Province, west-central China, plutonism was active from c. 197 Ma to 153 Ma, in northwestern China plutonism at c. 198 Ma is associated with gold mineralization (Qi et al. 2004) and, in southeast China, Jurassic plutons intruded in the interval 175-160 Ma and cover an area of 900 X 400 km (Li et al. 2003). Late TriassicEarly Jurassic plutonism in the range

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211-183 Ma is also recorded from the Korean Peninsula and Japan (Cluzel et al 1991; Arakawa et al 2000; Kamei et al 2000). Late Triassic magmatism (c. 220 Ma) is also evident in New Guinea (Crowhurst et al. 2004). In North America, peaks of kimberlite magmatism occur at c. 196 Ma and 180-176 Ma (Heaman & Kjarsgaard 2000), and at c. 172 Ma (Kopylova et al. 1998); the Yakutian kimberlite field of Siberia shows peaks in emplacement from 220 Ma to 240 Ma and from 170 Ma to 150 Ma (Heaman et al. 2003); in Africa, a pulse of kimberlite emplacement is recognized from 190 Ma to 150 Ma (Heaman et al. 2003); and, in Australia in the Northern Territories, kimberlites are recorded at c. 179 Ma (Belousova et al. 2001) and, in South Eastern Australia, at c. 170 Ma (with lamproite magmatism at c. 187 Ma) (Foden et al. 2002).

Geomagnetic field Geomagnetic field changes around the TriassicJurassic boundary are small and periods of extended polarity short, compared with what is recorded for the mid-Cretaceous period (e.g. Gallet & Hulot 1997). The Early Jurassic is relatively poorly known (Yang et al 1996); however, a several million-year period of anomalously low reversal rate has been observed for the Triassic-Jurassic boundary (Johnson et al 1995; Gradstein et al 2004). Although possibly not significant for changes at the core-mantle boundary, this is roughly an order of magnitude longer than polarity intervals during the preceding period of high reversal rate (mean c. 500 ka) in the Late Triassic (Kent & Olsen 1999).

Atmosphere and oceans Short-lived periods of extreme oceanic anoxia are recorded for the Late Triassic-Early Jurassic interval (Hesselbo et al 2000) with anomalous activity extending up to the early Middle Jurassic (Hesselbo et al 2003). For example, the global Early Toarcian (Early Jurassic) oceanic anoxic event (OAE) occurred at c. 183 Ma (Hesselbo et al. 2000) and carbon isotope excursions occurred at the TriassicJurassic boundary (Hesselbo et al 2002). Other OAEs also occurred at the end of the Norian (late Triassic; c. 204 Ma) (Sephton et al 2002) and in the Pliensbachian (Early Jurassic; 189-183 Ma) (Borrego et al 1996). CO2 levels, and probably atmospheric temperatures, were elevated substantially across the TriassicJurassic boundary (Beerling & Berner 2002),

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across the Pliensbachian-Toarcian boundary (Hesselbo et al 2000) and in the early Middle Jurassic (174-170 Ma) (Hesselbo et al 2003). Examination of carbon isotopes in wood from this time suggest huge influxes of isotopically light carbon into the upper oceans, biosphere and atmosphere, for which rapid release of methane hydrate from marine continental margin sediments has been implicated (Hesselbo et al. 2000; 2002; 2003). Indicating the scale of this event, a major enrichment is seen in 613C of hydrocarbons in the Late Triassic to Early Jurassic, one of only three such shifts seen in the Late Proterozoic and Phanerozoic (Andrusevich et al. 1998). Episodes of anoxia and accumulation of carbon-rich sediments resulted in formation of abundant oil source rocks, including Late Triassic-Middle Jurassic coal in Siberia (Il'ina & Shurygin 2000; Kirichkova & Travina 2000; Kirichkova & Kulikova 2002; Kirichkova et al. 2003). The early Toarcian was a period of sea-level rise and marine transgression (Schouten et al 2000) and the Triassic-Jurassic boundary has also recently been recognized as a time of sea-level rise (Hesselbo et al 2004). Biotic changes Two major mass extinction events occurred in the Late Triassic-Early Jurassic period. The largest, one of the 'Big Five' Phanerozoic mass extinctions (Newell 1967; Jablonski & Chaloner 1994; Hallam & Wignall 1999), occurred at the Triassic-Jurassic boundary (c. 200 Ma) (Tanner et al 2004) and the second across the Pliensbachian-Toarcian boundary (c. 183 Ma) (Aberhan & Fursich 2000), possibly in an extended period of extinctions lasting approximately seven million years (Little & Benton 1995). Eruption of large igneous provinces and extinction events have been linked causally but the relationships are not straightforward (Wignall 2001). Oceanic anoxia and extinction at the Pliensbachian-Toarcian boundary has been linked with eruption of the main phase of the Gondwana Large Igneous Province (Palfy & Smith 2000; Riley & Knight 2001; Riley et al. 2004) and the end-Triassic extinction event has been linked with the eruption of the Central Atlantic Magmatic Province (Wignall 2001; Guex etal 2004). Other extinctions are recognized in this period, including one in Late Norian times affecting Monotid bivalves and most ammonoids (c. 204 Ma) (Sephton et al. 2002) and two affecting ammonites in the Early Toarcian (c. 180 Ma) (Cecca & Macchioni 2004).

Distribution of mid-Cretaceous deformation around the palaeo-Pacific Ocean Antarctic Peninsula Mid-Cretaceous deformation affects rocks forming all three tectonostratigraphic terranes identified by Vaughan & Storey (2000) (AP: Fig. 2). In the Western Domain terrane (Vaughan & Storey 2000), the Fossil Bluff Group (Butterworth et al 1988; Crame & Hewlett 1988; Macdonald & Butterworth 1990) is interpreted to be an Early Jurassic to late Early Cretaceous forearc basin sequence (Doubleday & Storey 1998; Nichols & Cantrill 2002) overlying accretionary complex rocks on Alexander Island. Deposition of sediments in this basin terminated in the late Albian (c. 100 Ma) with shoaling and uplift (Doubleday et al 1993; Storey et al 1996; Nichols & Cantrill 2002) during dextral transpression (Doubleday & Storey 1998). In the Central Domain terrane (Vaughan & Storey 2000) in northwest Palmer Land, a major flower structure, forming a sinistral transpressional transfer zone, deforms Cretaceous gabbro, Late Triassic granodiorite and marble tectonic breccia of unknown age (Vaughan et al. 1999). Biotite cooling ages in deformed gabbro and mafic dykes from the flower structure date sinistral shear and thrusting there at c. 110 Ma (Vaughan et al. 1999). Mid-Cretaceous deformation is developed most strongly in the Eastern Domain terrane (Vaughan & Storey 2000). The eastern Palmer Land Shear Zone (Vaughan & Storey 2000) is a major ductile shear zone, up to 15 km wide, formed by dextral transpression and eastdirected orthogonal compression, with activity dated at c. 103 Ma (Vaughan et al 2002a). This biotite cooling age, dated by the Ar/Ar method, was obtained from ductilely thrusted midJurassic gabbro. This confirmed earlier work from eastern Palmer Land (Meneilly 1983; 1988) that suggested that east-directed overthrusts cut granodiorite dated at c. 113 Ma. In the Eastern Domain further south, in Ellsworth Land, Middle Jurassic to Early Cretaceous (Laudon & Fanning 2003) marine sediments of the Latady Formation (Laudon et al 1983) are thrust and folded openly (Vaughan et al 20026). The peak of deformation coincided with emplacement of a granitoid dyke dated by the U-Pb SHRIMP method at c. 107 Ma (Vaughan et al 20026). Sedimentary sequences of the Eastern Domain also show a paucity of sediments between late Berriasian and Barremian times (Hathway 2000), prior to deposition of an

EPISODICITY OF TERRANE ACCRETION

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Fig. 2. Distribution of areas preserving mid-Cretaceous deformation and magmatism on 100 Ma reconstruction of the Pacific basin. Areas of large igneous provinces or inferred oceanic plateaux outlined with dashed line; kimberlite magmatism marked with K; plutonism marked with a star. AF, Atacama Fault; AK, Anadyr-Koryak region; AP, Antarctic Peninsula; BC, British Columbia; BUB, Bunbury basalts; CA, Carpentaria Basin; CB, Coast Belt; CD, Cordillera Darwin; CJ, central Japan; DI, Duke Island; El, Elephant Island; FP, Fujian Province; GT, Guerran terrane; GU, Guatemala; HI, Hispaniola; HIP, Hikurangi plateau; HO, Honduras; JS, Java-Kalimantan-Sulawesi; LAP, Lesser Antilles plateau; LR, Lanterman Range; MAP, Manihiki plateau; MB, Marie Byrd Land; MD, Mirano diastrophic phase; ME, Mochica Event; NI, North Island, New Zealand; NJ, northern Japan; NV, northern Venezuela; OJP, Ontong-Java plateau; PB, Papuan Basin; PR, Peninsular Ranges; QS, Qilian Shan; SA, Sikhote-Alin Mountains; SB, Surat basin; SG, South Georgia; SH, Shikoku; SN, Sierra Nevada; SI, South Island, New Zealand, TS, Truong-Son belt; WC, western Columbia; WI, Whitsunday Islands; WP, western Philippines; WVP, Whitsunday Volcanic Province and associated magmatism.

Aptian to Eocene megasequence. This hiatus probably resulted from changes in deposition and erosion caused by mid-Cretaceous deformation. The continuation into the northern Antarctic Peninsula of the terranes identified by Vaughan & Storey (2000) in Palmer Land and Ellsworth Land is unclear and evidence for mid-Cretaceous deformation is more sparse; however, this may be a function of more recent and more intense mapping activity in the

southern Antarctic Peninsula. In northern Graham Land, west-directed, ocean-vergent folding and thrusting deformed mid-Cretaceous basaltic lavas of the Antarctic Peninsula Volcanic Group, dated at c. Ill Ma, and these structures were cut by granodiorite plutons dated at 96 Ma (Birkenmajer 1994). In northeast Graham Land, on the margin of the Larsen Basin, localized folds and faults formed during a mid-Cretaceous episode of sinistral

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transpression (Del Valle & Miller 2001). East of northern Graham Land, on James Ross Island, coarse conglomerate units were deposited and rafts of extrabasinal Jurassic strata incorporated in mostly mudstone Aptian to Albian back-arc sediments of the Larsen Basin (Ineson 1985, 1989; Riding & Crame 2002). Although this constitutes less direct evidence for deformation, these higher-energy events suggest that the arc was uplifted and rejuvenated in the midCretaceous. At the northern tip of the Antarctic Peninsula, along the South America-Antarctica plate boundary in the vicinity of Elephant Island, Grunow et al (1992) (El: Fig. 2) interpreted that mid-Cretaceous transpression had overprinted earlier evidence for subduction activity and Trouw et al. (2000) placed this deformation at 110-90 Ma. East of the Antarctic Peninsula, in the Weddell Sea, Geosat radar altimetry data reveal major changes in spreading during the Cretaceous normal polarity interval (Livermore & Woollett 1993). Other areas of Antarctica In this volume, Siddoway et al. (2005) described new dyke data from Marie Byrd Land that indicate an episode of dextral shear pre-dated formation of extensional gneiss domes in the period 105-94 Ma and prior to the onset of New Zealand-Antarctica rifting. Kinematic data from further east along the margin (Vaughan & Storey 2000) and interpretation of seafloor data (Sutherland & Hollis 2001) support midCretaceous dextral shear along this segment of the Gondwana margin. In East Antarctica, adjacent to the Rennick graben in the Lanterman Range of Victoria Land, Permo-Triassic Beacon Group sediments and Jurassic volcanic rocks are thrust and folded tightly (Roland & Tessensohn 1987), during an episode of sinistral transpression of possible mid-Cretaceous age (Tessensohn 1994) (LR: Fig. 2), although later authors have argued for Cenozoic-aged deformation (Rossetti et al. 2003). South America Dalziel (1986) noted the synchroneity of the first phase of mid-Cretaceous deformation (Mirano diastrophic phase, 112-105 Ma; MD: Fig. 2) along the western margin of South America from the Magellan Basin to the West Peruvian Trough, post-dating Aptian-Albian sediments and pre-dating intrusions and sediments formed from 105-80 Ma. Evidence of hiatuses in sedimentation is also widespread in western South America in mid-Aptian and

Aptian-Albian times (Macellari 1988). An unconformity related to one of these hiatuses cuts felsic tuff, dated at c. 104 Ma, in the Andes of central Chile (Charrier et al. 1996). Thrusting deformed Aptian-Albian sedimentary and ophiolitic rocks of the mid-Jurassic to Lower Cretaceous Rocas Verdes basin in South Georgia (SG: Fig. 2) (Macdonald et al. 1987) and in southern Patagonia (Dalziel 1986) and is linked to closure of a back-arc basin with associated ophiolite obduction (Dalziel 1981). In Cordillera Darwin (Kohn et al 1993) (CD: Fig. 2) amphibolite facies metamorphism of Palaeozoic basement rocks is associated with thrustrelated deformation of upper Aptian-Albian sedimentary rocks (Dott et al. 1977) with shortening of over 430 km from 110 Ma to 90 Ma (Halpern & Rex 1972) in the mid-Cretaceous (Kraemer 2003). Cunningham (1995) pointed out that this orogeny exhumes some of the highest metamorphic-grade rocks in the Andes south of Peru and demonstrated a transpressional component to deformation with strain partitioning. In northern Chile, the Atacama Fault system (AF: Fig. 2) exhibits transpressional movement and uplift which deformed plutonic and volcanic rocks older than c. Ill Ma (Thiele & Pincheira 1987) and arc activity ceased in the midCretaceous in that area (Brown et al. 1993). The fault system was inverted in post-Early Cretaceous times by contractional deformation that also resulted in widespread thin-skin thrust deformation of sediments (Grocott & Taylor 2002). Late Albian compression folded Lower Cretaceous submarine volcanic sequences in western Peru (Mochica deformation event of Megard 1984) (ME: Fig. 2). Late Albian uplift and erosion of mid-Albian sandstones is evident in western Columbia (Barrio & Coffield 1992) (WC: Fig. 2), and late Early Cretaceous metamorphism and deformation of Upper Jurassic to Lower Cretaceous arc rocks is widespread from Columbia to Mexico (Tardy et al. 1994; Dickinson & Lawton 2001). White mica and amphibole Ar/Ar method ages from high pressure metamorphic rocks in the Coastal Fringe/Margarita belt of northern Venezuela (part of the Caribbean-South American plate boundary, NV: Fig. 2) suggest that transpressional deformation occurred in mid-Cretaceous times at c. 96 Ma (Smith et al. 1999). Central America and Caribbean In Hispaniola, Puerto Rico and the Virgin Islands (HI: Fig. 2), Lower Cretaceous islandarc rocks, believed to have formed part of a

EPISODICITY OF TERRANE ACCRETION southwest-facing arc in the Early Cretaceous (Pindell & Barrett 1990), are truncated by a major erosional unconformity overlain by shallow-water, Aptian-Albian limestone (Lebron & Perfit 1993). Draper et al (1996) linked this to thrust emplacement of the Hispaniola peridotite belt in Aptian-Albian times. In Cuba, the Antilles arc accreted basaltic rocks of ocean plateau affinity in the period 112-91 Ma (Kerr et al 1999). Phengitic micas from the southern belt of the Motagua fault zone of Guatemala (GU: Fig. 2) yield Ar/Ar method deformation ages in the range 125113 Ma, recording mid-Cretaceous collision of the Chortis Block with western Mexico (Harlow et al 2004). Early to mid-Cretaceous uplift and erosion affects a Barremian to mid-Albian limestone sequence in Honduras (Weiland et al 1992) (HO: Fig. 2). Dickinson & Lawton (2001) described the accretion of an arc terrane to the Mexican margin in late Early Cretaceous time. In Mexico, greenschist facies metamorphism, linked to collision of the Guerrero terrane (Tardy et al 1994) (GT: Fig. 2) affects plutons dated to be between 112 Ma and 101 Ma (Stein et al 1994).

North America In southern California, ocean-vergent, southwest-directed overthrusting has deformed 118 Ma, magmatic-arc granodiorite of the western Peninsular Ranges batholith (PR: Fig. 2) with the formation of mylonite zones dated at 115 Ma (Thomson & Girty 1994). Mylonite and granodiorite are truncated by 105 Ma tonalite (Thomson & Girty 1994) and deformation was probably active in the period 115-108 Ma (Johnson etal 1999). West-directed (i.e. oceanward) compression of the western Peninsular Ranges batholith in southern and Baja California also pre-dated onset of emplacement of the eastern Peninsular Ranges batholith at 99 Ma (Kimbrough et al 2001). A peak of Franciscan melange blueschist metamorphism appears to have occurred at this time (Ernst 1993). In the Sierra Nevada of California (SN: Fig. 2) a major episode of thrusting is recognized deforming Jurassic and Cretaceous plutonic rocks between c. 110 Ma and c. 92 Ma (Mahan et al 2003). For example, contractional deformation on the Bench Canyon shear zone shows multiple phases from c. 101 Ma, with a peak at c. 95 Ma (McNulty 1995), and synkinematic sheared orthogneiss from the Tiefort Mountains is dated at c. 105 Ma (Schermer et al 2001). A major thrust system of mid-Cretaceous age is developed along much of the Coast Belt

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of northwestern North America (CB: Fig. 2) (Umhoefer & Miller 1996). Regional midCretaceous deformation is recognized in British Columbia (MacDonald et al 1996) (BC: Fig. 2), with, for example, ocean-vergent thrusting (Journeay & Friedman 1993) of upper Valanginian to middle Albian volcanic rocks. Thrusting was active from at least mid-Albian times to the emplacement of syn-kinematic plutons dated at c. 97 Ma and c. 96 Ma and deformed an accretionary complex that formed prior to 92 Ma (Chardon 2003). In Cretaceous sedimentary sequences of northwestern Canada, major, regionally extensive unconformities exist in mid-Aptian, and between Albian and Cenomanian, strata (Dixon 1993, 1997). On Duke Island, in southeast Alaska, a major and widespread episode of thrusting deformed ultramafic and arc rocks intruded from 111 Ma to 108 Ma (Butler et al 2001) (DI: Fig. 2). In central and northern Alaska, ductile, north-directed overthrusting was active with blueschist and amphibolite-facies metamorphism dated at c. 108 Ma. This formed during an episode of overall extension and metamorphic core-complex formation in the southern Brooks Range (Till et al 1993; Till & Snee 1995) and major thrusting and tectonic transportation of ophiolite terranes was coeval in the northern Brooks Range (Cole et al 1997).

East and Southeast Asia In the Anadyr-Koryak region of northeast Russia (AK: Fig. 2), Albian to Cenomanian volcanic rocks, and Barremian-Aptian and Albian clastic sequences, overlie unconformably folded Callovian to Hauterivian island-arc rocks and Palaeozoic to Hauterivian arc marginal volcano-sedimentary sequences (Stavsky et al 1990; Filatova & Vishnevskaya 1997). Deformation occurred in Barremian-Aptian (c. 120 Ma) and Albian (c. 105 Ma) times (Stavsky et al 1990; Filatova & Vishnevskaya 1997). In the Chukotka Peninsula of the Anadyr-Koryak region, compressional deformation from 124 Ma to 117 Ma, based on Ar/Ar method dating of polydeformed greenschistgrade phyllites and marbles, is followed by a period of metamorphic core-complex formation from 109 Ma to 104 Ma (Toro et al 2003). Ophiolitic thrusting in the Sikhote-Alin mountain range, eastern Siberia (SA: Fig. 2), is dated at c. 110 Ma by the Ar/Ar method (Faure et al 1995). Ophiolitic rocks in northern Japan (NJ: Fig. 2), containing a conformable sequence of Tithonian pelagic sedimentary rocks and Hauterivian to post-Aptian terrigenous forearc

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sedimentary rocks, were underthrust and accreted with uplift and metamorphism of arc rocks from Barremian times (130-125 Ma) to at latest the time of blueschist metamorphism dated at 110 Ma (Kimura et al 1994; Kimura 1997). Some evidence for a continuation of more rapid accretion is suggested by more intense metamorphism of ophiolitic units accreted between Aptian and Cenomanian times (Kiyokawa 1992; Ueda etal 2000). Sanbagawa metamorphism is estimated to have occurred between c. 120 Ma and c. 110 Ma, with a peak at c. 116 Ma, in central Shikoku, Japan (Isozaki & Itaya 1990; Okamoto et al 2004) (SH: Fig. 2). Associated metamorphism at c. 116 Ma is also recorded for the Japanese Median Tectonic Line in the South Kitakami and palaeo-Ryoke belts in central Japan (Sakashima et al 2003) (CJ: Fig. 2). In the Fujian Province of southern China (FP, Fig. 2), gabbroic and ultramafic rocks of a dismembered ophiolite suite (Wang & Lu 1998) were emplaced on Palaeozoic basement rocks by ocean-vergent thrusting coeval with c. 115 Ma gabbro and granitoid plutons (Wang 2002), and compressional shear joints dated at 107 Ma (Lu et al 1994). Fission track dating also suggests that uplift had ceased by c. 110 Ma (Chen et al 2002). Ar/Ar method cooling ages from the Changle-Nanao ductile shear zone in the same region suggest rapid uplift during the period 118-107 Ma (Wang & Lu 1997). In the Qilian Shan of western China (QS: Fig. 2), apatite and zircon fission track ages indicate a multiplephase history of exhumation in the period 115-90 Ma (George etal 2001) and the Bainang terrane in the Yarlung-Tsangpo suture of southern Tibet records post-late Aptian accretion of Tethyan oceanic rocks (Ziabrev et al 2004). Further west in the Karakoram, Tethyan margin, shallow-marine sediments of Aalenian to Lower Cretaceous age and crystalline basement rocks are incorporated in huge thrust sheets that are capped by thick deposits of midCretaceous conglomerate (Gaetani 1997). The Truong-Son belt, of north-central Vietnam (TS: Fig. 2), shows metamorphism and mylonite formation in the period 120-90 Ma with notable peaks at c. 115 Ma and between 106 Ma and 103 Ma (Lepvrier et al 1997). In the western Philippines (WP: Fig. 2), east-directed thrusting deforms Lower Cretaceous ophiolitic rocks overlain unconformably by a middle Eocene hemipelagic sedimentary sequence (Faure et al 1989), which is possibly related to the collision of the west Philippine block with the ChinaIndochina margin in latest Early Cretaceous times (Lapierre et al 1997). In Java, Sulawesi

and Kalimantan (JS: Fig. 2), an extensive high pressure-ultra-high pressure metamorphic basement terrane was recrystallized and uplifted during collision of a Gondwanan continental fragment with the Sundaland margin between 120 Ma and 115 Ma (Parkinson et al 1998).

Australia In the Surat basin of eastern Australia (SB: Fig. 2), uplifted and eroded upper Aptian marine sedimentary rocks are overlain unconformably by early Albian, paralic sedimentary rocks (Harrington & Korsch 1985; O'Sullivan et al 2000). This event also affected basins in northeastern Australia (Laura and Carpentaria) (CA: Fig. 2) and Papua New Guinea (Papuan) (PB: Fig. 1) (Haig & Lynch 1993). It is associated with increased Aptian-Albian tectonism in the eastern half of Australia, supported by fission track evidence for increased exhumation rates (Marshallsea et al 2000) and palaeobathymetric evidence for non-eustatic changes in base level (Henderson 1998; Campbell & Haig 1999). In the New England Orogeny of eastern Australia, mid-Cretaceous deformation (WI: Fig. 2) affected voluminous (1.4 X 106 km3) volcanic and plutonic rocks of the Whitsunday Volcanic Province and others (WVP: Fig. 2) dated at between 132 Ma and 95 Ma (Harrington & Korsch 1985; Bryan et al 2000), although the age of deformation may be constrained more tightly because magmatism peaked between 120 Ma and 95 Ma (Bryan et al 1997). Deformation is expressed largely as open folds associated with major faults (O'Sullivan et al 2000) and may be associated with a major phase of extension in eastern and southeastern Australia in the middle Cretaceous (O'Sullivan et al 1996; 2000).

New Zealand The tectonic situation of New Zealand is complex, with a terrane interpretation that has evolved rapidly (e.g. Leverenz & Ballance 2001; Kear & Mortimer 2003; Mortimer 2004) since early suggestions for South Island (Coombs et al 1976) and North Island (Sporli 1978) and comprehensive definition by Bradshaw (1989). Deformation of Early Cretaceous age was first recognized in New Zealand (the Rangitata II phase of Bradshaw et al 1981), where it was interpreted initially as a phase of deformation affecting convergent margin terranes. The age of deformation is based on widespread Aptian, Albian and intra-Albian angular unconformities

EPISODICITY OF TERRANE ACCRETION in sedimentary rocks of eastern New Zealand (Bradshaw 1989), with a mid- to late Albian age of last deformation (c. 105 Ma, Bradshaw 1989; c. 100 Ma, Laird & Bradshaw 2004). Since the key papers by Bradshaw et al. (1981) and Bradshaw (1989), much evidence has come to light for different-aged metamorphic events in different parts of New Zealand, for example, at c. 116-105 Ma in Fiordland (Klepeis et al 2004), 100-85 Ma in the Wellington region (Adams & Graham 1996; Kamp 2000), 86 Ma in the Southern Alps (Vry et al 2004) and 71 Ma in Otago (Mortimer & Cooper 2004). In addition, an episode of rifting has been identified associated with local high temperature extension and magmatism, which formed the Paparoa metamorphic core complex between 110 Ma and 90 Ma (Muir et al 1997; Spell et al 2000). Metamorphic events in the Southern Alps and Otago at c. 86 Ma and c. 71 Ma are probably related to final collision of the Hikurangi Plateau and the opening of the Tasman Sea (Mortimer & Cooper 2004; Vry et al 2004). These post-date contractional deformation being considered here. The older events, from Fiordland and the Wellington region, fall in the range of midCretaceous deformation from other parts of the Pacific basin margin. In the Fiordland region of southern South Island, contractional deformation was coeval with magma emplacement of the Median Batholith from 116 Ma to 105 Ma (Klepeis etal. 2004) (SI, Fig. 2). Mid-Cretaceous compression overlaps in age with formation of the hyper-extensional Paparoa metamorphic core complex which was magmatically active from 110 Ma to 90 Ma (Muir et al 1997); however, the earliest depositional ages from intercalated tuffs suggest that the switch from compression to extension was at c. 101-102 Ma (Bradshaw et al 1996; Muir et al 1997) and oroclinal bending associated with midCretaceous compression had probably ended by 110 Ma (Bradshaw et al 1996). In southern North Island (NI, Fig. 2), low temperature metamorphism of sandstones of the accretionary Pahau Subdivision of the Torlesse terrane (Adams & Graham 1996), and deformation, is dated by zircon fission track data to have occurred in the period 100-85 Ma (Kamp 2000). Deformation is represented, for example, by folded, Upper Jurassic to Aptian-Albian argillic and metabasic rocks in the Wellington region, interpreted to be part of a dismembered seamount (George 1993), which are cut by c. 99 Ma lamprophyre dykes.

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New Zealand and expression of plate boundary forces New Zealand is one of the best-studied areas, which has revealed that mid-Cretaceous deformation there has a complex expression. An examination of Cenozoic to Recent tectonic activity in New Zealand illustrates the kind of variation that can be seen along quite short stretches of an active margin. In particular, superimposed on the pre-rift (Laird & Bradshaw 2004) terrane complexity described by Bradshaw (1989) and Mortimer (2004) is the large-scale, c. 460 km (Sutherland et al 2000), Cenozoic movement of the Alpine Fault (Norris et al 1990). The fault is a distributed strike-slip transfer zone (Hall et al 2004) between eastdirected subduction to the south (Norris & Cooper 2003) and west-directed subduction to the north (Delteil et al 2003) and illustrates that the expression of plate boundary forces is not necessarily straightforward. In this context, the plate boundary forces, changes in subduction rate, potential slab capture and collision of the Hikurangi Plateau, all of which combined to drive mid-Cretaceous deformation in New Zealand, may have been expressed in different but co-existing ways. In South Island, one probably sees the most straightforward expression, where accelerated plate motion and heat flow are likely to have driven magmatism and contractional deformation in the period 116-105 Ma. In North Island, the situation was complicated by the approach of the Hikurangi Plateau (Vry et al 2004) and a possible ridge-trench collision (Luyendyk 1995). Contractional deformation appears to have stopped at about 100 Ma, followed by a period of slab rollback and extension prior to final collision of the Hikurangi Plateau at c. 86 Ma. This was followed by opening of the Tasman Sea (Vry et al 2004) and associated basins (Carter 1988) and onset of New Zealand-Antarctica rifting (Laird & Bradshaw 2004).

Other mid-Cretaceous events Magmatism Eruption of submarine basalt plateaux (areas of oceanic lithosphere with thickness >10 km) and the formation of basaltic large igneous provinces (LIP) is characteristic of the midCretaceous period (Kerr et al 2000). Six oceanic LIPs were erupted in the period 122-88 Ma (Coffin & Eldholm 1994; Mortimer & Parkinson 1996) with conspicuous peaks of magmatism at 122 Ma and 90-88 Ma (Kerr et al 2000). The

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Ontong-Java plateau (OJP: Fig. 2) is the largest known in the geological record (Coffin & Eldholm 1994) and erupted in close proximity to the Manihiki (MAP: Fig. 2) and Hikurangi (HIP: Fig. 2) plateaux. Fragments of an Aptian-Albian oceanic plateau that accreted to the Antilles arc (LAP: Fig. 2) in the period 112-91 Ma are preserved in Cuba (Kerr et al. 1999). Not all plume impact sites were under oceanic lithosphere. The Rajmahal Traps of eastern India were erupted at c. 118 Ma and were possibly related to emplacement of the chemically similar Kerguelen plateau at the same time (Kumar et al 2003) and the second phase of the Bunbury basalts of western Australia (BUB: Fig. 2), which erupted at c. 123 Ma (Courtillot & Renne 2003). All three are closely associated spatially on reconstructions of Gondwana and pre-date rifting-off of India (Kumar et al. 2003). New evidence suggests that the Bermuda plume was responsible for 2-3 km of mid-Cretaceous uplift with erosion and magmatism in the Mississippi Embayment (Cox & Van Arsdale 2002), and Maher (2001) has identified a largely terrestrial, mid-Cretaceous high-Arctic large igneous province, with magmatic ages currently poorly constrained in the range 135-90 Ma. As in the case of the Bermuda plume, the high-Arctic province is associated with uplift and denudation. For example, mid-Cretaceous apatite fission track cooling ages in Carboniferous sediments, probably associated with exhumation, have been recorded in East Greenland at this time (Johnson & Gallagher 2000); a widespread mid-Cretaceous unconformity is developed in the northern part of the mid-Norwegian margin (Lundin & Dore 1997); and an associated unconformity is developed as far south as the Atlantic Western Approaches Trough (Ruffell 1995). Granitic magmatism also shows a widespread mid-Cretaceous peak. The Lassiter Coast intrusive suite of the southern Antarctic Peninsula, which covers an area of c. 500 X 100 km, shows a major peak in magmatism between about 115 Ma and 95 Ma (Pankhurst & Rowley 1991). A major peak of I-type magmatism is evident in Thurston Island, with emplacement from 125 Ma to 110 Ma (Pankhurst et al. 1993), and a pulse of A-type granitoid magmatism is recorded in western Marie Byrd Land between 115 Ma and 95 Ma (Siddoway et al. 2005). Bruce et al (1991) identified a mid-Cretaceous peak of Andean magmatism occurring between 120 Ma and 70 Ma and the North Patagonian batholith shows a peak of emplacement between 120 Ma and 90 Ma (Pankhurst et al 1999). This is

supported by palaeomagnetic data that suggest that the major part of the North Patagonian batholith was intruded during the Cretaceous long normal interval (Beck et al. 2000). A pulse of plutonism is seen in central Mexico from 110 Ma to 100 Ma (Stein et al. 1994). The 800 km long eastern Peninsular Ranges batholith of southern and Baja California was intruded between 99 Ma and 92 Ma, with interpreted magma flow rates comparable to those estimated for flood basalt eruptions (Kimbrough et al 2001); the western Peninsular Ranges batholith of Baja California shows a peak of tonalite magmatism between 115 Ma and 103 Ma (Tate & Johnson 2000). In California, plutonism in the east-central Sierra Nevada batholith is concentrated in a brief period from 102 Ma to 86 Ma (Coleman et al 2003); the voluminous Tuolumne intrusive suite of the Sierra Nevada batholith was emplaced between c. 95 Ma and c. 85 Ma (Coleman et al 2004). A newly recognized episode of cauldron subsidence with emplacement of rhyolite domes occurred from 107 Ma to 104 Ma in central British Columbia (Maclntyre & Villeneuve 2001). Tin-associated magmatism in the Khingan-Okhotsk belt of eastern Siberia, extending over 400 km, occurred episodically in a short period around c. 95 Ma (Sato et al 2002). The adakitic, Separation Point magmatic suite and associated orthogneisses, the Karamea batholith and the Paparoa batholith, all of the Median batholith of New Zealand were emplaced in the period 120-105 Ma (Muir et al 1997; 1998; Waight et al. 1998; Mortimer et al 1999). Kimberlite magmatism also peaked in the mid-Cretaceous, with a period of increased emplacement rate in Russia from 105 Ma to 95 Ma, in North America from 103 Ma to 94 Ma and in Africa from 116 Ma to 70 Ma (with a brief hiatus from 100 Ma to 95 Ma) (Heaman et al 2003).

Post-compressional extension In general, mid-Cretaceous compression is followed by extension or transtension, usually from c. 100 Ma on, and often on a spectacular scale, probably as a function of the high heat flow associated with a superplume event. In New Zealand, Marie Byrd Land and eastern Siberia, metamorphic core complexes formed (Muir et al. 1997; Toro et al. 2003; Siddoway et al. 2004; 2005). Ductile extensional structures formed from 105 Ma to 94 Ma following thrusting of the Peninsular Ranges batholith in southern California (Thomson & Girty 1994), transtensional basins formed in the Canadian

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Cordillera (Bassett & Kleinspehn 1996) and major extensional basins formed following widespread thrusting in Alaska (Till et al 1993).

most prolific of six periods of accumulation of effective oil source rocks during the Phanerozoic (Andrusevich et al 1998).

Geomagnetic field

Biotic changes

The Cretaceous long normal polarity interval of the geomagnetic field (e.g. Gallet & Hulot 1997) lasted from 125-84 Ma and is one of the key pieces of evidence for the mid-Cretaceous superplume event (Larson 199lb). There is some disagreement over the intensity of the geomagnetic field during the Cretaceous long normal polarity interval, which some models suggest should have been stronger than normal (Tarduno et al. 2001). Measurements of geomagnetic field intensity are difficult to make (McElhinny & Larson 2003) and complicated by post-preservational changes (Doubrovine & Tarduno 2004). Although new evidence suggests that the geomagnetic poloidal field intensity was reduced just prior to (Pan et al 2004) and during (Biggin & Thomas 2003) the Cretaceous long normal interval, other authors (Tarduno et al 2001; 2002) present convincing evidence that the field was up to three time stronger than in the preceding Late Jurassic or following Cenozoic.

Apart from an extinction event at the Cenomanian-Turonian boundary (c. 94 Ma) (Harries & Little 1999), in terms of disappearance of major groups, the Cretaceous long normal polarity interval (125-84 Ma) is generally associated only with higher rates of turnover in mineralized plankton taxa (Leckie et al 2002). Rather than going extinct, faunal and floral groups experienced major radiations in this period, which appears to have been the peak of the so-called 'Mesozoic Marine Revolution' (Vermeij 1995; Bambach 1999; Alroy 2004). The angiosperm and pollinating insect taxa went through a major expansion at this time (Grimaldi 1999; Lupia 1999; Kress et al 2001), as did several fern groups (Nagalingum et al 2002). Many groups of tetrapods radiated into their main modern groups, including frogs, turtles, lizards, snakes and birds (Benton 1996). Among marine invertebrates, speciation was particularly marked among some groups of bryozoans (Jablonski et al 1997) and a peak in bivalve diversity is also recognized (Checa & Jimenez-Jimenez 2003).

Atmosphere and oceans Short-lived periods of extreme oceanic anoxia are recorded for the mid-Cretaceous interval (Leckie et al 2002; Handoh & Lenton 2003). For example, the global Selli ocean anoxic event (OAE-la) occurred between 120.5 Ma and 119.5 Ma (Larson & Erba 1999), and the global Bonarelli OAE-2 occurred at the CenomanianTuronian boundary (c. 93.5 Ma) (Pancost et al 2004). Other events that are more restricted in distribution within Tethys occurred at 113-109 Ma (OAE-lb) and c. 99.5 Ma (OAEId) (Leckie et al 2002). Maximum sea surface temperatures were 3-5° warmer than today, but with pronounced variability that possibly triggered OAEs (Wilson & Norris 2001). Reef drowning is also characteristic of this period (Wilson et al 1998) and the mid-Cretaceous period saw some of the highest sea-levels in the Phanerozoic (Haq et al 1987), with highstands in the Early Aptian (c. 118 Ma), Late Albian (c. 102 Ma) and Turonian (93-90 Ma) (Strasser et al 2001). Episodes of anoxia and accumulation of carbon-rich sediments resulted in major accumulation of oil source rocks (e.g. Nzoussi-Mbassani et al 2003), including abundant coal offshore Nigeria (e.g. Obaje & Hamza 2000), and the mid-Cretaceous is the

Discussion and model In this section, some key similarities between Late Triassic-Early Jurassic and mid-Cretaceous deformation events will be outlined and a unified conceptual model for break-up of large plates will be presented.

Late Triassic-Early Jurassic deformation, Pangaea and Mesozoic oceans Episodic periods of coincident continental rifting and marginal collision have been described for the Tethyan Ocean from Late Permian to Late Eocene times, with significant events in the mid-Late Triassic and end-Early Cretaceous (Kazmin 1991; Ricou 1994). What singles out the Late Triassic-Early Jurassic deformation is that it occurred at a time when Pangaea was starting to break-up, and when extension, rather than compression, would have been expected to dominate plate kinematics. The pre-break-up configuration of Pangaea only came into being in the Late Permian (Muttoni et al 2003) and Pangaea break-up models have the added feature of needing to

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accommodate the presence of a partially intrasupercontinent ocean, Tethys (Fig. 1), with differently evolving margins in its eastern and western parts (Ricou 1994). The relationship between Triassic Palaeo-Tethys and the Pangaea-surrounding, Panthallasan Ocean (Fig. 1) is comparable to the relationship between the Indian and Pacific oceans today. In terms of plate motions, apparent polar wander paths for North America show significant changes in velocity at c. 200 Ma and c. 160 Ma (Beck & Housen 2003), the earlier of which would have coincided with Late Triassic-Early Jurassic deformation on the Pangaea margin. Recently, it has been shown that a global pulse of ophiolite obduction events also occurred in the Late Triassic-Early Jurassic (Vaughan & Scarrow 2003) (Fig. 3), which is consistent with the widespread evidence for Pangaea-margin compression at this time. Although it is not possible to determine a cause and effect relationship with Pangaea events, formation of the Pacific Plate, dated at 175-170 Ma (Bartolini & Larson 2001; Fisk & Kelley 2002), followed a major oceanic plate reorganization at c. 190 Ma (Nakanishi et al 1992). This coincided with initial break-up of Pangaea associated with increased subduction rates at its outer margins (Bartolini & Larson 2001). Mid-Cretaceous deformation and Cretaceous oceans Mid-Cretaceous compressional deformation is not restricted to the margins of the palaeoPacific basin (Fig. 2) and is developed extremely widely, even more widely so than summarized here. For example, the northern margins of Neo-Tethys in the Mediterranean region also show deformation of this age, for example in northern Spain (Monie et at. 1994), Corsica (Malavieille et al 1998), Slovakia (Plasienka 2003), Hungary (Haas & Pero 2004) and Greece and Bulgaria (Zagorcev 1994). Although not directly associated with initiation of continental break-up, mid-Cretaceous deformation is followed closely in several areas by episodes of high rates of extension, as outlined above and, in the case of New Zealand, by rifting-off and drifting of a continental fragment from Gondwana. Major reorganization of the Pacific Plate occurred at 120 Ma (Nakanishi & Winterer 1998; Sutherland & Hollis 2001) coincident with eruption of magmas associated with the formation of the Ontong-Java plateau among others (Kerr et at. 2000). Apparent polar wander paths for North America show

significant changes in velocity at c. 125 Ma and c. 88-80 Ma (Beck & Housen 2003) and the period 125-84 Ma shows ocean crust production rates that are 50% to 75% higher than the rates before and after this period (Larson 19915). Superplume-lithospheric plate interaction The following features unify Late TriassicEarly Jurassic and mid-Cretaceous deformation: (1) both occurred at a time of elevated mantle heat flow, as evidenced by concentrated episodes of voluminous magmatism on the continents and in the oceans; (2) both appear to have occurred at times of anomalously low reversal rate of the geomagnetic field (although this is less clear for the Late Triassic-Early Jurassic event), further implicating the mantle in whatever caused them; (3) both are followed by periods of unusually high rates of continental extension, Pangaea/Gondwana break-up in the Late Triassic-Early Jurassic case and extensional core-complex formation in the midCretaceous case; (4) both are associated with oceanic plate reorganization and major changes in plate velocity; (5) both occurred at times of substantial environmental and biotic change. This co-occurrence of extreme global magmatic, geomagnetic, environmental and biotic events points to one particular causal mechanism, involving simultaneous ascent of multiple hot plumes of material from the deep mantle, a 'superplume event' (Larson 19916; 1992). The global distribution of hotspot magmatism, involving substantial continental and oceanic expressions in Late Triassic-Early Jurassic and mid-Cretaceous times as outlined above, makes more local mechanisms, such as thermal blanketing (e.g. Anderson 1998) less likely to have been responsible. Figure 3 illustrates some of the events coeval with Late Triassic-Early Jurassic and mid-Cretaceous deformation and shows the distribution of proposed superplume and geomagnetic events, major tectonism, ophiolite obduction, sea-level variation and biotic and environmental changes through the Phanerozoic. The relationship between deformation and superplume events will be discussed in detail below; however, it is worth touching briefly on the relationship between the rise of hot plumes of mantle material to shallow levels and the magmatic and palaeoenvironmental changes outlined above. The relationship between magmatism and advection of heat to the upper mantle is the easiest to understand. As described above, flood volcanism, plutonism and kimberlite emplacement appear to be

Fig. 3. Plot of likely superplume-related events for the Phanerozoie (modified from Vaughan & Scarrow (2003)). Cumulative frequency plot of ophiolite age data from Vaughan & Scarrow (2003). Significant Phanerozoie geomagnetic events from Algeo (1996) and modified Phanerozoie sea-level curve from Algeo (1996) and Algeo & Seslavinsky (1995). Major tectonic events from Vaughan (1995) and Windley (1995), and pre-Mesozoic proxy data from Doblas et al (1998), Lehmann et al (1995), Hennessy & Mossman (1996) and Connolly & Miller (2002). Data and age ranges for Mid-Cretaceous and Palaeozoic superplume events shown (Larson 1991fl, 6). Asterisks mark major changes in plate velocity discussed in text. LIP, large igneous province.

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coeval in the Triassic-Jurassic and midCretaceous periods. For example, the widespread distribution of A-type, I-type and adakitic plutonism around the Pacific basin in the period 125-90 Ma, coeval with oceanic LIP and kimberlite emplacement, points to elevated heat flow over a large area of the globe. The situation for palaeoenvironmental proxy data in the Late Triassic-Early Jurassic and midCretaceous period and links with high rates of volcanism is less clear, and the subject of much discussion (e.g. Larson & Erba 1999; Hesselbo et al. 2002). Much current interest is focused on the possibility that small, initial magmatismrelated phenomena can trigger larger palaeoenvironmental changes (Tanner et al. 2004; Beerling & Berner 2002). The trigger mechanisms include thermally driven elevation in crustal heat flow (Jahren 2002), fall in sea-level (Paull et al 1991), CO2-driven atmospheric warming (Jenkyns etal 2001; Beerling & Berner 2002), or uplift of ocean floor (Jahren 2002). These are invoked to trigger catastrophic, massive dissociation of methane hydrate deposits in shallow-marine settings (Jenkyns et al 2001) or at continental high latitudes (Kidder & Worsley 2004), with consequent large rises in atmospheric temperature, changes in ocean and atmosphere composition, and biotic radiations and extinctions (Jahren 2002; Retallack 2002; Tanner et al 2004). Late Triassic-Early Jurassic and midCretaceous deformation post-date the start of their respective superplume events by only a few million years, whereas ophiolite obduction pulses follow up to 10-20 Ma later. Vaughan & Scarrow (2003) suggested that this delay may be a consequence of the time it takes to close a back-arc basin, the commonest setting for ophiolitic source rocks, following a change in plate boundary forces. The evidence for both episodes of deformation outlined above, and summarized in Tables 1 and 2, suggests that deformation was concentrated in a short time period, of the order of 5 Ma. Late TriassicEarly Jurassic deformation appears to have occurred between 202 Ma and 197 Ma and mid-Cretaceous deformation in two periods c. 116-110 Ma in the west palaeo-Pacific and c. 105-99 Ma in the east palaeo-Pacific, with both events possibly represented in northeast Siberia (Stavsky et al. 1990; Filatova & Vishnevskaya 1997). The geological record indicates an abrupt onset for superplume events and consequently there has been much discussion of a triggering mechanism, usually involving changes in mantle convection state and geomagnetic field behaviour following sudden transfer

of ponded cold subducted material into the deep mantle (Solheim & Peltier 1994; Tackley et al 1994; Condie 1998). Associated cooling at the core-mantle boundary is argued to be a possible cause of perturbations of the geomagnetic field (e.g. Muller 2002). One of the main, large-scale effects inferred for a superplume event is heating of the upper mantle and thermal elevation of the overlying lithosphere (Larson & Kincaid 1996) and, developing the model proposed by Vaughan (1995), this paper will argue that this is the key process responsible for superplume-related deformation.

'Oceanic' case Rising mantle plumes associate with the midCretaceous superplume event largely impacted oceanic lithosphere. Vaughan (1995) argued that for deformation associated with the midCretaceous superplume event, which will be refered to here as the 'oceanic' case (Fig. 4), thermal uplift and rejuvenation of ocean floor during multiple hot plume impact at the base of the lithosphere resulted in increased gravity sliding from a broad (c. 10 000 km across according to Larson (1991^)) topographic high which increased ridge push. Figure 4a shows the pre-superplume state with a notional large ocean with single spreading ridge and a back-arc basin floored by oceanic lithosphere. The subducting margins of the large ocean are in a state of mild tension, which is the common observed situation globally today (Hamilton 1994). In Figure 4a, this is a result of rollback of the old and cold subducting slab at the trench. The margins of the smaller, younger back-arc basin are in mild compression because of ridgepush forces. The commonest model for the mechanism that triggers superplume events involves accumulation of cold subducted slab material at the 670 km discontinuity or phase boundary (Tackley et al 1994; Machetel 2003) and this accumulation is depicted in Figure 4a. Figure 4b shows the onset and effects of a superplume event. The 670 km discontinuity is inferred to be an endothermic phase boundary, i.e. with a negative Clapeyron slope (e.g. Mambole & Fleitout 2002), in other words energy is required to cross the boundary downwards which means it acts as a barrier to further descent of upper mantle material into the lower mantle. Models for the onset of superplume events suggest that gravitational potential energy of accumulating cold subducted material is eventually sufficient to overcome the resistance of this phase boundary (Solheim & Peltier 1994) and a mantle avalanche is generated that

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Fig. 4. (a) Non-superplume state for notional section through the Earth with large ocean and marginal basin ('oceanic' case); (b) superplume state with active rising plumes from the D" layer. CMB, core-mantle boundary.

descends rapidly to the core-mantle boundary (Fig. 4b). This has two effects on mantle convection: an initial broad return flow that elevates the 670 km discontinuity and the upper mantle, followed by deep-rooted mantle plumes generated from increased lateral thermal variation at the core-mantle boundary (Larson & Kincaid 1996). This heating and elevation of the upper mantle triggers large-scale adiabatic decompression melting and magmatism (Larson & Kincaid 1996). The advected heat to the upper mantle has the effect of uplifting and thermally rejuvenating oceanic lithosphere over

a broad area and, by topographically elevating the ridge, increasing ridge-push force. A larger ridge-push force increases the convergence rate at subduction zones and thermal rejuvenation of oceanic lithosphere increases the degree of coupling between the subducting and overriding plate, both of which combine to produce oceanmarginal compression. In Figure 4b, this is expressed as, for example, back-arc basin closure and arc collision or ophiolite obduction (Vaughan & Scarrow 2003), but is inferred to be the general cause for mid-Cretaceous deformation and widespread terrane accretion.

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'Continental' case Figure 5 illustrates the situation for deformation associated with a superplume event where rising mantle plumes impact largely beneath continental lithosphere. Apart from some evidence for oceanic hotspot magmatism at 227 Ma (Lapierre et al 2003) and 210 Ma (Kerr et al 2000), most Late Triassic-Early Jurassic large igneous province magmatism and inferred mantle plume impacts affected the continents. Deformation associated with this putative superplume event is refered to as the 'continental' case. In the pre-superplume state, as before,

subducted oceanic lithosphere accumulates at the 670km discontinuity, beneath subduction zones on the edges of a supercontinent. In Figure 5a the supercontinental margins are in tension (Hamilton 1994), as for the 'oceanic' case. In Figure 5b a superplume event is inferred to initiate by the same process as outlined above, but with the variation that heat advection to, and uplift of, the upper mantle occurs beneath continental lithosphere. This results in adiabatic decompression melting and magmatism, as before, and generation of a broad, thermally generated topographic high. Gravitational sliding away from this high

Fig. 5. (a) Non-superplume state for notional section through the Earth with large supercontinent ('continental' case); (b) superplume state with active rising plumes from the D" layer and onset of supercontinental break-up. CMB, core-mantle boundary.

EPISODICITY OF TERRANE ACCRETION increases the coupling between continental lithosphere in the hanging wall of subduction zones and the subducting slab, and may result in overriding of oceanic lithosphere by continental lithosphere, both of which combine to generate supercontinent margin deformation. At the same time, gravitational sliding away from the topographic high triggers the onset of supercontinental break-up. The temporal association between the onset of break-up and continent marginal compression provides support for active models of supercontinent break-up involving mantle plumes (Storey 1995) because passive models require plate margins to be in a state of tension. The environment of Late Triassic-Early Jurassic deformation is almost certainly made more complex by the presence of an effectively intra-continental palaeoTethyan ocean. The evidence for hotspot magmatism in this ocean in the Late Triassic (Okay et al 2002) raises the possibility that a situation akin to the mid-Cretaceous 'oceanic' case may also have been active, subjecting Pangaea to a 'double-whammy' from the point of view of superplume-related deformation.

Summary - a unifying mechanism Oceanic and continental lithospheric plates appear to behave in similar ways when impacted from beneath by multiple plumes of hot rising mantle material. In terms of plate-marginal deformation in the Late Triassic-Early Jurassic and mid-Cretaceous, this can be summarized as the result of increased plate coupling at subduction zones during gravitational sliding of lithosphere away from broad, superplumegenerated, thermally supported topographic highs. Evidence for Early Jurassic (Nakanishi et al 1992) and late Early Cretaceous (Nakanishi & Winterer 1998; Sutherland & Hollis 2001) changes in palaeo-Pacific plate geometry, and Early Jurassic evidence for continental breakup or mid-Cretaceous periods of anomalous periods of extension, suggests that multiple plume impacts at the base of lithospheric plates are also likely to trigger plate reorganization. This is expressed as formation of new spreading ridges and the death of existing ones in the case of oceanic lithosphere, and continental breakup or high-rate extension in the case of continental lithosphere. Both can be described in terms of 'break-up'. 'Oceanic break-up' (Fig. 4b) is comparable to continental break-up, but expressed more subtly, because breaking-up oceanic lithosphere, i.e. forming new spreading ridges, simply creates more oceanic lithosphere. Both can also be described in terms of plate

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reorganization: continental break-up can be viewed as plate reorganization of continental lithosphere. However, this break-up/reorganization relationship is the fundamental point that links mid-Cretaceous and Late Triassic-Early Jurassic deformation and that points to a unifying mechanism for superplume-plate interactions. This can be summarized in the following sentence: when lithospheric plates are impacted from beneath by multiple plumes of hot rising mantle material they break-up/reorganize associated with gravitational spreading away from a broad, thermally generated topographic high and with a resulting short-lived pulse of plate-marginal deformation. This simple mechanism provides a unifying framework for continental and oceanic plate behaviour during the Late Triassic-Early Jurassic and mid-Cretaceous superplume events, explains episodicity of terrane accretion and provides strong supporting evidence for active models of supercontinent break-up.

Conclusions A review of evidence for deformation on the Late Triassic-Early Jurassic margins of Pangaea and the mid-Cretaceous margins of the palaeoPacific ocean, and examination of the relationships between plate-marginal deformation and coeval probable superplume events, results in a number of conclusions. 1.

2.

3.

4.

Late Triassic-Early Jurassic deformation is developed widely on the margins of Pangaea (expressed as the Indosinian Orogeny in Eurasia) and appears to have occurred in a short period between 202 Ma and 197 Ma. Late Triassic-Early Jurassic deformation was coeval with eruption of the Central Atlantic Magmatic Province, onset of Pangaea break-up, a longer than usual period of normal magnetic polarity in the Rhaetian-Hettangian and a major mass extinction event, all possible expressions of a superplume event. Mid-Cretaceous deformation occurred in two brief periods, the first from c. 116 Ma to c. 110 Ma in the west palaeo-Pacific and the second from c. 105 Ma to c. 99 Ma in the east palaeo-Pacific, with both events possibly represented in northeast Siberia and also developed on the northern margin of neo-Tethys. This deformation was coeval with eruption of major oceanic plateaux, extensional core-complex formation and rifting of New

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Bay-Larsen Inlet area, north-east Graham Land (with particular reference to the Trinity Peninsula Series). British Antarctic Survey Scientific Reports, 51. ALGEO, T.J. 1996. Geomagnetic polarity bias patterns through the Phanerozoic. Journal of Geophysical Research-Solid Earth, 101 (B2), 2785-2814. ALGEO, T.J. & SESLAVINSKY, K.B. 1995. The Paleozoic 5. world: continental flooding, hypsometry, and sealevel. American Journal of Science, 295, 787-822. ALROY, J. 2004. Are Sepkoski's evolutionary faunas dynamically coherent? Evolutionary Ecology Research, 6, 1-32. ANDERSON, D.L. 1998. The scales of mantle convection. Tectonophysics, 284, 1-17. ANDRUSEVICH, V.E., ENGEL, M.H., ZUMBERGE, I.E. & 6. BROTHERS, L.A. 1998. Secular, episodic changes in stable carbon isotope composition of crude oils. Chemical Geology, 152, 59-72. ARAKAWA, Y., SAITO, Y. & AMAKAWA, L. 2000. Crustal development of the Hida belt, Japan: evidence from Nd-Sr isotopic and chemical characteristics of igneous and metamorphic rocks. Tectonophysics, 328 , 183-204. BAMBACH, R.K. 1999. Energetics in the global marine fauna: a connection between terrestrial diversification and change in the marine biosphere. Geobios, 32, 131-144. BARRIO, C.A. & COFFIELD, D.Q. 1992. Late Cretaceous stratigraphy of the Upper Magdalena Basin in the Payande-Chaparral segment This is a contribution to the British Antarctic Survey (western Girardot Sub-Basin), Columbia. Journal Core Project 'Superterranes in the Pacific-margin arc' of South American Earth Sciences, 5, 123-139. (SPARC). The authors would like to thank Roger Larson, Brendan Murphy and Phil Leat for helpful BARTOLINI, A. & LARSON, R.L. 2001. Pacific microplate and the Pangea supercontinent in the Early and constructive reviews. to Middle Jurassic. Geology, 29, 735-738. BASSETT, K.N. & KLEINSPEHN, K.L. 1996. MidCretaceous transtension in the Canadian References Cordillera: evidence from the Rocky Ridge ABBOTT, D.H. & ISLEY, A.E. 2002. The intensity, volcanics of the Skeena group. Tectonics, 15, occurrence, and duration of superplume events 727-746. and eras over geological time. Journal of Geo- BECK, M.E., JR, BURMESTER, R., CEMBRANO, I, dynamics, 34, 265-307. DRAKE, R., GARCIA, A., HERVE, E & MUNIZAGA, ABERHAN, M. & FURSICH, F.T. 2000. Mass origination F. 2000. Paleomagnetism of the North Patagonian versus mass extinction: the biological contribatholith, southern Chile: an exercise in shape bution to the Pliensbachian-Toarcian extinction analysis. Tectonophysics, 326, 185-202. event. Journal of the Geological Society, London, BECK, M.E., JR & HOUSEN, B.A. 2003. Absolute 157, 55-60. velocity of North America during the Mesozoic ACTON, S.L., SIMONY, PS. & HEAMAN, L.M. 2002. from paleomagnetic data. Tectonophysics, 377, Nature of the basement to Quesnel Terrane near 33-54. Christina Lake, southeastern British Columbia. BEERLING, D.J. & BERNER, R.A. 2002. BiogeochemiCanadian Journal of Earth Sciences, 39, 65-78. cal constraints on the Triassic-Jurassic boundary ADAMS, CJ. & GRAHAM, I.J. 1996. Metamorphic and carbon cycle event. Global Biogeochemical tectonic geochronology of the Torlesse Terrane, Cycles, 16,1-12, 1036, 10.1029/2001GB001637. Wellington, New Zealand. New Zealand Journal BELOUSOVA, E.A., GRIFFIN, W.L., SHEE, S.R., JACKSON, of Geology and Geophysics, 39, 157-180. S.E. & O'REILLY, S.Y. 2001. Two age populations ADAMS, C.J., BARLEY, M.E., MAAS, R. & DOYLE, M.G. of zircons from the Timber Creek kimberlites, 2002. Provenance of Permian-Triassic volcaniNorthern Territory, as determined by laserclastic sedimentary terranes in New Zealand: ablation ICP-MS analysis. Australian Journal of evidence from their radiogenic isotope characEarth Sciences, 48, 757-765. teristics and detrital mineral age patterns. New BENTON, M.J. 1996. Testing the roles of competition Zealand Journal of Geology and Geophysics, 45, and expansion in tetrapod evolution. Proceedings 221-242. of the Royal Society of London, Series B, 263, AITKENHEAD, N. 1975. The geology of the Duse 641-646. Zealand from Gondwana, the Cretaceous normal polarity epoch and a major radiation of flowering plants and several animal groups (pollinating insects, birds, frogs, snakes, lizards, turtles, bryozoans and bivalves), all linked with the midCretaceous superplume event. Plate-marginal deformation in the Late Triassic-Early Jurassic and mid-Cretaceous can be modelled conceptually as the result of increased plate coupling at subduction zones during gravitational sliding of lithosphere away from superplume-generated, thermally supported topographic highs. A simple unifying mechanism for superplume-plate interactions, and support for active models of supercontinent break-up, can be summarized thus: when large oceanic or continental plates are impacted from beneath by multiple plumes of hot rising mantle material they break-up/reorganize associated with gravitational spreading away from a thermally generated topographic high and with a resulting shortlived pulse of plate-marginal deformation, often associated with terrane accretion.

EPISODICITY OF TERRANE ACCRETION BIGGIN, A.J. & THOMAS, D.N. 2003. Analysis of longterm variations in the geomagnetic poloidal field intensity and evaluation of their relationship with global geodynamics. Geophysical Journal International, 152, 392^15. BIRKENMAJER, K. 1994. Geology of Cretaceous magmatic rocks at Paradise Harbour, Danco Coast, Antarctic Peninsula. Studia Geologica Polonica, 104, 7-39. BORREGO, A.G., HAGEMANN, H.W., BLANCO, C.G., VALENZUELA, M. & DECENTI, C.S. 1996. The Pliensbachian (Early Jurassic) 'anoxic' event in Asturias, northern Spain: Santa Mera Member, Rodiles Formation. Organic Geochemistry, 25, 295-309. BRADSHAW, J.D. 1989. Cretaceous geotectonic patterns in the New Zealand region. Tectonics, 8, 803-820. BRADSHAW, J.D., ANDREWS, PB. & ADAMS, C.J. 1981. Carboniferous to Cretaceous on the Pacific margin of Gondwana: the Rangitata Phase of New Zealand. In: CRESWELL, M.M. & VELLA, P. (eds) Gondwana Five: Papers and Abstracts of Papers Presented at the Fifth International Gondwana Symposium. Balkema, Rotterdam, 217-221. BRADSHAW, J.D., WEAVER, S.D. & MUIR, R.J. 1996. Mid-Cretaceous oroclinal bending of New Zealand terranes. New Zealand Journal of Geology and Geophysics, 39, 461-468. BROWN, M., DIAZ, F & GROCOTT, J. 1993. Displacement history of the Atacama fault system 25W S-27W S, northern Chile. Geological Society of America Bulletin, 105, 1165-1174. BRUCE, R.M., NELSON, E.P, WEAVER, S.G. & Lux, D.R. 1991. Temporal and spatial variations in the southern Patagonian batholith: constraints on magmatic arc development. Geological Society of America Special Papers, 265, 1-12. BRYAN, S.E., CONSTANTINE, A.E., STEPHENS, C.J., EWART, A., SCHON, R.W & PARIANOS, J. 1997. Early Cretaceous volcano-sedimentary successions along the eastern Australian continental margin: Implications for the break-up of eastern Gondwana. Earth and Planetary Science Letters, 153, 85-102. BRYAN, S.E., EWART, A., STEPHENS, C.J., PARIANOS, J. & DOWNES, PJ. 2000. The Whitsunday Volcanic Province, Central Queensland, Australia: lithological and stratigraphic investigations of a silicicdominated large igneous province. Journal of Volcanology and Geothermal Research, 99,55-78. BURN, R.W. 1984. The geology of the LeMay Group, Alexander Island. British Antarctic Survey Scientific Reports, 109. BUTLER, R.F., GEHRELS, G.E. & SALEEBY, IB. 2001. Paleomagnetism of the Duke Island, Alaska, ultramafic complex revisited. Journal of Geophysical Research, 106 (B9), 19259-19269. BUTLER, S. & PELTIER, WR. 1997. Internal thermal boundary layer stability in phase transition modulated convection. Journal of Geophysical Research, 102, 2731-2749. BUTTERWORTH, P.J., CRAME, J.A., HOWLETT, P.J. &

167

MACDONALD, D.I.M. 1988. Lithostratigraphy of Upper Jurassic-Lower Cretaceous strata of eastern Alexander Island, Antarctica. Cretaceous Research, 9, 249-264. CAMPBELL, R.J. & HAIG, D.W 1999. Bathymetric change during Early Cretaceous intracratonic marine transgression across the northeastern Eromanga Basin, Australia. Cretaceous Research, 20, 403-446. CANTRILL, DJ. 2000. A new macroflora from the South Orkney Islands, Antarctica: evidence of an Early to Middle Jurassic age for the Powell Island Conglomerate. Antarctic Science, 12, 185-195. CARTER, R.M. 1988. Post-breakup stratigraphy of the Kaikoura synthem (Cretaceous-Cenozoic), continental margin, southeastern New Zealand. New Zealand Journal of Geology and Geophysics, 31, 405-429. CECCA, F. & MACCHIONI, F. 2004. The two Early Toarcian (Early Jurassic) extinction events in ammonoids. Lethaia, 37, 35-56. CENTENO-GARCIA, E. & KEPPIE, J.D. 1999. Latest Paleozoic early Mesozoic structures in the central Oaxaca terrane of southern Mexico: deformation near a triple junction. Tectonophysics, 301, 231-242. CHANG, E.Z. 1996. Collisional orogenesis between North and South China and its eastern extension in the Korean peninsula. Journal of Southeast Asian Earth Sciences, 13, 261-211. CHARDON, D. 2003. Strain partitioning and batholith emplacement at the root of a transpressive magmatic arc. Journal of Structural Geology, 25, 91-107. CHARRIER,R.,WYSS, A.R.,FLYNN, J.J. ETAL. 1996. New evidence for late Mesozoic early Cenozoic evolution of the Chilean Andes in the Upper Tinguiririca Valley (35° S), Central Chile. Journal of South American Earth Sciences, 9, 393-422. CHECA, A.G. & JIMENEZ-JIMENEZ, A.P 2003. Evolutionary morphology of oblique ribs of bivalves. Palaeontology, 46, 709-724. CHEN, G, GRAPES, R. & ZHANG, K. 2003. A model for Mesozoic crustal melting and tectonic deformation in Southeast China. International Geology Review, 45, 948-957. CHEN, W.S., YANG, H.C., WANG, X. & HUANG, H. 2002. Tectonic setting and exhumation history of the Pingtan-Dongshan Metamorphic Belt along the coastal area, Fujian Province, Southeast China. Journal of Asian Earth Sciences, 20, 829-840. CLARK, A.H., FARRAR, E., KONTAK, DJ. ET AL. 1990. Geologic and geochronological constraints on the metallogenic evolution of the Andes of southeastern Peru. Economic Geology and the Bulletin of the Society of Economic Geologists, 85,1520-1583. CLUZEL, D., ALLES, W & LAPIERRE, H. 1991. The Triassic-Jurassic calc-alkaline acidic intrusions of the Ogcheon Belt (South-Korea): continent-based arc or post-collisional magmatism. Comptes Rendus de VAcademic Des Sciences Serie II, 313, 435-442.

168

A. P. M. VAUGHAN & R. A. LIVERMORE

COFFIN, M.E & ELDHOLM, O. 1994. Large igneous provinces: crustal structure, dimensions, and external consequences. Reviews of Geophysics, 32, 1-36. COLE, K, BIRD, K.J., TORO, J., ROURE, F, O'SULLIVAN, P.B., PAWLEWICZ, M. & HOWELL, D.G. 1997. An integrated model for the tectonic development of the frontal Brooks Range and Colville Basin 250km west of the Trans-Alaska Crustal Transect. Journal of Geophysical Research-Solid Earth, 102 (B9), 20 685-20 706. COLEMAN, D.S., BRIGGS, A.S., GLAZNER, A.F. & NORTHRUP, C.J. 2003. Timing of plutonism and deformation in the White Mountains of eastern California. Geological Society of America Bulletin, 115, 48-57. COLEMAN, D.S., GRAY, W. & GLAZNER, A.F. 2004. Rethinking the emplacement and evolution of zoned plutons: geochronologic evidence for incremental assembly of the Tuolumne Intrusive Suite, California. Geology, 32, 433-436. CONDIE, K.C. 1998. Episodic continental growth and supercontinents: a mantle avalanche connection? Earth and Planetary Science Letters, 163, 97-108. CONDIE, K.C. 2002. The supercontinent cycle: are there two patterns of cyclicity? Journal of African Earth Sciences, 35, 179-183. CONNOLLY, S.R. & MILLER, A.I. 2002. Global Ordovician faunal transitions in the marine benthos: ultimate causes. Paleobiology, 28, 26-40. COOMBS, D.S., LANDIS, C.A., NORRIS, R.J., SINTON, J.M., BORNS, D.J. & CRAW, D. 1976. The Dun Mountain ophiolite belt, New Zealand, its tectonic setting, constitution, and origin, with special reference to the southern portion. American Journal of Science, 276, 561-603. COURTILLOT, V.E. & RENNE, PR. 2003. On the ages of flood basalt events. Comptes Rendus Geoscience, 335,113-140. Cox, R.T. & VAN ARSDALE, R.B. 2002. The Mississippi Embayment, North America: a first order continental structure generated by the Cretaceous superplume mantle event. Journal of Geodynamics, 34, 163-176. CRAME, J.A. & HOWLETT, PJ. 1988. Late Jurassic and Early Cretaceous biostratigraphy of the Fossil Bluff Formation, Alexander Island. British Antarctic Survey Bulletin, 78, 1-35. CROWHURST, P.V., MAAS, R., HILL, K.C, FOSTER, D.A. & FANNING, C.M. 2004. Isotopic constraints on crustal architecture and Permo-Triassic tectonics in New Guinea: possible links with eastern Australia. Australian Journal of Earth Sciences, 51, 107-122. CUNNINGHAM, W.D. 1995. Orogenesis at the southern tip of the Americas: the structural evolution of the Cordillera Darwin metamorphic complex, southernmost Chile. Tectonophysics, 244,197-229. CURTIS, M.L. 2001. Tectonic history of the Ellsworth Mountains, West Antarctica: reconciling a Gondwana enigma. Geological Society of America Bulletin, 113, 939-958. CURTIS, M.L. 2002. Palaeozoic to Mesozoic polyphase deformation of the Patuxent Range, Pensacola

Mountains, Antarctica. Antarctic Science, 14, 175-183. DALZIEL, I.W.D. 1981. Back-arc extension in the southern Andes: a review and critical appraisal. Philosophical Transactions of the Royal Society, London, Series A, 300, 319-395. DALZIEL, I.W.D. 1982. The early (pre-Middle Jurassic) history of the Scotia Arc region: a review and progress report. In: CRADDOCK, C. (ed.) Antarctic Science. University of Wisconsin Press, Madison, 111-126. DALZIEL, I.W.D. 1986. Collision and Cordilleran orogenesis: an Andean perspective. In: COWARD, M.P & RIES, A.C. (eds) Collision Tectonics. Geological Society, London, Special Publications, 19, 389-404. DEL VALLE, R.A. & MILLER, H. 2001. Transpressional deformation along the margin of Larsen Basin: new data from Pedersen Nunatak, Antarctic Peninsula. Antarctic Science, 13, 158-166. DELTEIL, J., STEPHAN, J.F, DE LEPINAY, B.M. & RUELLAN, E. 2003. Wrench tectonics flip at oblique subduction. A model from New Zealand. Comptes Rendus Geoscience, 335, 743-750. DICKINSON, W.R. & LAWTON,T.F. 2001. Carboniferous to Cretaceous assembly and fragmentation of Mexico. Geological Society of America Bulletin, 113, 1142-1160. DIXON, J. 1993. Regional unconformities in the Cretaceous of North-West Canada. Cretaceous Research, 14, 17-38. DIXON, J. 1997. Cretaceous stratigraphy in the subsurface of Great Slave Plain, southern Northwest Territories. Bulletin of Canadian Petroleum Geology, 45, 178-193. DOBLAS, M., OYARZUN, R., LOPEZ-RUIZ, J. ETAL. 1998. Permo-Carboniferous volcanism in Europe and northwest Africa: a superplume exhaust valve in the centre of Pangaea? Journal of African Earth Sciences, 26, 89-99. DOTT, R.H., JR, WINN, R.D., JR, DE WIT, M.J. & BRUHN, R.L. 1977. Tectonic and sedimentary significance of Cretaceous Tekenika Beds of Tierra del Fuego. Nature, 266, 620-622. DOUBLEDAY, P.A. & STOREY, B.C. 1998. Deformation history of a Mesozoic forearc basin sequence on Alexander Island, Antarctic Peninsula. Journal of South American Earth Sciences, 11,1-21. DOUBLEDAY, PA., MACDONALD, D.I.M. & NELL, P.A.R. 1993. Sedimentology and structure of the trench-slope to fore-arc basin transition in the Mesozoic of Alexander Island, Antarctica. Geological Magazine, 130, 737-754. DOUBROVINE, P.V. & TARDUNO, J.A. 2004. Selfreversed magnetization carried by titanomaghemite in oceanic basalts. Earth and Planetary Science Letters, 222, 959-969. DRAPER, G, GUTIERREZ, G. & LEWIS, J.F. 1996. Thrust emplacement of the Hispaniola peridotite belt: erogenic expression of the mid-Cretaceous Caribbean arc polarity reversal? Geology, 24, 1143-1146. ERNST, W.G. 1993. Metamorphism of Franciscan tectonostratigraphic assemblage, Pacheco Pass

EPISODICITY OF TERRANE ACCRETION area, east-central Diablo Range, California Coast Ranges. Geological Society of America Bulletin, 105, 618-636. EVENCHICK, C.A. & McNicoLL, V.J. 2002. Stratigraphy, structure, and geochronology of the Any ox Pendant, northwest British Columbia, and implications for mineral exploration. Canadian Journal of Earth Sciences, 39,1313-1332. FAURE, M, MARCHADIER, Y. & RANGIN, C. 1989. PreEocene synmetamorphic structure in the Mindoro-Romblon-Palawan Area, West Philippines, and implications for the history of Southeast-Asia. Tectonics, 8, 963-979. FAURE, M., NATALIN, B.A., MONIE, P., VRUBLEVSKY, A.A., BORUKAIEV, C. & PRIKHODKO, V. 1995. Tectonic evolution of the Anuy metamorphic rocks (Sikhote-Alin, Russia) and their place in the Mesozoic geodynamic framework of east Asia. Tectonophysics, 241, 279-301. FILATOVA, N.I. & VISHNEVSKAYA, V.S. 1997. Radiolarian stratigraphy and origin of the Mesozoic terranes of the continental framework of the northwestern Pacific (Russia). Tectonophysics, 269, 131-150. FISK, M. & KELLEY, K.A. 2002. Probing the Pacific's oldest MORB glass: mantle chemistry and melting conditions during the birth of the Pacific Plate. Earth and Planetary Science Letters, 202, 741-752. FODEN, I, SONG, S.H., TURNER, S., ELBURG, M., SMITH, P.B., VAN DER STELDT, B. & VAN PENGLIS, D. 2002. Geochemical evolution of lithospheric mantle beneath SE South Australia. Chemical Geology, 182, 663-695. GAETANI, M. 1997. The Karakorum Block in Central Asia, from Ordovician to Cretaceous. Sedimentary Geology, 109, 339-359. GALLET, Y. & HULOT, G. 1997. Stationary and nonstationary behaviour within the geomagnetic polarity time scale. Geophysical Research Letters, 24, 1875-1878. GEORGE, A.D. 1993. Radiolarians in off scraped seamount fragments, Aorangi Range, New Zealand. New Zealand Journal of Geology and Geophysics, 36, 185-199. GEORGE, A.D., MARSHALLSEA, S.J., WYRWOLL, K.H., CHEN, J. & Lu, Y.C. 2001. Miocene cooling in the northern Qilian Shan, northeastern margin of the Tibetan Plateau, revealed by apatite fission-track and vitrinite-reflectance analysis. Geology, 29, 939-942. GOLONKA, J. 2004. Plate tectonic evolution of the southern margin of Eurasia in the Mesozoic and Cenozoic. Tectonophysics, 381, 235-273. GRADSTEIN, F.M., OGG, J.G., SMITH, A.G & 36 OTHERS. 2004. A Geologic Time Scale 2004. Cambridge University Press, Cambridge. GRIMALDI, D. 1999. The co-radiations of pollinating insects and angiosperms in the Cretaceous. Annals of the Missouri Botanical Garden, 86, 373^06. GROCOTT, J. & TAYLOR, G.K. 2002. Magmatic arc fault systems, deformation partitioning and emplacement of granitic complexes in the Coastal

169

Cordillera, north Chilean Andes (25°30' S to 27W S). Journal of the Geological Society, London, 159, 425-442. GRUNOW, A.M., DALZIEL, I.W.D., HARRISON, T.M. & HEIZLER, M.T. 1992. Structural geology and geochronology of subduction complexes along the margin of Gondwanaland: new data from the Antarctic Peninsula and southernmost Andes. Geological Society of America Bulletin, 104, 1497-1514. GUEX, J., BARTOLINI, A., ATUDOREI, V. & TAYLOR, D. 2004. High-resolution ammonite and carbon isotope stratigraphy across the Triassic-Jurassic boundary at New York Canyon (Nevada). Earth and Planetary Science Letters, 225, 29-41. HAAS, J. & PERO, C. 2004. Mesozoic evolution of the Tisza Mega-unit. International Journal of Earth Sciences, 93, 297-313. HAIG, D.W. & LYNCH, D.A. 1993. A late early Albian marine transgressive pulse over northeastern Australia, precursor to epeiric basin anoxia: foraminiferal evidence. Marine Micropaleontology, 22, 311-362. HALL, L.S., LAMB, S.H. & MAC NIOCAILL, C. 2004. Cenozoic distributed rotational deformation, South Island, New Zealand. Tectonics, 23, 1-16, TC2002,10.1029/2002TC001421. HALLAM, A. & WIGNALL, P.B. 1999. Mass extinctions and sea-level changes. Earth-Science Reviews, 48, 217-250. HALPERN, M. & REX, D.C. 1972. Time of folding of the Yahgan formation and age of the Teckenika bed, southern Chile, South America. Geological Society of America Bulletin, 83, 1881-1886. HAMILTON, W.B. 1994. Subduction systems and magmatism. In: SMELLIE, J.L. (ed.) Volcanism associated with extension at consuming plate margins. Geological Society, London, Special Publications, 81, 3-28. HANDOH, I.C. & LENTON, T.M. 2003. Periodic midCretaceous oceanic anoxic events linked by oscillations of the phosphorus and oxygen biogeochemical cycles. Global Biogeochemical Cycles, 17, 1-13, 1092, 10.1029/2003GB002039. HANSEN, V.L. & DUSEL-BACON, C. 1998. Structural and kinematic evolution of the Yukon-Tanana upland tectonites, east-central Alaska: a record of late Paleozoic to Mesozoic crustal assembly. Geological Society of America Bulletin, 110, 211-230. HAQ, B.U, HARDENBOL, J. & VAIL, PR. 1987. Chronology of fluctuating sea-levels since the Triassic. Science, 235, 1156-1167. HARLOW, G.E., HEMMING, S.R., LALLEMANT, H.G.A., SISSON, VB. & SORENSEN, S.S. 2004. Two highpressure-low-temperature serpentinite-matrix melange belts, Motagua fault zone, Guatemala: a record of Aptian and Maastrichtian collisions. Geology, 32, 17-20. HARRIES, P.J. & LITTLE, C.T.S. 1999. The early Toarcian (Early Jurassic) and the CenomanianTuronian (Late Cretaceous) mass extinctions: similarities and contrasts. Palaeogeography, Palaeoclimatology, Palaeoecology, 154, 39-66.

170

A. P. M. VAUGHAN & R. A. LIVERMORE

HARRINGTON, HJ. & KORSCH, RJ. 1985. Late Permian to Cainozoic tectonics of the New England Orogen. Australian Journal of Earth Sciences, 32, 181-203. HARRIS, M.J., SYMONS, D.T.A., BLACKBURN, W.H., HART, C.J.R. & VILLENEUVE, M. 2003. Travels of the Cache Creek Terrane: a paleomagnetic, geobarometric and 40Ar/39Ar study of the Jurassic Fourth of July Batholith, Canadian Cordillera. Tectonophysics, 362,137-159. HATHWAY, B. 2000. Continental rift to back-arc basin: Jurassic-Cretaceous stratigraphical and structural evolution of the Larsen Basins, Antarctic Peninsula. Journal of the Geological Society, London, 157, 417-432. HEAMAN, L.M. & KJARSGAARD, B.A. 2000. Timing of eastern North American kimberlite magmatism: continental extension of the Great Meteor hotspot track? Earth and Planetary Science Letters, 178, 253-268. HEAMAN, L.M., KJARSGAARD, B.A. & CREASER, R.A. 2003. The timing of kimberlite magmatism in North America: implications for global kimberlite genesis and diamond exploration. Lithos, 71, 153-184. HENDERSON, R.A. 1998. Eustatic and palaeoenvironmental assessment of the mid- Cretaceous Bathurst Island Group of the Money Shoals platform, northern Australia. Palaeogeography, Palaeoclimatology, Palaeoecology, 138, 115-138. HENNESSY, J.F. & MOSSMAN, DJ. 1996. Geochemistry of Ordovician black shales at Meductic, southern Miramichi Highlands, New Brunswick. Atlantic Geology, 32, 233-245. HERVE, F. & FANNING, C.M. 2001. Late Triassic detrital zircons in meta-turbidites of the Chonos Metamorphic Complex, southern Chile. Revista Geologica de Chile, 28, 91-104. HERVE, F, FANNING, C.M. & PANKHURST, RJ. 2003. Detrital zircon age patterns and provenance of the metamorphic complexes of southern Chile. Journal of South American Earth Sciences, 16, 107-123. HESSELBO, S.P., GROCKE, D.R., JENKYNS, H.C., BJERRUM, C.J., FARRIMOND, P., BELL, H.S.M. & GREEN, O.R. 2000. Massive dissociation of gas hydrate during a Jurassic oceanic anoxic event. Nature, 406, 392-395. HESSELBO, S.P., ROBINSON, S.A., SURLYK, F. & PIASECKI, S. 2002. Terrestrial and marine extinction at the Triassic-Jurassic boundary synchronized with major carbon cycle perturbation: a link to initiation of massive volcanism? Geology, 30, 251-254. HESSELBO, S.P., MORGANS-BELL, H.S., MC£LWAIN, J.C., REES, P.M., ROBINSON, S.A. & Ross, C.E. 2003. Carbon-cycle perturbation in the Middle Jurassic and accompanying changes in the terrestrial paleoenvironment. Journal of Geology, 111, 259-276. HESSELBO, S.P, ROBINSON, S.A. & SURLYK, F. 2004. Sea-level change and facies development across potential Triassic-Jurassic boundary horizons, SW Britain. Journal of the Geological Society, London, 161, 365-379.

HOLDSWORTH, B.K. & NELL, P.A.R. 1992. Mesozoic radiolarian faunas from the Antarctic Peninsula: age, tectonic and paleoceanographic significance. Journal of the Geological Society, London, 149, 1003-1020. IL'INA, V.I. & SHURYGIN, B.N. 2000. The Ilansk formation and its stratotype (Lower Jurassic, southern Siberia). Geologiya I Geofizika, 41, 1195-1202. INESON, J.R. 1985. Submarine glide blocks from the Lower Cretaceous of the Antarctic Peninsula. Sedimentology, 32, 659-670. INESON, J.R. 1989. Coarse-grained submarine fan and slope apron deposits in a Cretaceous back-arc basin, Antarctica. Sedimentology, 36, 793-819. INGER, S., SCOTT, R.A. & GOLIONKO, B.G 1999. Tectonic evolution of the Taimyr Peninsula, northern Russia: implications for Arctic continental assembly. Journal of the Geological Society, London, 156, 1069-1072. ISOZAKI, Y. & ITAYA, T. 1990. Chronology of Sambagawa metamorphism. Journal of Metamorphic Geology, 8, 401-411. JABLONSKI,D. & CHALONER, W.G. 1994. Extinctions in the fossil record. Philosophical Transactions of the Royal Society of London Series B, 344,11-16. JABLONSKI, D., LIDGARD, S. & TAYLOR, P.D. 1997. Comparative ecology of bryozoan radiations: Origin of novelties in cyclostomes and cheilostomes. Palaios, 12, 505-523. JACQUES, J.M. 20030. A tectonostratigraphic synthesis of the Sub-Andean basins: implications for the geotectonic segmentation of the Andean Belt. Journal of the Geological Society, London, 160, 687-701. JACQUES, J.M. 2003/?. A tectonostratigraphic synthesis of the Sub-Andean basins: inferences on the position of South American intraplate accommodation zones and their control on South Atlantic opening. Journal of the Geological Society, London, 160, 703-717. JAHREN, A.H. 2002. The biogeochemical consequences of the mid-Cretaceous superplume. Journal of Geodynamics, 34, 177-191. JENKYNS, H.C., GROCKE, D.R. & HESSELBO, S.P. 2001. Nitrogen isotope evidence for water mass denitrification during the early Toarcian (Jurassic) oceanic anoxic event. Paleoceanography, 16, 593-603. JOHNSON, C. & GALLAGHER, K. 2000. A preliminary Mesozoic and Cenozoic denudation history of the North East Greenland onshore margin. Global and Planetary Change, 24, 261-274. JOHNSON, C.L. 2004. Polyphase evolution of the East Gobi basin: sedimentary and structural records of Mesozoic-Cenozoic intraplate deformation in Mongolia. Basin Research, 16, 79-99. JOHNSON, H.P., VANPATTEN, D., TIVEY, M. & SAGER, W.W. 1995. Geomagnetic polarity reversal rate for the Phanerozoic. Geophysical Research Letters, 22, 231-234. JOHNSON, S.E.,TATE, M.C. & FANNING, C.M. 1999. New geologic mapping and SHRIMP U-Pb zircon data in the Peninsular Ranges batholith, Baja

EPISODICITY OF TERRANE ACCRETION California, Mexico: Evidence for a suture? Geology, 27, 743-746. JOLIVET, M., BRUNEL, M., SEWARD, D. ET AL. 2001. Mesozoic and Cenozoic tectonics of the northern edge of the Tibetan plateau: fission-track constraints. Tectonophysics, 343, 111-134. JOURNEAY, J.M. & FRIEDMAN, R.M. 1993. The Coast Belt thrust system: evidence of Late Cretaceous shortening in southwest British Columbia. Tectonics, 12, 756-775. KAMEI, A., OWADA, M., HAMAMOTO, T, OSANAI, Y, YUHARA, M. & KAGAMI, H. 2000. Isotopic equilibration ages for the Miyanohara tonalite from the Higo metamorphic belt in central Kyushu, Southwest Japan: implications for the tectonic setting during the Triassic. Island Arc, 9, 97-112. KAMP, PJJ. 2000. Thermochronology of the Torlesse accretionary complex, Wellington region, New Zealand. Journal of Geophysical Research, 105 (B8), 19 253-19 272. KAZMIN, V.G. 1991. Collision and rifting in the Tethys Ocean: geodynamic implications. Tectonophysics, 196, 371-384. KEAR, D. & MORTIMER, N. 2003. Waipa Supergroup, New Zealand: a proposal. Journal of the Royal Society of New Zealand, 33, 149-163. KELLOGG, K.S. & ROWLEY, P.D. 1989. Structural geology and tectonics of the Orville Coast region, southern Antarctica Peninsula, Antarctica. US Geological Survey Professional Papers, 1498. KELLY, S.R.A., DOUBLEDAY, P.A., BRUNTON, C.H.C., DICKINS, J.M., SEVASTOPULO, G.D. & TAYLOR, P.D. 2001. First Carboniferous and ?Permian marine macrofaunas from Antarctica and their tectonic implications. Journal of the Geological Society, London, 158, 219-232. KENT, D.V. & OLSEN, P.E. 1999. Astronomically tuned geomagnetic polarity timescale for the Late Triassic. Journal of Geophysical Research, 104 (B6), 12 831-12 841. KERR, A.C., ITURRALDE-VINENT, M.A., SAUNDERS, A.D., BABBS,T.L. & TARNEY, J. 1999. A new plate tectonic model of the Caribbean: implications from a geochemical reconnaissance of Cuban Mesozoic volcanic rocks. Geological Society of America Bulletin, III, 1581-1599. KERR, A.C., WHITE, R.V. & SAUNDERS, A.D. 2000. LIP reading: Recognizing oceanic plateaux in the geological record. Journal of Petrology, 41, 1041-1056. KIDDER, D.L. & WORSLEY, T.R. 2004. Causes and consequences of extreme Permo-Triassic warming to globally equable climate and relation to the Permo-Triassic extinction and recovery. Palaeogeography, Palaeo climatology, Palaeoecology, 203, 207-237. KIMBROUGH, D.L., SMITH, D.P, MAHONEY, J.B. ET AL. 2001. Forearc-basin sedimentary response to rapid Late Cretaceous batholith emplacement in the Peninsular Ranges of southern and Baja California. Geology, 29, 491-494. KIMURA, G. 1997. Cretaceous episodic growth of the Japanese Islands. Island Arc, 6, 52-68. KIMURA, G, SAKAKIBARA, M. & OKAMURA, M. 1994.

171

Plumes in central Panthalassa: deductions from accreted oceanic fragments in Japan. Tectonics, 13, 905-916. KlRICHKOVA, A.I., KOSTINA, E.I. & TlMOSHINA, N.A.

2003. Comprehensive substantiation of Jurassic phytostratigraphy and correlations in the Kansk coal basin, Siberia. Stratigraphy and Geological Correlation, 11, 244-260. KIRICHKOVA, A.I. & KULIKOVA,N.K. 2002. Correlation of different-facies Triassic deposits of the eastern Urals and Siberia. Stratigraphy and Geological Correlation, 10, 488-502. KIRICHKOVA, A.I. & TRAVINA, T.A. 2000. Phytostratigraphy of Jurassic coal-bearing formations of the Irkutsk Basin. Stratigraphy and Geological Correlation, 8, 608-620. KIYOKAWA, S. 1992. Geology of the Idonnappu Belt, Central Hokkaido, Japan - evolution of a Cretaceous accretionary complex. Tectonics, 11, 1180-1206. KLEPEIS, K.A., CLARKE, G.L., GEHRELS, G. & VERVOORT, J. 2004. Processes controlling vertical coupling and decoupling between the upper and lower crust of orogens: results from Fiordland, New Zealand. Journal of Structural Geology, 26, 765-791. KOHN, M.J., SPEAR, F.S. & DALZIEL, I.W.D. 1993. Metamorphic P-T paths from Cordillera Darwin, a core complex in Tierra del Fuego, Chile. Journal of Petrology, 34, 519-542. KOPYLOVA, M.G, RUSSELL, J.K. & COOKENBOO, H. 1998. Upper-mantle stratigraphy of the Slave craton, Canada: insights into a new kimberlite province. Geology, 26, 315-318. KRAEMER, P.E. 2003. Orogenic shortening and the origin of the Patagonian orocline (56° S). Journal of South American Earth Sciences, 15, 731-748. KRESS, W.J., PRINCE, L.M., HAHN, WJ. & ZIMMER, E.A. 2001. Unravelling the evolutionary radiation of the families of the Zingiberales using morphological and molecular evidence. Systematic Biology, 50, 926-944. KUMAR, A., DAYAL, A.M. & PADMAKUMARI, V.M. 2003. Kimberlite from Rajmahal magmatic province: Sr-Nd-Pb isotopic evidence for Kerguelen plume derived magmas. Geophysical Research Letters, 30, 1-4, 2296, 10.1029/ 2003GL018518. KUSKY, T.M. & BRADLEY, D.C. 1999. Kinematic analysis of melange fabrics: examples and applications from the McHugh Complex, Kenai Peninsula, Alaska. Journal of Structural Geology, 21,1773-1796. LAIRD, M.G. & BRADSHAW, J.D. 2004. The break-up of a long-term relationship: the Cretaceous separation of New Zealand from Gondwana. Gondwana Research, 7, 273-286. LAPIERRE, H., JAHN, B.M., CHARVET, J. & Yu, Y.W. 1997. Mesozoic felsic arc magmatism and continental olivine tholeiites in Zhejiang province and their relationship with the tectonic activity in southeastern China. Tectonophysics, 274,321-338. LAPIERRE, H., BOSCH, D., TARDY, M. & STRUIK, L.C. 2003. Late Paleozoic and Triassic plume-derived

172

A. P. M. VAUGHAN & R. A. LIVERMORE

magmas in the Canadian Cordillera played a key role in continental crust growth. Chemical Geology, 201, 55-89. LARSON, R.L. 1991a. Geological consequences of superplumes. Geology, 19, 963-966. LARSON, R.L. 19916. Latest pulse of Earth: evidence for a mid-Cretaceous superplume. Geology, 19, 547-550. LARSON, R.L. 1992. Latest pulse of Earth: evidence for a mid-Cretaceous superplume and geological consequences of superplumes - Reply. Geology, 20, 476-477. LARSON, R.L. 1997. Superplumes and ridge interactions between Ontong Java and Manihiki plateaus and the Nova-Canton trough. Geology, 25, 779-782. LARSON, R.L. & ERBA, E. 1999. Onset of the midCretaceous greenhouse in the BarremianAptian: Igneous events and the biological, sedimentary, and geochemical responses. Paleoceanography, 14, 663-678. LARSON, R.L. & KINCAID, C 1996. Onset of midCretaceous volcanism by elevation of the 670 km thermal boundary layer. Geology, 24, 551-554. LAUDON,T.S. & FANNING, C.M. 2003. SHRIMP U-Pb age characteristics of detrital and magmatic zircons, eastern Ellsworth Land, Abstracts of 9th International Symposium of Antarctic Earth Sciences (Potsdam), Alfred-Wegener Stiftung, Berlin, 156. LAUDON, T.S., THOMSON, M.R.A., WILLIAMS, PL., MILIKEN, K.L., ROWLEY, P.D. & BOYLES, J.M. 1983. The Jurassic Latady Formation, southern Antarctic Peninsula. In: OLIVER, R., JAMES, PR. & JAGO, J.B. (eds) Antarctic Earth Science. Australian Academy of Science, Canberra, 398-414. LEBRON, M.C. & PERFIT, M.R. 1993. Stratigraphic and petrochemical data support subduction polarity reversal of the Cretaceous Caribbean island arc. Journal of Geology, 101, 389-396. LECKIE, R.M., BRALOWER, T.J. & CASHMAN, R. 2002. Oceanic anoxic events and plankton evolution: Biotic response to tectonic forcing during the mid-Cretaceous. Paleoceanography, 17, 1-12, 1041, 10.1029/2001PA000623. LEHMANN, D., BRETT, C.E., COLE, R. & BAIRD, G. 1995. Distal sedimentation in a peripheral foreland basin: Ordovician black shales and associated flysch of the western Taconic foreland, New York State and Ontario. Geological Society of America Bulletin, 107, 708-724. LEPVRIER, C., MALUSKI, H., VAN VUONG, N, ROGUES, D., AXENTE, V. & RANGIN, C. 1997. Indosinian NW-trending shear zones within the Truong Son belt (Vietnam) 40Ar/39Ar Triassic ages and Cretaceous to Cenozoic overprints. Tectonophysics, 283, 105-127. LEVERENZ, A. & BALLANCE, P.F. 2001. Terrane affiliation and terrane boundaries of Mesozoic accretionary complexes, northeastern North Island, New Zealand: some implications from recycled elastics. New Zealand Journal of Geology and Geophysics, 44, 589-599.

Li, X.H., CHEN, Z., Liu, D.Y. & Li, W.X. 2003. Jurassic gabbro-granite-syenite suites from southern Jiangxi province, SE China: age, origin, and tectonic significance. International Geology Review, 45, 898-921. LITTLE, C.T.S. & BENTON, M.J. 1995. Early Jurassic mass extinction: a global long-term event. Geology, 23, 495-498. LITTLE, T. A., MORTIMER, N. & MCWILLLAMS, M. 1999. An episodic Cretaceous cooling model for the Otago-Marlborough Schist, New Zealand, based on 40Ar/39Ar white mica ages. New Zealand Journal of Geology and Geophysics, 42, 305-325. LIVERMORE, R.A. & WOOLLETT, R.W. 1993. Sea-floor spreading in the Weddell Sea and southwest Atlantic since the Late Cretaceous. Earth and Planetary Science Letters, 117, 475^95. Lu, H., JIA, D., WANG, Z., Guo, L., SHI, Y. & ZHANG, Q. 1994. Tectonic evolution of the Dongshan Terrane, Fujian Province, China. Journal of South American Earth Sciences, 7, 349-365. LUNDIN, E.R. & DORE, A.G. 1997. A tectonic model for the Norwegian passive margin with implications for the NE Atlantic: Early Cretaceous to break-up. Journal of the Geological Society, London, 154, 545-550. LUPIA, R. 1999. Discordant morphological disparity and taxonomic diversity during the Cretaceous angiosperm radiation: North American pollen record. Paleobiology, 25, 1-28. LUYENDYK, B.P. 1995. Hypothesis for Cretaceous rifting of East Gondwana caused by subducted slab capture. Geology, 23, 373-376. MACDONALD, A.J., LEWIS, P.D., THOMPSON, J.F.H. ET AL. 1996. Metallogeny of an early to middle Jurassic arc, Iskut river area, northwestern British Columbia. Economic Geology and the Bulletin of the Society of Economic Geologists, 91, 1098-1114. MACDONALD, D.I.M. & BUTTERWORTH, PJ. 1990. The stratigraphy, setting and hydrocarbon potential of the Mesozoic sedimentary basins of the Antarctic Peninsula. In: ST JOHN, B. (ed.) Antarctica as an Exploration Frontier. Studies in Geology. American Association of Petroleum Geologists, 31, 102-125. MACDONALD, D.I.M., STOREY, B.C. & THOMSON, J.W. 1987. South Georgia. BAS GEOMAP Series, Sheet 1, 1:250000, Geological map and supplementary text. British Antarctic Survey, Cambridge. MACELLARI, C.E. 1988. Cretaceous palaeogeography and depositional cycles of western South America. Journal of South American Earth Sciences, 1, 373-418. MACHETEL, P. 2003. Global thermal and dynamical perturbations due to Cretaceous mantle avalanche. Comptes Rendus Geoscience, 335, 91-97. MAClNTYRE, D.G.

& VlLLENEUVE,

M.E.

2001.

Geochronology of mid-Cretaceous to Eocene magmatism, Babine porphyry copper district, central British Columbia. Canadian Journal of Earth Sciences, 38, 639-655.

EPISODICITY OF TERRANE ACCRETION MAHAN, K.H., HARTLEY, J.M., COLEMAN, D.S., GLAZNER, A.R & CARL, B.S. 2003. Sheeted intrusion of the synkinematic McDoogle pluton, Sierra Nevada, California. Geological Society of America Bulletin, 115, 1570-1582. MAHER, H.D. 2001. Manifestations of the Cretaceous High Arctic large igneous province in Svalbard. Journal of Geology, 109, 91-104. MALAVIEILLE, J., CHEMENDA, A. & LARROQUE, C. 1998. Evolutionary model for Alpine Corsica: mechanism for ophiolite emplacement and exhumation of high-pressure rocks. Terra Nova, 10, 317-322. MAMBOLE, A. & FLEITOUT, L. 2002. Petrological layering induced by an endothermic phase transition in the Earth's mantle. Geophysical Research Letters, 29, 1-4, 2044, 10.1029/2002GL014674. MARSHALLSEA, S.J., GREEN, P.P. & WEBB, J. 2000. Thermal history of the Hodgkinson Province and Laura Basin, Queensland: multiple cooling episodes identified from apatite fission track analysis and vitrinite reflectance data. Australian Journal of Earth Sciences, 47, 779-797.

173

MILLER, H. 1983. The position of Antarctica within Gondwana in the light of Palaeozoic erogenic development. In: OLIVER, R.L., JAMES, PR. & JAGO, J.B. (eds) Antarctic Earth Science. Australian Academy of Science, Canberra, 579-581. MONIE, P., SOLIVA, I, BRUNEL, M. & MALUSKI, H. 1994. Mylonitic shear zones in the Millas Granite (Pyrenees, France): 40Ar/39Ar Cretaceous ages and tectonic interpretation. Bulletin de la Societe Geologique de France, 165, 559-571. MORTIMER, N. 1986. Late Triassic, arc-related, potassic igneous rocks in the North American Cordillera. Geology, 14, 1035-1038. MORTIMER, N. 1987. The Nicola-Group - Late Triassic and Early Jurassic subduction-related volcanism in British-Columbia. Canadian Journal of Earth Sciences, 24, 2521-2536. MORTIMER, N. 1993. Jurassic tectonic history of the Otago Schist, New Zealand. Tectonics, 12, 237-244. MORTIMER, N. 2004. New Zealand's geological foundations. Gondwana Research, 1, 261-272. MORTIMER, N. & COOPER, A.F. 2004. U-Pb and SmMARTIN, M.W., KATO,TT, RODRIGUEZ, C, GODOY, E., Nd ages from the Alpine Schist, New Zealand. DUHART, P., MCDONOUGH, M. & CAMPOS, A. New Zealand Journal of Geology and 1999. Evolution of the late Paleozoic accretionary Geophysics, 47, 21-28. complex and overlying forearc-magmatic arc, MORTIMER, N. & PARKINSON, D. 1996. Hikurangi south central Chile (38° S-41° S): Constraints for Plateau: a Cretaceous large igneous province in the tectonic setting along the southwestern the southwest Pacific Ocean. Journal of Geophysmargin of Gondwana. Tectonics, 18, 582-605. ical Research, 101 (Bl), 687-696. MARZOLI, A., RENNE, PR., PICCIRILLO, E.M., MORTIMER, N, TULLOCH, A.J., SPARK, R.N., WALKER, ERNESTO, M., BELLIENI, G. & DE MIN, A. 1999. N.W., LADLEY, E., ALLIBONE, A. & KIMBROUGH, Extensive 200-million-year-old continental flood D.L. 1999. Overview of the Median batholith, basalts of the Central Atlantic Magmatic New Zealand: a new interpretation of the Province. Science, 284, 616-618. geology of the Median Tectonic Zone and MCELHINNY, M.W. & LARSON, R.L. 2003. Jurassic adjacent rocks. Journal of African Earth Sciences, dipole low defined from land and sea data. Eos, 29, 257-268. 84, 362-366. MUIR, R.J., IRELAND, T.R., WEAVER, S.D., BRADSHAW, McNuLTY, B.A. 1995. Shear zone development during J.D., WAIGHT,T.£., JONGENS, R. & EBY, G.N. 1997. magmatic arc construction: the Bench Canyon SHRIMP U-Pb geochronology of Cretaceous shear zone, central Sierra Nevada, California. magmatism in northwest Nelson-Westland, Geological Society of America Bulletin, 107, South Island, New Zealand. New Zealand Journal 1094-1107. of Geology and Geophysics, 40, 453-463. MEGARD, F. 1984. The Andean erogenic period and its MUIR, R.J., IRELAND, T.R., WEAVER, S.D., BRADSHAW, major structures in central and northern Peru. J.D., EVANS, J.A., EBY, G.N. & SHELLEY, D. 1998. Journal of the Geological Society, London, 141, Geochronology and geochemistry of a Mesozoic 893-900. magmatic arc system, Fiordland, New Zealand. MENEILLY, A.W. 1983. Deformation of granitic plutons Journal of the Geological Society, London, 155, in eastern Palmer Land. British Antarctic Survey 1037-1052. Bulletin, 61, 75-79. MULLER, R.A. 2002. Avalanches at the core-mantle MENEILLY, A.W. 1988. Reverse-fault step at Engel boundary. Geophysical Research Letters, 29,1-4, Peaks, Antarctic Peninsula. Journal of Structural 1935,10.1029/2002GL015938. Geology, 10, 393-403. MUTTONI, G., KENT, D.V., GARZANTI, E., BRACK, P., MENG, Q.R. & ZHANG, G.W 2000. Geologic framework ABRAHAMSEN, N. & GAETANI, M. 2003. Early and tectonic evolution of the Qinling orogen, Permian Pangea 'B' to Late Permian Pangea 'A. central China. Tectonophysics, 323,183-196. Earth and Planetary Science Letters, 215,379-394. METCALFE, I. 2000. The Bentong-Raub Suture Zone. NAGALINGUM, N.S., DRINNAN, A.N., LUPIA, R. & Journal of Asian Earth Sciences, 18, 691-712. MCLOUGHLIN, S. 2002. Fern spore diversity and MILLAR, I.L., PANKHURST, R.J. & FANNING, C.M. 2002. abundance in Australia during the Cretaceous. Basement chronology of the Antarctic Peninsula: Review of Palaeobotany and Palynology, 119, recurrent magmatism and anatexis in the Palaeo69-92. zoic Gondwana Margin. Journal of the Geological NAKANISHI, M. & WINTERER, E.L. 1998. Tectonic Society, London, 159, 145-157. history of the Pacific-Farallon-Phoenix triple

174

A. P. M. VAUGHAN & R. A. LIVERMORE

junction from late Jurassic to early Cretaceous: an abandoned Mesozoic spreading system in the central Pacific basin. Journal of Geophysical Research, 103 (B6), 12453-12468. NAKANISHI, M., TAMAKI, K. & KOBAYASHI, K. 1992. A new Mesozoic isochron chart of the northwestern Pacific Ocean: paleomagnetic and tectonic implications. Geophysical Research Letters, 19, 693-696. NEWELL, N.D. 1967. Revolutions in the history of life. Geological Society of America Special Paper, 89, 63-91. NICHOLS, GJ. & CANTRILL, DJ. 2002. Tectonic and climatic controls on a Mesozoic forearc basin succession, Alexander Island, Antarctica. Geological Magazine, 139, 313-330. NOBLE, S.R., ASPDEN, J.A. & JEMIELITA, R. 1997. Northern Andean crustal evolution: New U-Pb geochronological constraints from Ecuador. Geological Society of America Bulletin, 109, 789-798. NOMADE, S., POUCLET, A. & CHEN, Y. 2002. The French Guyana doleritic dykes: geochemical evidence of three populations and new data for the Jurassic Central Atlantic Magmatic Province. Journal of Geodynamics, 34, 595-614. NORRIS, R.J. & COOPER, A.F. 2003. Very high strains recorded in mylonites along the Alpine Fault, New Zealand: implications for the deep structure of plate boundary faults. Journal of Structural Geology, 25, 2141-2157. NORRIS, R.J., KOONS, P.O. & COOPER, A.F. 1990. The obliquely convergent plate boundary in the South Island of New Zealand: implications for ancient collision zones. Journal of Structural Geology, 12, 715-725. NZOUSSI-MBASSANI, P., DISNAR, J.R. & LAGGOUNDEFARGE, F. 2003. Organic matter characteristics of Cenomanian-Turonian source rocks: implications for petroleum and gas exploration onshore Senegal. Marine and Petroleum Geology, 20, 411-427. OBAJE, N.G. & HAMZA, H. 2000. Liquid hydrocarbon source-rock potential of mid-Cretaceous coals and coal measures in the Middle Benue Trough of Nigeria. International Journal of Earth Sciences, 89,130-139. OKAMOTO, K., SHINJOE, H., KATAYAMA, I.,TERADA, K., SANO, Y. & JOHNSON, S. 2004. SHRIMP U-Pb zircon dating of quartz-bearing eclogite from the Sanbagawa Belt, south-west Japan: implications for metamorphic evolution of subducted protolith. Terra Nova, 16, 81-89. OKAY, A.I., MONOD, O. & MONIE, P. 2002. Triassic blueschists and eclogites from northwest Turkey: vestiges of the Paleo-Tethyan subduction. Lithos, 64, 155-178. O'SULLIVAN, P.B., FOSTER, D.A., KOHN, B.P. & GLEADOW, A.J.W. 1996. Multiple postorogenic denudation events: an example from the eastern Lachlan fold belt, Australia. Geology, 24, 563-566. O'SULLIVAN, P.B., MITCHELL, M.M., O'SULLIVAN, A.J., KOHN, B.P. & GLEADOW, A.J.W. 2000. Thermo-

tectonic history of the Bassian Rise, Australia: implications for the breakup of eastern Gondwana along Australia's southeastern margins. Earth and Planetary Science Letters, 182, 31-47. OTTO, S.C. 1997. Mesozoic-Cenozoic history of deformation and petroleum systems in sedimentary basins of Central Asia: implications of collisions on the Eurasian margin. Petroleum Geoscience, 3, 327-341. PALFY, J. & SMITH, PL. 2000. Synchrony between Early Jurassic extinction, oceanic anoxic event, and the Karoo-Ferrar flood basalt volcanism. Geology, 28, 747-750. PAN, Y.X., HILL, M.J., ZHU, R.X. & SHAW, J. 2004. Further evidence for low intensity of the geomagnetic field during the early Cretaceous time: using the modified Shaw method and microwave technique. Geophysical Journal International, 157, 553-564. PANCOST, R.D., CRAWFORD, N., MAGNESS, S., TURNER, A., JENKYNS, H.C. & MAXWELL, J.R. 2004. Further evidence for the development of photic-zone euxinic conditions during Mesozoic oceanic anoxic events. Journal of the Geological Society, London, 161, 353-364. PANKHURST, R.J. & ROWLEY, P.D. 1991. Rb-Sr study of Cretaceous plutons from southern Antarctic Peninsula and eastern Ellsworth Land, Antarctica. In: THOMSON, M.R.A., CRAME, J.A. & THOMSON, J.W. (eds) Geological Evolution of Antarctica. Cambridge University Press, Cambridge, 387-394. PANKHURST, R.J., MILLAR, I.L., GRUNOW, A.M. & STOREY, B.C. 1993. The Pre-Cenozoic magmatic history of the Thurston Island crustal block, West Antarctica. Journal of Geophysical Research, 98 (B7), 11835-11849. PANKHURST, R.J., WEAVER, S.D., HERVE, F. & LARRONDO, P. 1999. Mesozoic-Cenozoic evolution of the North Patagonian batholith in Aysen, southern Chile. Journal of the Geological Society, London, 156, 673-694. PANKHURST, R.J., RILEY, T.R., FANNING, C.M. & KELLEY, S.P 2000. Episodic silicic volcanism in Patagonia and the Antarctic Peninsula: chronology of magmatism associated with the break-up of Gondwana. Journal of Petrology, 41, 605-625. PARKINSON, C.D., MIYAZAKI, K., WAKITA, K., BARBER, AJ. & CARSWELL, D.A. 1998. An overview and tectonic synthesis of the pre-Tertiary very-highpressure metamorphic and associated rocks of Java, Sulawesi and Kalimantan, Indonesia. Island Arc, 1, 184-200. PAULL, C.K., USSLER, W. & DILLON, WP. 1991. Is the extent of glaciation limited by marine gas hydrates? Geophysical Research Letters, 18, 432-434. PE-PIPER, G. & REYNOLDS, PH. 2000. Early Mesozoic alkaline mafic dykes, southwestern Nova Scotia, Canada, and their bearing on Triassic-Jurassic magmatism. Canadian Mineralogist, 38, 217-232. PINDELL, J.L. & BARRETT, S. 1990. Geological evolution of the Caribbean region; a plate tectonic perspective. In: DENGO, G. & CASE, J. (eds) The

EPISODICITY OF TERRANE ACCRETION Caribbean Region, The Geology of North America. Geological Society of America, Boulder, Colorado, H, 405-432. PLASIENKA, D. 2003. Development of basementinvolved fold and thrust structures exemplified by the Tatric-Fatric-Veporic nappe system of the Western Carpathians (Slovakia). Geodinamica Acta, 16, 21-38. PROKOPH, A., ERNST, R.E. & BUCHAN, K.L. 2004. Time-series analysis of large igneous provinces: 3500 Ma to present. Journal of Geology, 112, 1-22. Qi, J.Z., YUAN, S.S., Liu, Z.J. ET AL. 2004. U-Pb SHRIMP dating of zircon from quartz veins of the Yangshan gold deposit in Gansu Province and its geological significance. Acta Geologica Sinica-English Edition, 78, 443-451. QUINN, M.J., WRIGHT, I.E. & WYLD, S.J. 1997. Happy Creek igneous complex and tectonic evolution of the early Mesozoic arc in the Jackson Mountains, northwest Nevada. Geological Society of America Bulletin, 109, 461-482. RAPELA, C.W., PANKHURST, R.J. & HARRISON, S.M. 1992. Triassic Gondwana granites of the Gastre district, North Patagonian massif. Transactions of the Royal Society of Edinburgh-Earth Sciences, 83, 291-304. RETALLACK, GJ. 2002. Carbon dioxide and climate over the past 300 Myr. Philosophical Transactions of the Royal Society of London Series A, 360, 659-673. Ricou, L.E. 1994. Tethys reconstructed: plates, continental fragments and their boundaries since 260-Ma from Central America to south-eastern Asia. Geodinamica Acta, 7, 169-218. RIDING, IB. & CRAME, J.A. 2002. Aptian to Coniacian (Early-Late Cretaceous) palynostratigraphy of the Gustav Group, James Ross Basin, Antarctica. Cretaceous Research, 23, 739-760. RILEY, T.R. & KNIGHT, K.B. 2001. Age of pre-breakup Gondwana magmatism. Antarctic Science, 13, 99-110. RILEY, T.R., LEAT, P.T., PANKHURST, R.J. & HARRIS, C. 2001. Origins of large volume rhyolitic volcanism in the Antarctic Peninsula and Patagonia by crustal melting. Journal of Petrology, 42, 1043-1065. RILEY, T.R., MILLAR, I.L., WATKEYS, M.K., CURTIS, M.L., LEAT, P.T., KLAUSEN, M.B. & FANNING, CM. 2004. U-Pb zircon (SHRIMP) ages for the Lebombo rhyolites, South Africa: refining the duration of Karoo volcanism. Journal of the Geological Society, London, 161, 547-550. ROGER, R, MALAVIEILLE, I, LELOUP, P.H., CALASSOU, S. & Xu, Z. 2004. Timing of granite emplacement and cooling in the Songpan-Garze Fold Belt (eastern Tibetan Plateau) with tectonic implications. Journal of Asian Earth Sciences, 22, 465-481. ROLAND, N.W. & TESSENSOHN, F. 1987. Rennick faulting: an early phase of Ross Sea rifting. Geologisches Jahrbuch, B66, 203-229. ROSSETTI, F, LlSKER, F, STORTI, F & LAUFER, A.L.

2003. Tectonic and denudational history of the

175

Rennick Graben (North Victoria Land): Implications for the evolution of rifting between East and West Antarctica. Tectonics, 22, 1-22, 1016, 10.1029/2002TC001416. RUFFELL, A. 1995. Evolution and hydrocarbon prospectivity of the Brittany Basin (Western Approaches Trough), offshore north-west France. Marine and Petroleum Geology, 12, 387-407. SAKASHIMA, T, TERADA, K., TAKESHITA, T. & SANO, Y. 2003. Large-scale displacement along the Median Tectonic Line, Japan: evidence from SHRIMP zircon U-Pb dating of granites and gneisses from the South Kitakami and paleo-Ryoke belts. Journal of Asian Earth Sciences, 21, 1019-1039. SATO, K., VRUBLEVSKY, A.A., RODIONOV, S.M., ROMANOVSKY, N.P. & NEDACHI, M. 2002. Mid-

Cretaceous episodic magmatism and tin mineralization in Khingan-Okhotsk volcano-plutonic belt, Far East Russia. Resource Geology, 52,1-14. SCHERMER, E.R., STEPHENS, K.A. & WALKER, J.D. 2001. Paleogeographic and tectonic implications of the geology of the Tiefort Mountains, northern Mojave Desert, California. Geological Society of America Bulletin, 113, 920-938. SCHEUBER, E. & GONZALEZ, G. 1999. Tectonics of the Jurassic-Early Cretaceous magmatic arc of the north Chilean Coastal Cordillera (22° S-26° S): a story of crustal deformation along a convergent plate boundary. Tectonics, 18, 895-910. SCHOUTEN, S., VAN KAAM-PETERS, H.M.E., RIJPSTRA, W.I.C., SCHOELL, M. & DAMSTE, J.S.S. 2000. Effects of an oceanic anoxic event on the stable carbon isotopic composition of Early Toarcian carbon. American Journal of Science, 300, 1-22. SEPHTON, M.A., AMOR, K., FRANCHI, I.A., WIGNALL, P.B., NEWTON, R. & ZONNEVELD, J.P. 2002. Carbon and nitrogen isotope disturbances and an endNorian (Late Triassic) extinction event. Geology, 30,1119-1122. SIDDOWAY, C.S., BALDWIN, S.L., FITZGERALD, P.G., FANNING, C.M. & LUYENDYK, B.P. 2004. Ross Sea mylonites and the timing of intracontinental extension within the West Antarctic rift system. Geology, 32, 57-60. SIDDOWAY, C.S., SASS III, L.C. & ESSER, R.P. 2005. Kinematic history of the Marie Byrd Land terrane, West Antarctica: Direct evidence from Cretaceous mafic dykes. In: VAUGHAN, A.P.M., LEAT, PT. & PANKHURST, R.J. (eds) Terrane processes at the margin of Gondwana. Geological Society, London, Special Publications, 246, 417-438. SILVA-ROMO, G, ARELLANO-GIL, I, MENDOZAROSALES, C. & NIETO-OBREGON, J. 2000. A submarine fan in the Mesa Central, Mexico. Journal of South American Earth Sciences, 13, 429-442. SMELLIE, J.L. 1981. A complete arc-trench system recognized in Gondwana sequences of the Antarctic Peninsula region. Geological Magazine, 118, 139-159. SMELLIE, J.L., ROBERTS, B. & HIRONS, S.R. 1996. Very low- and low-grade metamorphism in the Trinity

176

A. P. M. VAUGHAN & R. A. LIVERMORE

Peninsula Group (Permo-Triassic) of northern Graham Land, Antarctic Peninsula. Geological Magazine, 133, 583-594. SMITH, C.A., SISSON, V.B., LALLEMANT, H.G.A. & COPELAND, P. 1999. Two contrasting pressuretemperature-time paths in the Villa de Cura blueschist belt, Venezuela: possible evidence for Late Cretaceous initiation of subduction in the Caribbean. Geological Society of America Bulletin, 111, 831-848. SOLHEIM, L.P & PELTIER, W.R. 1994. Avalanche effects in phase transition modulated thermal convection: a model of Earth's mantle. Journal of Geophysical Research, 99 (B4), 6997-7018. SPELL, T.L., MCDOUGALL, I. & TULLOCH, AJ. 2000. Thermochronologic constraints on the breakup of the Pacific Gondwana margin: The Paparoa metamorphic core complex, South Island, New Zealand. Tectonics, 19, 433^51. SPORLI, K.B. 1978. Mesozoic tectonics, North Island, New Zealand. Geological Society of America Bulletin, 89, 415-425. STAVSKY, A.P., CHEKHOVITCH, V.D., KONONOV, M.V. & ZONENSHAIN, L.P. 1990. Plate tectonics and palinspastic reconstructions of the AnadyrKoryak Region, northeast USSR. Tectonics, 9, 81-101. STEIN, G, LAPIERRE, H., MONOD, O., ZIMMERMANN, J.L. & VIDAL, R. 1994. Petrology of some Mexican Mesozoic-Cenozoic plutons: sources and tectonic environments. Journal of South American Earth Sciences, 7, 1-7. STEVENS, R.A. & ERDMER, P. 1996. Structural divergence and transpression in the Teslin tectonic zone, southern Yukon Territory. Tectonics, 15, 1342-1363. STOREY, B.C. 1995. The role of mantle plumes in continental breakup: case histories from Gondwanaland. Nature, 377, 301-308. STOREY, B.C. & KYLE, PR. 1997. An active mantle mechanism for Gondwana breakup. South African Journal of Geology, 100, 283-290. STOREY, B.C., THOMSON, M.R.A. & MENEILLY, A.W. 1987. The Gondwanian Orogeny within the Antarctic Peninsula: a discussion. In: McKENZiE, G.D. (ed.) Gondwana Six: Structure, Tectonics and Geophysics. Geophysical Monograph Series. American Geophysical Union, Washington DC, 41, 191-197. STOREY, B.C., HOLE, M.J., PANKHURST, R.J., MILLAR, I.L. & VENNUM, W. 1988. Middle Jurassic withinplate granites in West Antarctica and their bearing on the break-up of Gondwanaland. Journal of the Geological Society, London, 145, 999-1007. STOREY, B.C., BROWN, R.W., CARTER, A., DOUBLEDAY, PA., HURFORD, A.J., MACDONALD, D.I.M. & NELL, P.A.R. 1996. Fission-track evidence for the thermotectonic evolution of a MesozoicCenozoic fore-arc, Antarctica. Journal of the Geological Society, London, 153, 65-82. STOVBA, S.M. & STEPHENSON, R.A. 1999. The Donbas Foldbelt: its relationships with the uninverted Donets segment of the Dniepr-Donets Basin, Ukraine. Tectonophysics, 313, 59-83.

STRASSER, A., CARON, M. & GJERMENI, M. 2001. The Aptian, Albian and Cenomanian of Roter Sattel, Romandes Prealps, Switzerland: a high-resolution record of oceanographic changes. Cretaceous Research, 22, 173-199. SUTHERLAND, R. & HOLLIS, C. 2001. Cretaceous demise of the Moa plate and strike-slip motion at the Gondwana margin. Geology, 29, 279-282. SUTHERLAND, R., DAVEY, F. & BEAVAN, J. 2000. Plate boundary deformation in South Island, New Zealand, is related to inherited lithospheric structure. Earth and Planetary Science Letters, 177, 141-151. TACKLEY, P.J., STEVENSON, D.J., GLATZMAIER, G.A. & SCHUBERT, G. 1994. Effects of multiple phasetransitions in a 3-dimensional spherical model of convection in Earth's mantle. Journal of Geophysical Research, 99 (B8), 15 877-15 901. TANNER, L.H., LUCAS, S.G. & CHAPMAN, M.G. 2004. Assessing the record and causes of Late Triassic extinctions. Earth-Science Reviews, 65, 103-139. TARDUNO, J.A., COTTRELL, R.D. & SMIRNOV, A.V. 2001. High geomagnetic intensity during the midCretaceous from Thellier analyses of single plagioclase crystals. Science, 291,1779-1783. TARDUNO, J.A., COTTRELL, R.D. & SMIRNOV, A.V. 2002. The Cretaceous superchron geodynamo: Observations near the tangent cylinder. Proceedings of the National Academy of Sciences of the United States of America, 99,14 020-14 025. TARDY, M., LAPIERRE, H., FREYDIER, C. ET AL. 1994. The Guerrero suspect terrane (western Mexico) and coeval arc terranes (the Greater Antilles and the Western Cordillera of Colombia): a Late Mesozoic intraoceanic arc accreted to cratonal America during the Cretaceous. Tectonophysics, 230, 49-73. TATE, M.C. & JOHNSON, S.E. 2000. Sub volcanic and deep-crustal tonalite genesis beneath the Mexican peninsular ranges. Journal of Geology, 108, 721-728. TESSENSOHN, F. 1994. The Ross Sea region: structural interpretation in relation to the evolution of the Southern Ocean. Terra Antartica, 1, 553-558. THIELE, R. & PINCHEIRA, M. 1987. Tectonica transpresiva y movimiento de desgarre en el segmento sur de la Zona de la Falla Atacama, Chile. Revista Geologica de Chile, 31, 77-94. THOMAS, J.C., COBBOLD, PR., SHEIN, V.S. & LE DOUARAN, S. 1999. Sedimentary record of late Paleozoic to Recent tectonism in central Asia analysis of subsurface data from the Turan and south Kazak domains. Tectonophysics, 313, 243-263. THOMSON, B., AFTALION, M., MC!NTYRE, R.M. & RICE, C. 1995. Geochronology and tectonic setting of silicic dike swarms and related silver mineralization at Candelaria, western Nevada. Economic Geology and the Bulletin of the Society of Economic Geologists, 90, 2182-2196. THOMSON, C.N. & GIRTY, GH. 1994. Early Cretaceous intra-arc ductile strain in Triassic-Jurassic and Cretaceous continental-margin arc rocks, Peninsular Ranges, California. Tectonics, 13,1108-1119.

EPISODICITY OF TERRANE ACCRETION THOMSON, M.R.A. & TRANTER, T.H. 1986. Early Jurassic fossils from central Alexander Island and their geological setting. British Antarctic Survey Bulletin, 70, 23-39. THOMSON, S.N. & HERVE, F. 2002. New time constraints for the age of metamorphism at the ancestral Pacific Gondwana margin of southern Chile (42° S-52° S). Revista Geologica de Chile, 29, 255-271. TILL, A.B. & SNEE, L.W. 1995. 40Ar/39Ar evidence that formation of blueschists in continental-crust was synchronous with foreland fold-and-thrust belt deformation, western Brooks Range, Alaska. Journal of Metamorphic Geology, 13, 41-60. TILL, A.B., Box, S.E., ROESKE, S.M. & PATTON, W.W. 1993. Mid-Cretaceous extensional fragmentation of a Jurassic-Early Cretaceous compressional orogen, Alaska - Comment. Tectonics, 12, 1076-1081. TORO, J., AMATO, J.M. & NATAL'IN, B. 2003. Cretaceous deformation, Chegitun River area, Chukotka Peninsula, Russia: implications for the tectonic evolution of the Bering Strait region. Tectonics, 22, 1-17, 1021, 10.1029/2001TC001333. TORSVIK, T.H. & ANDERSEN, T.B. 2002. The Taimyr fold belt, Arctic Siberia: timing of prefold remagnetisation and regional tectonics. Tectonophysics, 352, 335-348. TRANTER, T.H. 1987. The structural history of the LeMay Group of central Alexander Island, Antarctic Peninsula. British Antarctic Survey Bulletin, 77, 61-80. TROUW, R.A.J., PASSCHIER, C.W., SIMOES, L.S.A., ANDREIS, R.R. & VALERIANO, CM. 1997. Mesozoic tectonic evolution of the South Orkney microcontinent, Scotia arc, Antarctica. Geological Magazine, 134, 383-401. TROUW, R.A.J., PASSCHIER, C.W., VALERIANO, C.M., SIMOES, L.S.A., PACIULLO, F.V.P. & RIBEIRO, A. 2000. Deformational evolution of a Cretaceous subduction complex: Elephant Island, South Shetland Islands, Antarctica. Tectonophysics, 319, 93-110. TULLOCH, A.J.,

KlMBROUGH, D.L.,

LANDIS,

C.A.,

MORTIMER, N. & JOHNSTON, M.R. 1999. Relationships between the Brook Street terrane and Median Tectonic Zone (Median batholith): evidence from Jurassic conglomerates. New Zealand Journal of Geology and Geophysics, 42, 279-293. UEDA, H., KAWAMURA, M. & NIIDA, K. 2000. Accretion and tectonic erosion processes revealed by the mode of occurrence and geochemistry of greenstones in the Cretaceous accretionary complexes of the Idonnappu Zone, southern central Hokkaido, Japan. Island Arc, 9, 237-257. UMHOEFER, P.J. & MILLER, R.B. 1996. Mid-Cretaceous thrusting in the southern Coast Belt, British Columbia and Washington, after strike-slip fault reconstruction. Tectonics, 15, 545-565. VAUGHAN,A.PM. 1995. Circum-Pacific mid-Cretaceous deformation and uplift - a superplume-related event? Geology, 23, 491-494. VAUGHAN, A.P.M. & SCARROW, J.H. 2003. Ophiolite

177

obduction pulses as a proxy indicator of superplume events? Earth and Planetary Science Letters, 213, 407^16. VAUGHAN, A.P.M. & STOREY, B.C. 2000. The eastern Palmer Land shear zone: a new terrane accretion model for the Mesozoic development of the Antarctic Peninsula. Journal of the Geological Society, London, 157,1243-1256. VAUGHAN, A.P.M., MILLAR, I.L. & THISTLEWOOD, L. 1999. The Auriga Nunataks shear zone: Mesozoic transfer faulting and arc deformation in northwest Palmer Land, Antarctica. Tectonics, 18, 911-928. VAUGHAN, A.P.M., KELLEY, S.P. & STOREY, B.C. 20020. Mid-Cretaceous ductile deformation on the Eastern Palmer Land Shear Zone, Antarctica, and implications for timing of Mesozoic terrane collision. Geological Magazine, 139, 465-471. VAUGHAN, A.P.M., PANKHURST, R.J. & FANNING, C.M. 2002ft. A mid-Cretaceous age for the Palmer Land event, Antarctic Peninsula: implications for terrane accretion timing and Gondwana palaeolatitudes. Journal of the Geological Society, London, 159, 113-116. VERMEIJ, GJ. 1995. Economics, volcanoes, and Phanerozoic revolutions. Paleobiology, 21, 125-152. VlELREICHER, R.M., VlELREICHER, N.M., HAGEMANN,

S.G. & JONES, G. 2003. Fault zone evolution and its controls on ore-grade distribution at the Jian Cha Ling gold deposit, western Qinling region, central China. Mineralium Deposita, 38, 538-554. VRY, J.K., BAKER, J., MAAS, R., LITTLE, T.A., GRAPES, R. & DIXON, M. 2004. Zoned (Cretaceous and Cenozoic) garnet and the timing of high grade metamorphism, Southern Alps, New Zealand. Journal of Metamorphic Geology, 22, 137-157. WAIGHT, T.E., WEAVER, S.D. & MUIR, R.J. 1998. MidCretaceous granitic magmatism during the transition from subduction to extension in southern New Zealand: a chemical and tectonic synthesis. Lithos, 45, 469-482. WANG, Z.H. 2002. The origin of the Cretaceous gabbros in the Fujian coastal region of SE China: implications for deformation-accompanied magmatism. Contributions to Mineralogy and Petrology, 144, 230-240. WANG, Z.H. & Lu, H.F. 1997.40Ar/39Ar geochronology and exhumation of mylonitized metamorphic complex in Changle-Nanao ductile shear zone. Science in China, Series D-Earth Sciences, 40, 641-647. WANG, Z.H. & Lu, H.F. 1998. Geology, petrology and geochemistry of the mafic-ultramafic rocks in the Fujian coastal region, southeastern China, and their genesis. Ofioliti, 23,1-6. WEILAND,T.J., SUAYAH, I.B. & FINCH, R.C. 1992. Petrologic, stratigraphic and tectonic significance of Mesozoic volcanic-rocks in the Rio-Wampu area, Eastern Honduras. Journal of South American Earth Sciences, 6, 309-325. WEVER, H.E., STOREY, B.C. & LEAT, P. 1995. Peraluminous granites in NE Palmer Land, Antarctic Peninsula: early Mesozoic crustal melting in a

178

A. P. M. VAUGHAN & R. A. LIVERMORE

magmatic arc. Journal of the Geological Society, London, 152, 85-96. WIGNALL, PB. 2001. Large igneous provinces and mass extinctions. Earth-Science Reviews, 53,1-33. WILSON, P.A. & NORRIS, R.D. 2001. Warm tropical ocean surface and global anoxia during the midCretaceous period. Nature, 412, 425-429. WILSON, P.A., JENKYNS, H.C., ELDERFIELD, H. & LARSON, R.L. 1998. The paradox of drowned carbonate platforms and the origin of Cretaceous Pacific guyots. Nature, 392, 889-894. WINDLEY, B.E 1995. The Evolving Continents. Wiley, Chichester. WYLD, S.J., ROGERS, J.W. & WRIGHT, IE. 2001. Structural evolution within the Luning-Fencemaker fold-thrust belt, Nevada: progression from backarc basin closure to intra- arc shortening. Journal of Structural Geology, 23,1971-1995. XIAO, WI, WINDLEY, B.E, CHEN, H.L., ZHANG, G.C. & Li, J.L. 2002. Carboniferous-Triassic subduction and accretion in the western Kunlun, China: Implications for the collisional and accretionary tectonics of the northern Tibetan Plateau. Geology, 30, 295-298. YANG, Z.Y., MOREAU, M.G., BUCHER, H., DOMMERGUES,

J.L. & TROUILLER, A. 1996. Hettangian and Sinemurian magnetostratigraphy from Paris Basin. Journal of Geophysical Research-Solid Earth, 101 (B4), 8025-8042. YONG, L., ALLEN, PA., DENSMORE, A.L. & QIANG, X. 2003. Evolution of the Longmen Shan Foreland Basin (Western Sichuan, China) during the Late Triassic Indosinian orogeny. Basin Research, 15, 117-138. ZAGORCEV, I.S. 1994. Structure and tectonic evolution of the Pirin-Pangaion structural zone (Rhodope Massif, southern Bulgaria and northern Greece). Geological Journal, 29, 241-268. ZHANG, X.H., Li, T.S. & Pu, Z.P. 2002. 40Ar/39Ar thermochronology of two ductile shear zones from Yiwulushan, West Liaoning Region: age constraints on the Mesozoic tectonic events. Chinese Science Bulletin, 47,1113-1118. ZlABREV, S.V., AlTCHISON, 1C, ABRAJEVITCH, A.V.,

BADENGZHU, DAVIS, A.M. & Luo, H. 2004. Bainang Terrane, Yarlung-Tsangpo suture, southern Tibet (Xizang, China): a record of intraNeotethyan subduction-accretion processes preserved on the roof of the world. Journal of the Geological Society, London, 161, 523-538.

New Zealand tectonostratigraphy and implications from conglomeratic rocks for the configuration of the SW Pacific margin of Gondwana A. M. WANDRES & J. D. BRADSHAW Department of Geological Sciences, University of Canterbury, Private Bag 4800, Christchurch, New Zealand (e-mail: anekant. [email protected]) Abstract: The active margin of Gondwana is presently preserved in the southwest Pacific region in the formerly continuous Gondwana fragments of Australia, Antarctica and New Zealand. The Phanerozoic tectonic history of New Zealand is interpreted in terms of progressive Pacific-ward growth by accretion of arc-trench systems and the basement rocks are described in terms of a number of volcano-sedimentary accreted terranes, suites and batholiths that intrude the terranes. The age of these basement rocks ranges from Early Cambrian to late Early Cretaceous. The origin of the magmatic and sedimentary rocks and the time of accretion of the New Zealand terranes to the Gondwana margin are important for the understanding of Phanerozoic Pacific tectonics. Geochronological research over the last decade on igneous rocks and conglomeratic units shows that the Tutoko Complex/Amundsen Province plutons are major contributors of detritus to the Pahau depositional basin and that the Antarctic sector of the Panthalassan Gondwana margin has to be (re)considered as the likely source for the Permo-Triassic Rakaia sediments. Igneous clast data have greatly improved understanding of the evolution of the New Zealand microcontinent and have put tighter constraints on its Mesozoic tectonic setting within the southwest Pacific margin of Gondwana.

Terrane-based terminology is used now very widely in the context of New Zealand rocks of mid-Cretaceous age and older and periodic overviews have been given over the last thirty years (Coombs et al 1976; Bradshaw et al 1981; Bradshaw 1989; Sporli & Ballance 1989; Mortimer 2004). The terranes conform to the basic definition of fault-bounded slices of regional scale, each with their own distinctive geological history. The terranes are thin slices from much larger entities, prompting a search for related rocks on other continents. It is clear that the terranes are allochthonous in the sense that very considerable volumes of rock have been displaced from between currently adjacent terranes. It is less clear which of the terranes are genuinely exotic and which are the results of large-scale strike-slip displacement with the same complex convergent margin. Mortimer et al. (2002) presented an interpretation of a marine geophysical transect that crosses most of the terrane boundaries and shows that the terranes make up the full thickness of the continental crust. The provenance of New Zealand terranes and, in particular, the Torlesse terranes has occupied geologists for many years (Landis & Bishop 1972; Andrews et al 1976; Coombs et al

1976; Howell 1980; MacKinnon 1983; Korsch & Wellman 1988; Bradshaw 1989; Mortimer 1995; Cawood et al 2002; Adams 2004). The time of accretion of the Median Tectonic Zone (MTZ) and all the Eastern Province terranes to the Gondwana margin is important for the understanding of the Palaeozoic-Mesozoic evolution of the southwest Pacific Gondwana margin. Over the last two decades the field of provenance analysis has undergone a revolution with the development of single-crystal isotope dating techniques. Many attempts were made to determine the provenance of sedimentary rocks using silt- to sand-sized single minerals (Ireland 1992; Pell et al 1997; Adams & Kelley 1998; Ireland et al 1998; Cawood et al 1999; Sircombe 1999; Pickard et al 2000). Coarse-grained rocks, notably conglomerates, have been used to trace sources of the Torlesse terranes (Wandres et al 20040, b). The aim of this paper is first to present a minireview of the New Zealand terranes and to examine the role and validity of the MTZ (or Median Batholith) in the context of the tectonostratigraphic framework of New Zealand and the geochronological record acquired from New Zealand rocks over the last decade. The second part summarizes and uses data from Cambrian

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246,179-216. 0305-8719/$15.00 © The Geological Society of London 2005.

Fig. 1. Present-day distribution of continental crust around the southwest Pacific margin. AP, Amundsen Province; RP, Ross Province. Inset: New Zealand microcontinent indicating the main features discussed in the text.

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to Early Cretaceous igneous clasts in conglomerates to reconstruct New Zealand and constrain the tectonic setting within the southwest Pacific margin of Gondwana.

Regional setting of New Zealand It is widely accepted that the terranes of New Zealand formed along Gondwana's Panthalassan plate margin, representing one of the most extensive orogenic belts in Earth history. The margin was established originally during Neoproterozoic rifting (Borg & De Paolo 1991) and became a long-lived site of subduction and plate convergence from the early Palaeozoic (e.g. Flottmann et al. 1993) until the midCretaceous (Bradshaw 1989; Luyendyk 1995). Accretion of arc-trench systems and, possibly, micro-continental fragments formed a series of eastward-younging orogenic belts (Fig. 1): the formerly continuous Neoproterozoic to Ordovician Ross-Delamerian Orogen (e.g. Stump et al. 1986; Coney et al 1990; Flottmann et al 1993), the Cambrian to Carboniferous Lachlan and Thomson orogens (e.g. Coney et al 1990; Foster & Gray 2000), the Silurian to Cretaceous New England Fold Belt (e.g. Powell & Li 1994) and the Permian to Cretaceous Rangitata Orogen (Suggate et al 1978). Remnants of these orogenic belts are preserved in the Gondwana fragments of Australia, Antarctica and New Zealand (Gibson & Ireland 1996, Fig. 1). Original relative positions in the New Zealand sector of the margin have been disrupted by the opening of the Tasman Sea and the Southern Ocean after 85 Ma (Wood 1994), followed by further disruption and re-configuration of the New Zealand microcontinent along the presently active plate boundary (Sutherland 1999). Some 90% of New Zealand continental crust lies beneath the sea (inset Fig. 1). Direct sampling of offshore basement is limited to a few dozen dredge hauls and oil exploration wells. The major bathymetric feature of western New Zealand is the Lord Howe Rise, which is the continental extension of the Challenger Plateau (Fig. 1). The northern region from the Norfolk Ridge to the Kermadec Trench to the north of the North Island is dominated by Cenozoic tectonics associated with subduction and back-arc basin development (Mortimer et al 1998). The Norfolk Ridge is the continental link between New Zealand and New Caledonia (Bade 1988; Cluzel et al 2001), where correlatives of the Murihiku, possibly the Brook Street (Fig. 2; Black 19960) and the Maitai terranes (Aitchison et al 1998) are found. The

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prominent offshore features of eastern New Zealand are the Hikurangi Plateau, the Chatham Rise, the Campbell Plateau and the Bounty Trough, a basin formed by crustal extension prior to break-up between New Zealand and West Antarctica. The Hikurangi Plateau lies mostly at depths of 2500-3500 m and is thought to be a Cretaceous large igneous province (LIP) similar to the Ontong Java Plateau and the Manihiki Plateau (Wood & Davy 1994; Mortimer & Parkinson 1996). The Chatham Rise and the Campbell Plateau are continental extensions of southern New Zealand (e.g. Lebrun et al 2003). Deformed and metamorphosed Torlesse basement rocks have been dredged from the crest of the Chatham Rise and are exposed on the Chatham Islands (Wood et al 1989; Ireland 1992; Wood & Herzer 1993). Western Province rocks occur on Campbell Island and in petroleum exploration drill holes on the Campbell Plateau (Cook et al 1999).

The role of the Median Tectonic Zone The New Zealand pre-Early Cretaceous basement rocks are traditionally grouped into three provinces, the Western Province, the MTZ and the Eastern Province (Coombs et al 1976; Bishop et al 1985; Bradshaw 1989; Fig. 2). The Western Province comprises two terranes and is made up largely of Lower Palaeozoic metasedimentary rocks cut by Devonian, Carboniferous and Early Cretaceous granitoids (Figs 3-6; Cooper 19890; Muir et al 1994; 1997; Waight et al 1997; Allibone & Tulloch 2004), with minor volcanic and metamorphic rocks of Cambrian age (Gibson & Ireland 1996; Miinker & Cooper 1997; 1999; Ireland & Gibson 1998). The Eastern Province consists of arc, forearc and accretionary complex rocks that relate to Permian to Cretaceous plate convergence. The MTZ separates the Western and Eastern provinces and consists of suites of Carboniferous to Early Cretaceous subduction-related calc-alkaline plutons with subordinate volcanic and sedimentary rocks (Kimbrough et al 19945; Muir et al 1998; Mortimer et al 19990). Muir et al (1995; 1997; 1998) separated the JurassicEarly Cretaceous MTZ plutons into the DarranSuite and the Separation Point Suite and Tulloch & Kimbrough (2003) termed the plutons LoSY (low Sr/Y ratio) and HiSY (high Sr/Y ratio) respectively (Figs 3, 5). Landis & Coombs (1967) proposed that the New Zealand Eastern and Western provinces were a paired metamorphic belt separated by a Median Tectonic Line (MTL). Cooper (19895)

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Fig. 2. Simplified tectonostratigraphic terrane map of New Zealand. Outliers of pre-Cretaceous rocks: NL, Northland allochthon; RP, Raukumara Peninsula; C, Clent Hill Group; Ky, Kyeburn and Horse Range Formation; B, Barretts Formation.

grouped the Western Province terranes (Buller and Takaka) into the Tuhua terrane and used the MTL to separate these rocks from the Eastern Province. Further research has shown

that the line is a broad zone (MTZ, Bradshaw 1993; Kimbrough et al. 1993) and that it plays an important role in tectonic reconstructions because it was immediately outboard of the

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Fig. 3. Geological map of Nelson showing the distribution of igneous and sedimentary rocks. palaeo-Gondwana margin. Geochronological research over the last ten years has identified magmatic activity in the MTZ that started in the Carboniferous and culminated in the Early Cretaceous (Mattinson et al. 1986; Kimbrough et al 1993; 19946; Muir et al 1995; 1997; 1998; Mortimer et al 19990; Hollis et al 2003; Allibone & Tulloch 2004; Brathwaite et al 2004). The plutons conform to the general definition of a batholith and Mortimer et al (19996) proposed the name Median batholith. Their definition included plutons such as the Separation Point batholith that cuts the Takaka terrane and had been interpreted previously as a stitching pluton between the Western Province and the MTZ (Muir et al 1995). The

term Median Tectonic Zone is probably misleading and the term Median batholith is not exactly equivalent. Recently published evidence (Kimbrough et al 1993; 19946; Muir et al 1995; 1997; 1998; Mortimer et al 19990; Clark et al 2000; Hollis et al 2003; Allibone & Tulloch 2004; Brathwaite et al 2004) indicates that the MTZ is best regarded as a magmatic arc complex that formed within, and adjacent to, the edge of the Western Province. It is proposed here that the rocks formerly assigned to the MTZ be regarded as part of the Western Province based on evidence summarized below. The name Tutoko Complex' is suggested after the highest peak in Fiordland. The present complex is probably a relict of an originally wider

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Araded sandstone, siltstone and

Fig. 4. Major rock units of the Buller and Takaka terranes of the Nelson region showing various tectonic events that have affected the region (modified after Rattenbury et al. 1998).

magmatic belt similar to, or continuous with, the Amundsen Province of West Antarctica (Bradshaw et al. 1997; Pankhurst et al. 19980). It includes both the non-metamorphosed Early Jurassic to Early Cretaceous plutonic and volcanic rocks of the old MTZ and the high pressure metamorphic equivalents in Fiordland (Fig. 5). The high pressure rocks imply thrusting and obduction within the complex (Clark et al. 2000; Daczko et al. 2002; Hollis et al. 2003; Klepeis et al. 2003; 2004).

Tutoko Complex (new name) -part of the Western Province Based on field observations and the presence of Carboniferous plutonic rocks in both the MTZ and the Western Province, Mortimer et al. (19990) suggested that Triassic to Cretaceous plutons of the MTZ were intruded along the autochthonous margin of Gondwana as the Median batholith. Although the chemistry of

the Carboniferous plutons of the Western Province (mainly A-type) differ in composition from the penecontemporaneous calc-alkaline Itype plutons of the MTZ, their relationship is (maybe) similar to that observed in the Early Cretaceous Electric Granite (A-type) and the Darran Suite (calc-alkaline I-type) of Fiordland (Fig. 5) - plutons that were emplaced in close proximity and under the same tectonic conditions. The model of Mortimer et al. (19990) contrasts with the allochthonous model of Muir et al. (1995; 1998), who proposed that the MTZ is itself a collided arc, with a back-arc basin between the arc and the Takaka terrane (Gondwana margin), despite the absence of evidence of such a basin. The latter model is based on the primitive Sr and Nd isotopic ratios of the Separation Point and Darran Suite granitoids and the conspicuous absence of inherited zircons, suggesting that these magmas have undergone little, if any, interaction with felsic crust and were emplaced outboard of the

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Fig. 5. Geological map of Fiordland and Stewart Island showing the distribution of igneous rocks (Stewart Island geology from Allibone & Tulloch 2004).

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Fig. 6. Time-space sequence across New Zealand also showing tectonics.

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Western Province. In their model, Muir et al. (1995) argued that the mafic Darran Suite rocks (LoSY, 157-131 Ma) have the appropriate chemical and isotopic compositions to generate, when subducted under the Western Province, the Western Fiordland Orthogneiss (WFO; at 125 Ma, HiSY) and the higher-level Separation Point plutons (at 117 ± 2 Ma, HiSY). The timing of emplacement of the Arthur River Complex (ARC, LoSY) at 136-129 Ma (Hollis et al. 2003), immediately following the last main phase of the Darran Suite magmatism and immediately prior to the emplacement of the WFO and Separation Point Suite (126-119 Ma and 124-111 Ma respectively, Mattinson et al. 1986; Muir et al. 1998) is consistent with a continuum of Jurassic to Cretaceous magmatic activity. Hollis et al. (2003) argued that the ARC was emplaced at mid-crustal levels and then buried to deeper crustal levels due to the convergence of the MTZ and the continental margin. The timing and change from calc-alkaline to adakitic magmatism in the Tutoko Complex, as noted by Wandres et al. (1998), coincided with the emplacement of the Darran plutons and contemporaneous adakitic sheets at 137 Ma, and magma loading was proposed as petrogenetic model for the adakites (Brown 1996; Wandres et al. 1998). The Arthur River Complex and the Western Fiordland Orthogneiss are juxtaposed but separated from Western Province metasedimentary rocks by the Anita Shear Zone (Hill 1995; Hollis et al. 2003) and the Doubtful Sound Shear Zone (Oliver 1980; Gibson et al. 1988), respectively (Fig. 5). However, the presence of Palaeozoic Orthogneiss entrained in the ARC requires that the ARC of the MTZ and the Western Province were adjacent when the ARC was emplaced. The Crow Granite (137 ± 3 Ma, Muir et al. 1997) and the Copperstain Creek Granodiorite (134 ± 1 Ma, Brathwaite et al. 2004) in Nelson (Fig. 3) are in intrusive contact with Western Province sediments, again indicating that the MTZ overlaped the Western Province at 137 Ma. Intrusive relationships between Carboniferous to Cretaceous plutons from the Western Province and MTZ are best preserved on Stewart Island (Fig. 5), where basement rocks are relatively unaffected by Cenozoic tectonism (Fig. 2). The Deceit Granite is in intrusive contact with the Carboniferous Knob Granite (c. 145 Ma and c. 305 Ma, respectively; Allibone & Tulloch 2004), the latter intruding the Western Province, showing a tectonic linkage at c. 145 Ma. A tectonic linkage is further

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supported by isotope data from conglomerate clasts reported by Wandres et al. (2004fc) and will be discussed below.

Tectonostratigraphic framework of New Zealand Western Province (Tuhua sequence) The Western Province consists of two distinct, north-south-trending, Middle Cambrian to Lower Devonian tectonostratigraphic terranes - the Buller terrane to the west and Takaka terrane to the east (Cooper 19890). In Nelson, where the terranes are best developed, they are separated by the Anatoki Fault (Figs 2, 3). The two terranes have also been recognized in Dusky Sound, Fiordland (Fig. 5), were they are separated by the Old Quarry Fault (Ward 1986), and on Stewart Island (Allibone & Tulloch 2004, references therein). Cooper (19890) inferred a Middle Devonian terrane amalgamation and suggested the term Tuhua terrane for the composite unit. Both terranes show the same folding which pre-dates the 375 ± 5 Ma Karamea Suite granitoids (Muir et al. 199 19960; refer to Fig. 6 for dating methods). Buller terrane. The Buller terrane comprises a relatively uniform suite of latest Cambrian to Early Ordovician quartz-rich turbidites, the Greenland Group, that have a passive continental margin geochemical character (Roser et al. 1996). Detrital zircons (Ireland & Gibson 1998) show a typical Palaeozoic Gondwana margin assemblage, with major U-Pb zircon peaks at 500-600 Ma (Ross-Delamerian orogen) and 1000-1200 Ma ('Grenvillian'). Initial metamorphism and cleavage formation of the Greenland Group is dated at 450 ± 10 Ma (Adams 2004). Along the eastern margin of the terrane, the group is overlain by a Middle Ordovician graptolitic mudstone and quartzite succession (Golden Bay Group) (Fig. 4). A marginal-marine to shallow-marine, Early Devonian succession (Reefton Group, Fig. 3, Bradshaw 1995) rests in tectonic contact on the Greenland Group (Figs 4, 6). Takaka terrane. Cambrian rocks predominate in the western part and Ordovician to Devonian rocks in the eastern part of the Takaka terrane. In the western part, the Cambrian rocks form ten fault-bounded slices, each with a coherent internal stratigraphy that differs from those of adjacent slices (Rattenbury et al. 1998; Mtinker & Cooper 1999). Five discrete assemblages can be recognized in the Cambrian: (i) a Middle to

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Late Cambrian island arc (Devil River Volcanics Group) with primitive, probably intra-oceanic, geochemistry; (ii) arc-derived sediments (Haupiri Group, Miinker & Cooper 1999; Wombacher & Miinker 2000); (iii) Middle Cambrian, non-volcanic, quartzofeldspathic sediments that geochemically resemble the Greenland Group (Roser et al. 1996) and have the typical Gondwana margin zircon assemblage (Jongens et al. 2003); (iv) syn- to post-arc conglomerates; (v) intrusive sedimentary melange (Fig. 4). Detrital zircon patterns from the quartzofeldspathic sandstone and the melange are similar to the adjacent Duller terrane and other terranes from the Gondwana margin (Wysoczanski et al. 1997; Jongens et al. 2003). The age of the Devil River arc rocks is constrained by fossil ages from interfingering Haupiri Group sediments. The Devil River Volcanics Group includes the main arc rocks which show a progressive evolution from low-K to high-K magmas over time. The Group also includes basalt lavas of Mataki Volcanic Formation and shallow intrusions of Cobb Igneous Complex (Miinker & Cooper 1999). A U-Pb age of 515 ± 7 Ma was obtained for zircons from a plagiogranite that intrudes the Cobb Igneous Complex. Both the Mataki basalts and the Cobb Igneous Complex are considered to have back-arc affinities (Miinker & Cooper 1999) but, due to faulting, the original relationship of these units cannot be established. The quartzofeldspathic turbidites are confined to one fault-bounded slice. The late- to post-arc conglomerates are well developed (>1000 m thick) and recent research suggests that a significant proportion of the clasts were derived from a continental margin source (M. Gutjahr, pers. comm. 2002). The intrusive sedimentary melange is similar in composition to the quartzofeldspathic turbidites (Jongens et al. 2003). This whole assemblage appears to range from early Middle Cambrian to early Late Cambrian and Jongens et al. (2003) suggest that the origin and emplacement of the melange is evidence of the Ross-Delamerian Orogeny in New Zealand. Magmatic rocks in Fiordland with ages of 501 ± 8 and 481 ± 8 Ma (Figs 5 & 6, Gibson & Ireland 1996) have also been attributed to this orogeny. The latest Cambrian to Early Devonian rocks that form the eastern half of the Takaka terrane are nowhere seen in sedimentary contact with the older Cambrian rocks. They comprise a thick (>6 km) continent-derived, quartzose sandstone, mudstone and carbonate succession (Cooper 19890). Carbonate sediments are particularly abundant in the Ordovician and

quartzite is common in the Silurian (Fig. 4). This succession was folded in the Middle Devonian (Bradshaw 2000) and cut by the mafic Riwaka Complex at 377 ± 5 Ma (Muir et al. 19960). Both terranes have small outliers of postCarboniferous rocks. The Permo-Triassic Parapara Group rests on the Takaka terrane and consists of thermally metamorphosed conglomerates, sandstones and black slates (Figs 3 & 4; Landis & Coombs 1967; Smale et al. 1996). The Parapara sediments are compositionally distinct from contemporaneous strata in the Eastern Province (Rakaia terrane, see below). However, a U-Pb zircon geochronological study showed that the younger part is dominated by young detrital zircons with ages close to 240 Ma (Wysoczanski etal. 1997). The Topfer Formation is a non-marine Triassic sandstone resting on the Buller terrane and cut by the Kirwans Dolerite (Fig. 3). The Topfer Formation has been compared with the Beacon Supergroup and the dolerite with the Ferrar magmatic province of Gondwana (Mortimer et al. 1995).

Post-Cambrian magmatism in the Western Province Nelson. Post-Cambrian plutons within the Buller and Takaka terranes were emplaced during Middle Devonian, Carboniferous and Cretaceous times (Figs 3, 6). The Devonian Karamea Suite was emplaced in the Buller terrane over a short period (375 ± 5 Ma) and includes a range of T- and 'S'-type granitoids (Muir et al. 1994; 19960) and, probably, the coeval but more mafic Riwaka Complex of the eastern Takaka terrane (Muir et al. 19966). Small Carboniferous plutons in the Buller terrane include members of the Cape Foulwind Supersuite such as the Cape Foulwind and Windy Point granites (327 ± 6 Ma and 328 ± 4 Ma, Muir et al. 1994) which form isolated plutons of weakly per aluminous, high-K calcalkaline, A-type monzogranites (Muir et al. 19966). Other, apparently unrelated, Carboniferous plutons include the c. 310 Ma highly fractionated, metaluminous to weakly peraluminous high-K, calc-alkaline Toropuihi Granite (encountered in Toropuihi-1 drillhole, Cooper & Tulloch 1992; Mortimer et al. 1997) and the c. 300-330 Ma Paringa Tonalite (Hurley et al. 1962; Aronson 1965; Cooper & Tulloch 1992). Carboniferous rocks also occur on Pepin Island (Platform Gneiss, Echinus Granite, Kimbrough et al. 1993; Beresford et al. 1996). Kimbrough et al. (19940) reported a discordant U-Pb age of 157 ± 21 Ma for a migmatite

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from the Eraser Complex (Rattenbury 1991) and argued that these rocks are similar to those described from the Charlston Complex and the Victoria Ranges that have been interpreted as lower plate assemblages of metamorphic core complexes. The Early Cretaceous plutonic rocks of the Hohonu batholith were emplaced between 114 ± 2 Ma and 110 ± 2 Ma (Muir et al 1997; Waight et al. 1997). The Paparoa batholith forms the central region of a metamorphic core complex developed during mid-Cretaceous crustal extension (Tulloch & Kimbrough 1989; Spell et al. 2000) and has Cretaceous granitoids related to the Hohonu Suite in its core. The youngest of these, the Buckland Granite, is dated at 110 ± 2 Ma (Muir et al. 1997) and shows some ductile deformation (Tulloch & Kimbrough 1989). Late Cretaceous magmatism is represented by the Atype French Creek Granite (82 ± 2 Ma) which is broadly coincident with the opening of the Tasman Sea (Waight et al. 1997).

quartz diorite (230 Ma) and the Pahia Intrusives (207 Ma). K-feldspar Ar/Ar incremental heating ages indicate that most of the plutons in the Longwood Range had cooled below 175 °C by the Middle Jurassic (170-180 Ma) and experienced no subsequent reheating (Mortimer et al. 19990).

Fiordland. South of the Alpine Fault, three distinct geological regions are recognized in the Fiordland massif (Fig. 5) - the western and eastern belts and the SW Fiordland block (Oliver & Coggon 1979). The latter comprises mainly low-grade metasedimentary rocks and granitoids that are regarded as the faultdisplaced extension of the lower Palaeozoic terranes of the Western Province in the northwest of the South Island (Cooper 19890, Fig. 2). The eastern belt comprises mainly Mesozoic plutonic rocks, some of which form part of the Tutoko Complex. The Kellard Point Orthogneiss gives U-Pb zircon ages of 481 Ma (Gibson & Ireland 1996). Two metasedimentary cover rocks, the Townly calc-silicate (502 Ma) and a quartzofeldspathic schist (334 Ma), have been regarded as the correlatives of the Australian Delamerian Fold Belt and the Lachlan Fold Belt, respectively (Gibson & Ireland 1996). Late Devonian to Early Permian granitoids range in age from 358-291 Ma and include the Hauroko, Kakapo and Pomona granites, Poteriteri pluton and Roxburgh Tonalite (Kimbrough et al. 19945; Muir et al. 1998). There is a distinct magmatic lull in the Permian followed by minor Late Triassic magmatism with the emplacement of the Mistake Diorite (224 Ma and 226 ± 3 Ma, Kimbrough et al. 19945; Muir et al. 1998) and the Slip Hill Diorite (229 Ma, Kimbrough et al. 19946) in northern Fiordland. To the southeast in Southland the Longwood Complex (Kimbrough et al. 19946; Mortimer et al. 19996) consists of the Late Triassic Oraka Hybrids

The Tutoko Complex magmatism extended from Early Jurassic to Early Cretaceous times and consists of the Darran Suite (LoSY, 168-131 Ma, Kimbrough et al. 19946; Muir et al. 1998), the Western Fiordland Orthogneiss (HiSY, 126-116 Ma, Mattinson et al. 1986; Gibson et al. 1988; Gibson & Ireland 1995) and its upper crustal equivalent, the Separation Point Suite (124-111 Ma, Muir et al. 1997; 1998), as well as the Arthur River Complex (LoSY, 136-129 Ma, Hollis et al. 2003). The penecontemporaneous LoSY and HiSY plutons (Tulloch & Kimbrough 2003; Allibone & Tulloch 2004) of Stewart Island and the Bounty Island granite are included in the Tutoko Complex (Fig. 6).

Stewart Island. The Carboniferous peraluminous orthogneiss (Ridge and Table Hill c. 345 Ma, Allibone & Tulloch 1997; 2004) intruded the Western Province Pegasus Group schists south of the Escarpment Fault (Fig. 5). Carboniferous granitoids north of the Escarpment Fault (Ruggedy and Neck granitoids, Knob, Freds Camp, Big Glory and Forked Plutons) are metaluminous to mildly alkaline rocks whose character is inconsistent with the involvement of older Palaeozoic crust in their genesis (Allibone & Tulloch 2004).

Tutoko Complex

Nelson. The Late Jurassic to Early Cretaceous I-type, calc-alkaline plutonic rocks forming the Rotoroa Complex (156 Ma, Kimbrough et al. 1993) in Nelson (Fig. 3) are probably equivalent to the Darran Suite rocks in Eastern Fiordland (Muir et al. 1998). The Separation Point Batholith (Kimbrough et al. 1993; Muir et al. 1995) intrudes both the Rotoroa Complex and the Lower Palaeozoic rocks of the Takaka terrane at 117 ± 2 Ma. The Copperstain Creek Granodiorite (LoSY, 134 ± 1 Ma, Brathwaite et al. 2004) intrudes the Takaka terrane, and the Crow Granite (LoSY, 137 ± 3 Ma, Muir et al. 1997), emplaced adjacent to the Anatoki Fault (Fig. 3), is in intrusive contact with the Buller terrane Ordovician metasediments. Other isolated Separation Point-type plutons intruding the Buller terrane are the Gouland

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Granodiorite (119 ± 2 Ma, Muir et al 1997) and The focus of plutonism during the Early the Mount Olympus pluton (111 ± 2 Ma, Muir Cretaceous moved southwards into the Western et al. 1994). Province. The Escarpment Fault currently separates the LoSY plutons (north; c. 170-125 Ma) Fiordland. The bulk of the plutonic rocks in from the HiSY plutons (south, c. 125-105 Ma, eastern Fiordland (Fig. 5) are I-type, calc- Tulloch & Kimbrough 2003). alkaline granitoids and diorites that range from Middle Jurassic to Early Cretaceous in age Bounty Islands. Cook et al. (1999) placed the (168-137 Ma) and are collectively referred to as Bounty Islands (Fig. 1) tentatively within the the Darran Suite. The Darran Suite is cut by Median Tectonic Zone. The peraluminous, calcseveral plutons of the Separation Point Suite alkaline Bounty Island granite was dated at 194 that give ages of c. 124 Ma (Muir et al. 1998). ± 5 Ma (T. Ireland, pers. comm.). The Darran Complex (part of the Darran Suite) is also cut by adakitic pegmatites (Wandres et al. 1998) indistinguishable in age (137 ± 2 Ma, Muir Early Jurassic to Early Cretaceous et al. 1998) from the Darran Complex. The sedimentary sequences gabbroic to dioritic orthogneisses of the Arthur River Complex were emplaced between 136 Ma Five volcano-sedimentary units are recognized to 129 Ma and represent the youngest phase of within the Tutoko Complex: the Drumduan and magmatism associated with the Darran Suite Teetotal groups in Nelson, the Largs terrane and Loch Burn Formation in Eastern Fiordland, (Hollis et al. 2003). The rocks of the western belt of Fiordland and the Paterson Group on Stewart Island. (Fig. 5) have been interpreted as the deep These volcano-sedimentary rocks are of Jurassic crustal levels of a metamorphic core complex of or earliest Cretaceous age and rest unconCretaceous age (Gibson et al. 1988). They are formably on, or contain clasts of, older plutonic termed collectively the Western Fiordland rocks (e.g. Williams 1978; Johnston et al. 1987; Orthogneiss (HiSY) and comprise Early Johnston 1990; Bradshaw 1993; Kimbrough et Cretaceous granulite-facies rocks (core) that are al. 1993,19946; Mortimer etal 19990, b\ Tulloch overlain structurally by Palaeozoic amphibolite- et al. 1999). facies rocks (Gibson et al. 1988). The magmatic precursors of the orthogneisses were emplaced Nelson. The Teetotal Group (Fig. 3) includes at c. 120-130 Ma (Mattinson et al. 1986; Gibson the Rainy River Conglomerate that forms et al. 1988; Gibson & Ireland 1995; Muir et al. a narrow belt within the Tutoko Complex 1998). The granulite-facies rocks of the Western and comprises conglomerates with minor Fiordland Orthogneiss are considered to be the sandstone-siltstone-mudstone and sparse coal lower crustal equivalent of the Separation Point seams. It was considered originally to be cut by Suite (Muir et al. 1995). Granitic enclaves in the the 228+| Ma Buller Diorite, but Tulloch et al. Western Fiordland Orthogneiss have been (1999) showed that it contained Late Jurassic dated at 380 Ma (Bradshaw & Kimbrough granitoid clasts and rested unconformably on 1991), which appears to correlate with the the diorite and was cut by the later 147+f Ma Karamea Suite rocks (Fig. 3). In general, most One Mile Gabbronorite. The Drumduan Group of the strongly deformed orthogneisses and (Fig. 3) is made up of sandstone and mudstone paragneisses in the Western Province are with probable Jurassic plant fossils. It is of Cretaceous metamorphic core complexes volcanic provenance but lacks magmatic rocks. (Gibson et al. 1988; Kimbrough & Tulloch 1989; The Early Cretaceous Cable Granodiorite cuts Tulloch & Kimbrough 1989; Kimbrough et al. it (Kimbrough et al. 1993). 19940) and not older Precambrian basement as argued by Adams (1975). Fiordland. The Largs terrane (Fig. 5) of eastern Fiordland was thought previously to be of Stewart Island. Apart from Carboniferous gran- Permian or Triassic age based on intrusion by the itoids, Allibone & Tulloch (2004) identified Late Triassic Mistake Diorite. Parts, however, three major episodes of plutonism that occurred must be younger, as indicated by a 140 ± 2 Ma at c. 170 Ma, c. 150-130 Ma and 130-105 Ma U-Pb zircon date for the Largs Ignimbrite (Figs 5, 6). The northern limit of the Middle (Mortimer et al. 19990). In the same area, the Jurassic magmatism is the Freshwater Fault Loch Burn Formation may also be composite, System, whereas the latest Jurassic and earliest with both Early and Late Jurassic rocks indicated Cretaceous magmatism produced the basement by detrital zircon U-Pb ages (Fig. 5, Kimbrough north of the Freshwater Fault System (Fig. 5). et al. 19946, T. Ewing, pers. comm.).

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polyphase deformation and complex structure is typical and, where developed, schistosity commonly post-dates one or more phases of folding (Mortimer 1993). Bradshaw et al. (1981) ascribed the polyphase deformation of these terranes to the Rangitata Orogeny Phase I (Late Triassic) and Phase II (Early Cretaceous). The three terranes west of the Livingstone Fault (Brook Street, Murihiku and Dun MountainMaitai) are dominated by volcanoclastic sediEastern boundary of the Western Province ments and have been referred to as the Central The boundary between the Western Province Arc terranes (Campbell 20005) or the Hokonui and Eastern Province is largely fault-controlled Assemblage (Landis et al. 1999). Volcanoclastic and the process of amalgamation of the two composition, however, is also characteristic of provinces remains unclear. However, Mortimer some terranes east of the Livingstone Fault et al. (19990) proposed that accretion of the (Caples terrane, Chrystalls Beach Complex and Brook Street terrane to the Western Province Waipapa), but the bulk of eastern South Island (and Gondwana) occurred at 245-230 Ma, is made up of quartzofeldspathic sediments, based on isotopic compositions and intrusive which comprise the assemblage of Torlesse relationships between the two units in the terranes (Rakaia, Aspiring, Te Akatarawa, Longwood Range (Fig. 5). Dated granitoid Kakahu and Pahau) and the Esk Head Melange clasts from the Rainy River Conglomerate, (Landis et al. 1999). An alternative nomenclawhich lies within the Tutoko Complex in ture has been proposed that groups the Caples Nelson, and from the Barretts Formation and the Waipapa terranes with other quart(Fig. 2, see below) of the Brook Street terrane zofeldspathic terranes east of the Livingstone in Southland, constrain the depositional ages of Fault into a Torlesse superterrane, and all both units to be no older than c. 170 Ma previous Torlesse subterranes are accorded (Tulloch et al. 1999; Adams et al. 2002). The ages terrane status (Campbell 20005). and chemistry of the granitoid clasts are broadly compatible with derivation from rocks that are Brook Street terrane. The foundation of the now represented by Triassic plutons of the Brook Street terrane is an Early Permian Western Province (Tulloch et al. 1999). Early oceanic volcanic arc (Sivell & Rankin 1983) that Jurassic ages as young as 180 Ma have been includes a 14-16 km thick sequence of moderobtained too. Based on similarities in strati- ately metamorphosed submarine volcanics and graphic age, depositional characteristics, grani- volcanoclastics of mainly basaltic-andesitic toid clast ages and composition between the composition with minor rhyolitic and dacitic Rainy River Conglomerate and the Barretts lithologies. The arc succession is overlain by a Formation, Tulloch et al. (1999) suggested that largely non-volcanic Late Permian succession, they are broadly correlative and that they lie the Productus Creek Group, which includes collectively within a combined Brook Street thick limestone. The Brook Street terrane is cut terrane-Western Province before the Late by Late Permian plutons dated at 265 Ma Jurassic. (Tulloch et al. 1999) and 261 Ma (Kimbrough et al. 1992). The Permian rocks are overlain unconformably by the Jurassic Barretts Formation (Landis et al. 1999), a thick, predomiEastern Province - terranes west of the nantly sandstone unit which hosts conglomerate Livingston Fault lenses with large granite boulders that have The Eastern Province is an assemblage of yielded U-Pb (TIMS) ages of 237 Ma to 180 Ma accreted allochthonous terranes making up (Tulloch et al. 1999). In Southland, Triassic almost all of northern New Zealand and the strata of the Murihiku terrane (see below) have eastern part of the South Island (Fig. 2). The been thrust westward (Letham Ridge Thrust, terranes can be divided into two groups based Figs 2, 6) over the Barretts Formation (Landis on either structure or geochemistry and prove- et al. 1999). nance. The two structural groups are separated In eastern Fiordland and Nelson, the Brook by the Livingstone Fault (Fig. 2). To the south- Street terrane-Western Province contact is west of the fault, structures are relatively marked by Cenozoic faulting (Mortimer et al. simple, axial planes are steeply inclined and 1999a). The early Late Triassic Mistake Diorite schistosity is not developed. To the northeast, (Fig. 5) has been interpreted as intruding the

Stewart Island. The Paterson Group (Fig. 5) comprises a diverse range of intermediate to acid volcanic rocks and volcanoclastic conglomerate, sandstone and mudstone, variably metamorphosed and locally schistose. The Paterson Group includes rocks cut by a Late Jurassic meta-rhyolite dated at 146+^ Ma (Kimbrough et al 19945).

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volcanoclastic rocks of the Brook Street terrane (Williams & Harper 1978). The Late Permian Pourakino Trondhjemite, the Hekeia Gabbro and the Colac Granite in the Longwood Range have been assigned, based on Sr and Nd initial ratios, to the Brook Street terrane (Mortimer et al. 19990). These plutons are intruded by dioritic dykes and plutons of the Western Province (Mortimer et al. 19990). A Middle-Late Triassic amalgamation of the Brook Street terrane to the Western Province has been proposed (Mortimer et al. 19990). Murihiku terrane. The Murihiku terrane is the least structurally deformed and stratigraphically most coherent of the New Zealand basement terranes. It comprises a thick (15 km) succession of weakly metamorphosed, mainly volcanoclastic sediments (Ballance & Campbell 1993). Campbell et al. (2001) argued that the Permian rocks of the Kuriwao Group are in gradational contact with Triassic strata, suggesting that sedimentation in the Murihiku terrane extends from Late Permian to late Early Cretaceous. In Southland, the Permian to early Middle Triassic basal Malakoff Hill Group is marine in origin with epiclastic andesitic sandstone and siltstone containing abundant interbedded tuff. Petrographic and geochemical data, including Nd isotopes, imply derivation from a nearby oceanic arc that lacked an older crystalline basement (Frost & Coombs 1989; Landis et al. 1999). In contrast, the overlying Middle and Late Triassic Murihiku strata show compelling evidence for a more Andean-type volcanic arc source (Boles 1974; Frost & Coombs 1989). The Murihiku rocks are considered generally to have formed parallel to and along the Gondwana margin in a forearc basin setting (e.g. Coombs etal. 1976), although there is some evidence in favour of a back-arc location (Coombs et al. 1996). The rocks coarsen to the southwest and the Brook Street terrane was considered originally to be the source. This may be true for the earliest Murihiku rocks, but the abundance of airfall deposits points to a source that was volcanically active through Triassic and Jurassic times whereas significant volcanism in the Brook Street terrane is restricted to the Early Permian. This conclusion is supported by evidence from isotope, trace element and palaeomagnetic data that indicates that the terranes are unrelated to each other (Grindley et al. 1981; Ballance & Campbell 1993; Bradshaw 1994). It is possible, however, that the Barretts Formation (Brook Street) and Murihiku sediments shared a similar source and might have been in close proximity by the

Jurassic (Landis et al. 1999). Granitic boulders in the Late Triassic (Murihiku) Moeatoa Conglomerate on the North Island have a Rb-Sr whole-rock isochron age of 226 ± 6 Ma (Graham & Korsch 1990). They are very similar in composition to plutons from the Western Province (N. Mortimer, pers. comm. 2004) and similar in age to the older clasts in the Barretts Formation of the Brook Street terrane. A reduction of vitric volcanic debris in the Middle to Late Jurassic Murihiku sediments has been reported by Ballance et al. (1981) and Black et al. (1993). Dun Mountain-Maitai terrane. The Early Permian Dun Mountain Ophiolite Belt (DMOB) comprises a discontinuous exposure of mafic and ultramafic rocks up to 4 km thick which extends for a length of >1000 km. The DMOB has been dated at 285-275 Ma (Kimbrough et al. 1992) and a similar Ndisotope model age of 278 ± 4 Ma has been obtained from a plagiogranite (Sivell & McCulloch 2000). The Maitai Group, a 6 km thick, moderately metamorphosed, volcanoclastic, sedimentary succession, rests on the DMOB. Three distinct suites in the DMOB were recognized, based on petrological and geochemical data, by Sivell & McCulloch (2000), who inferred a forearc setting for the DMOB, suggesting that oceanic crust was entrapped above a subduction zone which may have driven eruptive activity in the Brook Street terrane. The lower 1000-1500 m of the Maitai Group is of Late Permian age and consists of three units: Upukeroa Breccia, Wooded Peak Limestone and Tramway Formation. They comprise redeposited sandstone and bioclastic carbonate lithologies dominated by molluscan prismatic calcite shell debris attributed to unidentified atomodesmatinid bivalves. Thick lenses of polymictic breccia and bioclastic limestone of the basal Maitai Group rest locally in primary depositional contact on deformed ophiolite within the DMOB. Apart from the lower 1000-1500 m, the bulk (Stephens Supergroup, 4500-5000 m, Aitchison et al. 1988) is Early to Middle Triassic in age (Campbell 2000&). A granite clast from the Maitai Group has been dated at 265 Ma (Kimbrough et al. 1992). A carbon isotope study by Krull et al. (2000) on the marine organic matter of the Maitai Group strongly indicates a high-palaeolatitude setting for the terrane and deposition at c. 400 m depth within a volcanic arc-related basin. The Livingstone Fault marks the tectonic eastern margin of the Maitai terrane. It separates the

NEW ZEALAND TECTONOSTRATIGRAPHY structurally simple arc terranes to the southwest from highly deformed Caples, Waipapa, and Torlesse terranes to the northeast.

Eastern Province - terranes east of the Livingston Fault Caples terrane. The Caples terrane of Otago crops out in an arcuate belt stretching from the east Otago coast to the Alpine Fault (Fig. 2). The terrane includes a subdivided sequence (Caples Group) in the west and an undifferentiated sequence (Tuapeka Group) in the east. Offset by the Alpine Fault, similar rocks occur in the Nelson area and are called the Pelorus Group. In Otago, the Caples terrane is metamorphosed, from prehnite-pumpellyite facies to lower greenschist facies (Bishop et al. 1976). The northern boundary of the Caples terrane is a cryptic suture that lies within the Otago Schist and has been defined geochemically (Mortimer & Roser 1992). At the northern end of the Remarkables Range, Central Otago, psammitic schists of the Caples terrane pass downwards through a 300 m thick structural transitional zone into pelitic schists of the Aspiring lithologic association (Cox 1991). Mortimer & Roser (1992) demonstrated that the Aspiring lithological association is part of the Rakaia terrane. Structural geometry and shear criteria indicate that the Caples terrane overthrusts the Rakaia terrane from the south and west (Cox 1991; Mortimer 1993). Sedimentation is inferred to have been as fan deposits in structurally controlled lower trenchslope basins and on the trench floor adjacent to an active arc (Turnbull 19790, b). Geochemical compositions indicate that the sediments were derived from a relatively evolved, calc-alkaline arc system (Roser et al. 1993), perhaps intermediate between oceanic and continental arcs (Roser & Cooper 1990). Comparison of the Caples samples with data from nearby terranes of similar age (Torlesse, Maitai, Murihiku and Brook Street) reveals consistent compositional differences, suggesting that provenance linkages are unlikely (Roser & Korsch 1986, 1988). The greater incidence of more mature sandstone in the east supports depositional models that propose eastwardincreasing cratonic Gondwana influence and a westwards-increasing arc influence on the Caples sedimentary provenance (MacKinnon 1983; Korsch & Wellman 1988). The Caples terrane is virtually unfossiliferous but contains allochthonous boulders of limestone consisting of atomodesmatinid shell debris (Turnbull

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1979a) and Kungurian (Early Permian) conodonts (Fischer 1998; Ford et al. 1999). The Middle Triassic (Ito et al. 2000) Chrystalls Beach-Brighton coastal block (Fig. 2), south of Dunedin, lies within the Caples terrane but is of distinctly different petrofacies from that of the Caples terrane. On the basis of age (Middle Triassic), petrography and geochemical data, the Chrystalls Beach-Brighton metasediments cannot be correlated with known formations of the Caples terrane (Coombs et al. 2000). The southern part of the complex contains Tethyan and non-Tethyan radiolarian faunas of possible Southern Hemisphere high-latitude origin (Ito et al. 2000). Torlesse terranes. The Torlesse sedimentary rocks constitute the Permian to Late Triassic Rakaia and the Late Jurassic to Early Cretaceous Pahau terranes and crop out over large parts of New Zealand outside the Haast Schist (Fig. 2). In the North Island, the volcanoclastic Waioeka petrofacies forms a Late Jurassic to Early Cretaceous Waioeka terrane that is not present in the south (Mortimer 1995). The terranes consist of well-bedded but structurally complex felsarenites and mudstones and are inferred commonly as representing accreted subduction complexes (e.g. Howell 1980; MacKinnon 1983; George 1992; Mazengarb & Harris 1994). Oceanic associations of basalt, limestone and chert are preserved as regionally minor components and usually occur in fault or melange zones (Bradshaw 1973; Silberling et al. 1988; Mortimer 1995). Fossils are rare in the Torlesse rocks, but several fossil zones have been recognized (Campbell & Warren 1965; Speden 1974; 1976). Fossil content of the Rakaia terrane indicates a depositional age range of Permian to Triassic, with the youngest Triassic fossils being Rhaetian radiolarians (Campbell 2000b). However, Kamp (2001) argued, based on zircon fission track data and 40Ar/39Ar mica ages, that the Rakaia terrane sandstone sedimentation in the South Island continued into the Early Jurassic (Pliensbachian). The Esk Head Melange (Bradshaw 1973) separates the Rakaia terrane from the Pahau terrane. It has unclear boundaries but may be thought of as a zone of intense deformation of weak sediments between the two terranes. The sandstone and mudstone components of the melange contain Jurassic fossils (Campbell & Warren 1965) and compositional studies show that the Esk Head matrix is essentially of Pahau origin. Allochthonous blocks include Triassic pillow lava, Triassic limestone and Jurassic

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siliceous hemipelagite that are thought to represent the ocean floor and/or seamounts in the original substrate of the Pahau terrane (Silberling et al. 1988). The youngest Rakaia sediments and the oldest Pahau sediments differ in age by c. 45 Ma (Pliensbachian to Kimmeridgian; Fig. 6). The Late Jurassic to Early Cretaceous Pahau terrane has traditionally been interpreted as an accretionary complex made up predominantly of submarine fan turbidites deposited on ocean crust (cf. MacKinnon 1983). A detailed sedimentological study of the Pahau type section, however, gives support to a marginal-marine depositional environment based on the presence of numerous beds with rootlets in both the conglomerate facies and in the mudstone facies in the Pahau River (Bassett & Orlowski 2004). Detrital zircon ages (121 ± 3 Ma) of Pahau conglomerate matrix (Ethelton) indicate a Barremian maximum stratigraphic age for the Pahau sandstone at Ethelton (Wandres et al. 2005). A provenance link between the Rakaia and the Pahau terranes was first proposed by MacKinnon (1983) and also, based on heavy minerals, by Smale (1997). Adams & Graham (1996) concluded that initial 87Sr/86Sr ratios of the younger Pahau terrane (at the time of metamorphism) are inconsistent with the Rakaia terrane being its dominant sediment source. However, Wandres et al. (2005) demonstrate that sandstone clasts in Pahau conglomerates are recycled Permian to early Late Triassic Rakaia rocks, a conclusion supported by detrital zircon age data (Pickard et al. 2000; Cawood et al. 2002) from both terranes. On the basis of petrographic evidence and chemical composition, Mortimer (1995) subdivided the sandstones of the northern quarter of the Torlesse terrane into four new petrofacies: (1) Rakaia; (2) Pahau; (3) Waioeka; and (4) Omaio petrofacies. A comparison of these petrofacies with existing South Island Torlesse classifications indicates continuation of the Triassic Rakaia subterrane and the Late Jurassic to Early Cretaceous Pahau subterrane into the central part of the North Island. A U-Pb age of 99 ± 2 Ma of a detrital zircon from the Omaio petrofacies approximates the age of deposition of these rocks (Cawood et al. 1999). Within the Rakaia terrane, two small tectonic enclaves - the Akatarawa (Hada & Landis 1995; Cawood et al. 2002) and the Kakahu (Hitching 1979; Bishop et al. 1985) terranes (Fig. 2) - comprise more varied Carboniferous Permian rocks, including

volcanic rocks, cherts and fossiliferous limestones. The two tectonic enclaves record the same K-Ar age patterns as the surrounding Rakaia terrane (Adams 1975). Also within the Rakaia terrane, the Late Jurassic Clent Hills Formation consists of shallow-water sediments of very low metamorphic grade (zeolite facies or lower) and commonly has a tectonic contact with older Torlesse rocks (Oliver et al. 1982). This post-Jurassic tectonism is ascribed to a younger, Early Cretaceous phase (Phase II) of the Rangitata Orogeny (Bradshaw et al. 1981) and pre-dates the mid-Cretaceous Mount Somers Volcanics (99-95 Ma, Tappenden et al. 2002). Waipapa terrane. The Waipapa terrane is a diverse assemblage of largely terrigenous turbidites lying west of the Torlesse terranes in the central and northern North Island (Sporli 1978). Nomenclature, internal subdivision and external correlation of the Waipapa terrane are controversial. Black (1996&) considered it to be a superterrane divisible into the OmahutaPuketi (late Palaeozoic to early Mesozoic), Western Bay of Islands (Permian to Early Jurassic) and Manaia Hill subterranes (Late Jurassic to Early Cretaceous). The Manaia Hill subterrane consists of two contrasting facies, one, the younger Morrinsville facies, is compositionally similar to the coeval Waioeka terrane (Torlesse) in the eastern North Island (Mortimer 1995; Kear & Mortimer 2003) and the older Hunua facies includes a melange zone containing slices of Jurassic ocean floor. Campbell (20000) restricted the Waipapa terrane to the Bay of Islands subterrane and assigned the Omahuta rocks to the Caples terrane, the latter supported by Rb-Sr wholerock isochron ages and initial 87Sr/86Sr ratios (Adams 2004). Apart from the OmahutaPuketi block, the other Waipapa rocks have a consistent age and an isotopic signature that differs from that of the Permian to Cretaceous Torlesse metasediments, lending support to the retention of a Waipapa terrane (Adams 2004). Kear & Mortimer (2003) divided the Waipapa terrane into the Bay of Island terrane (Hunua facies) and the Waipa Supergroup, a Late Jurassic to Early Cretaceous, Tutoko Complexderived, volcanoclastic blanket deposited across six older Eastern Province terranes (Mortimer 2004). The boundary between the Waipapa and Murihiku terranes in the North Island coincides with a magnetic anomaly (Junction Magnetic Anomaly). Rare, small serpentine outcrops indicate that the Dun Mountain Ophiolite Belt

NEW ZEALAND TECTONOSTRATIGRAPHY

may be present, although on a very reduced scale (Black 19960). Haast Schist metamorphic overprint. The Haast Schist is a polyphase pumpellyite-actinolite to amphibolite facies metamorphic belt, containing the Otago, Marlborough and Alpine Schists, which has overprinted the Caples, Torlesse and Waipapa terranes (Mortimer 2004, and references therein). In Otago, Rb-Sr whole-rock data (Graham & Korsch 1989; Graham & Mortimer 1992) support K-Ar dating (Adams et al. 1985) indicating a Late Triassic-Early Jurassic age for schist metamorphism, with younger ages (to 115 Ma) representing either long-continued exhumation or a second stage of metamorphism (Adams & Graham 1997). Interpretation of Ar/Ar dates on white micas indicates that peak metamorphism actually occurred in the Middle Jurassic (180-170 Ma) and a second peak (135 ± 5 Ma) represents the onset of rapid exhumation and uplift (Little et al. 1999). Adams & Graham (1997) found that the Otago Schist north of Dunedin has 87Sr/86Sr initial ratios ranging from 0.7064 to 0.7092, similar to the Rakaia terrane. The Chrystalls Beach-Brighton metasediments have 87Sr/86Sr initial ratios that range from 0.7052 to 0.7064. These are lower than the Rakaia rocks yet higher than the Caples-type rocks elsewhere in Otago (0.7035 to 0.7055) (Graham & Mortimer 1992). Initial metamorphism and deformation of the Otago Schist has been attributed to the amalgamation of the Caples and the Rakaia terranes and, over large areas, the Caples terrane is structurally above Rakaia rocks. The present broad arc of schistose rocks has been explained more controversially as a core complex (Forster & Lister 2003), with the core of the Otago Schist exhumed from beneath lowangle ductile shear zones that formed from 112 Ma to 109 Ma. A more detailed discussion of the Caples terrane and the Otago Schist can be found in Roser et al. (1993) and Mortimer (1993).

End ofsubduction and accretion With the exception of the northeastern margin of the Mesozoic New Zealand crust, the midCretaceous and older basement rocks are separated from Late Cretaceous and Cenozoic rocks by a regional unconformity. In many sections, the rocks resting on basement are non-marine and clearly post-date substantial subaerial erosion. The unconformity separates rocks generated by convergent plate margin activity

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from those that relate to continental break-up. The new sedimentary basins, including offshore basins, overlap the terrane boundaries and do not show significantly younger displacements other than those related to the present plate boundary. The situation is less clear in sections close to the northeastern Mesozoic subduction margin (present Raukumara Peninsula (RF), Fig. 2). A break can be identified that separates rocks showing pervasive deformation, attributed to trench slope processes, from similar rocks that lack this deformation. The difference in age is typically small (within one stage) and differs from section to section (Laird & Bradshaw 2004). During the Early Cretaceous, the Phoenix plate was being subducted along the New Zealand sector of the active Gondwana margin and subduction ceased when the Phoenix-Pacific ridge crest and, possibly, the Hikurangi Plateau approached the margin (Bradshaw 1989; Luyendyk 1995). Analysis of a wide range of data indicates that subduction ceased at 105 ± 5 Ma (Bradshaw 1989). On the other hand, the youngest zircon ages in Torlesse-type rocks (Cawood et al. 1999) in the northeast indicate that subduction continued locally until 100 Ma (Kamp 2000). Late subduction can be attributed best to small transformbound relics of the Phoenix plate that were still actively subducted. Inboard of the trench, however, crustal extension started before the end of subduction, forming the Western Province and Otago Schist core complexes (Spell et al. 2000; Forster & Lister 2003). Tuffs of Albian age occur in newly formed basins near the margins of both the Paparoa (Muir et al. 1997) and Haast Schist metamorphic cores (Adams & Raine 1988). Within-plate plutonic and volcanic rocks cut the Pahau terrane at c. 100 Ma (Weaver & Pankhurst 1991). The offshore extensional Great South Basin contains a very thick, nonmarine Cretaceous succession (Cook et al. 1999) and Cenomanian pollen palynomorphs (Raine et al. 1993) occur near the bottom of a drill hole (Tora 1). Seismic interpretation, however, indicates that a further kilometre of sediments lie between the base of the hole and the basement and it is likely that basin development started in the Albian. This period of radical tectonic change is seen also in the originally contiguous part of West Antarctica (Weaver et al. 1994; Storey et al. 1999) and is part of a regional pattern of subduction cessation that moved eastward to the Antarctic Peninsula in the Cenozoic (McCarron & Larter 1998).

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Conglomerate clast geochronology and chemistry The Torlesse terranes are the most outboard and their provenance is controversial (Coombs et al. 1976; MacKinnon 1983; Korsch & Wellman 1988; Bradshaw 1989; Mortimer 1995; Cawood etal. 2002). Sandstone petrography, sedimentary geochemistry, detrital mineral geochronology and isotope geochemical studies of Torlesse sediments have all been employed to address this problem and establish the nature of the Torlesse sources (e.g. MacKinnon 1983; Roser 1986; Roser & Korsch 1986; 1988; 1999; Frost & Coombs 1989; Ireland 1992; Mortimer 1995; Adams & Graham 1996; Smale 1997; Cawood et al. 1999; Kamp 2000; 2001; Pickard et al. 2000; Grapes et al. 2001; Leverenz & Ballance 2001; Adams 2004). Isotopic evidence shows distinct differences between the Caples and the Rakaia terranes, with initial 87Sr/86Sr ratios at the time of metamorphism higher (Adams & Graham 1996) and ENd lower (Frost & Coombs 1989) in the Rakaia terrane. Sr isotopes also show that initial 87Sr/86Sr ratios for the Rakaia terrane are consistent with a postulated continental arc/cratonic source (Adams & Graham 1996). Detrital zircon age distribution in the Rakaia terrane rocks (Ireland 1992; Pickard et al. 2000; Cawood et al. 2002; Wandres et al. 2005) shows a distinct peak at c. 260-240 Ma, suggestive of an active magmatic arc source of Permian age, and various sediment sources have been proposed for the Rakaia terrane, including a provenance east of present-day New Zealand (Andrews et al. 1976), Marie Byrd Land in Antarctica (Korsch & Wellman 1988; Cawood et al. 2002; Wandres et al. 2005) and the New England Fold Belt in eastern Australia (Adams & Kelley 1998; Pickard et al. 2000; Adams 2004). More local sources have also been proposed for the younger Pahau terrane, such as the Rakaia terrane (MacKinnon 1983; Smale 1997; Roser & Korsch 1999; Wandres et al. 2005) or the Tutoko Complex (Mortimer 1995; Wandres etal 2004ft). Barnes & Korsch (1991) demonstrated that there is a wide range of lithologies within the Torlesse source area and that the clasts exhibit trace element characteristics indicative of subduction-related magmatism. The importance and advantages of using igneous conglomerate clasts to determine provenance are discussed by Wandres et al. (20040, ft). Conglomerate clasts, which probably were transported relatively short distances (Kodama 19940, ft; Ferguson et al. 1996), may be used to trace proximal sources. In addition, the pebble- to boulder-sized clasts provide a hand specimen of their source. Fingerprints such as petrography, geochemistry,

isotopes (e.g. Sr-Nd) and crystallization age can be obtained. A detailed rock sampling programme and geochronological, geochemical and Sr-Nd isotope analyses of igneous clasts from Torlesse conglomerates have been presented by Wandres et al. (20040, ft) and salient points are summarized here.

Rakaia terrane igneous clasts The four Rakaia conglomerates at Boundary Creek, Te Moana, McKenzie Pass and Lake Hill (Fig. 2) are dominated by beds of poorly to moderately sorted, matrix- to clast-supported conglomerate interbedded with massive to thinly bedded, medium- to coarse-grained sandstones in graded turbidites (Wandres et al. 2005). All conglomerates are dominated by rounded to well-rounded pebbles and cobbles, but boulders and large angular blocks occur. Clast populations in conglomerates are dominated by sandstones (McKenzie Pass and Lake Hill) and volcanics (Boundary Creek and Te Moana). All conglomerates were deposited in a middle to upper fan environment. SHRIMP U-Pb zircon ages of igneous clasts from four conglomerates define three distinct periods of magmatic crystallization (Fig. 6). The first period ranges in age from 292 to 243 Ma (Permian to Middle Triassic) with two clusters recognizable: a minor Early Permian one ranging in age from 292 Ma to 277 Ma, and a major Late Permian to Middle Triassic one from 258 Ma to 243 Ma. The subduction-related calcalkaline to high-K calc-alkaline, metaluminous to peraluminous clasts range in lithology from andesite to rhyolite and their plutonic equivalents. The second period comprises Carboniferous, calc-alkaline, metaluminous to weakly peraluminous mainly granitoid and rhyolite clasts, ranging in age from 356 Ma to 325 Ma. The third group consists of two Cambrian clasts, a monzogranite (c. 497 ± 8 Ma) and a dacite (c. 517 Ma). Wandres et al (20040) concluded that the geochronology, geochemistry and Sr-Nd isotopes of Rakaia igneous clasts correlate broadly with those of Permian to Triassic plutons and volcanics from the Amundsen and Ross provinces of Marie Byrd Land (Bradshaw et al. 1997; Pankhurst et al 19980). This contrasts with a proposed New England Fold Belt source (Australia) for the Rakaia sediments (Adams & Kelley 1998; Pickard et al 2000; Adams & Maas 2004).

Pahau terrane igneous clasts The three Pahau conglomerates investigated by Wandres et al. (2004ft, Mount Saul, Ethelton and

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Fig. 7. 8Ncj(i37) versus 87Sr/86Sr(137) for igneous clasts from the three conglomerate locations from the Pahau terrane, for the Darran Suite (recalculated from Muir et al. 1998), and Greenland Group (recalculated from analyses presented by Waight et al. 1998). Also shown is a calculated simple mixing curve between the Darran Suite component and a Greenland Group metasediment. Darran Suite calculated as the average of the analyses presented by Muir et al (1998) (Sr(137) = 0.70399, eNd(137) = +3.6, with Sr - 400 ppm and Nd = 20 ppm). Greenland Group is taken as an average from analyses presented by Waight et aL (1998) (Sr(137) = 0.74295, 8 Nd(i3?) - -11-4, with Sr = 94 ppm and Nd = 36 ppm). Tick marks represent 10% increments of mixing Greenland Group with the Darran Suite component. Present day mantle array and depleted mantle from Dickin (1995).

Kekerengu, Fig. 2) are dominated by beds of poorly to moderately sorted, matrix- to clastsupported conglomerate interbedded with massive to thinly bedded, often graded, sandstones. Pebbles and boulders in all conglomerates range in size from <1 cm to 50 cm and are usually rounded to well rounded, but large angular blocks of sandstones (rip-ups) also occur. Clast populations in all conglomerates are dominated by sandstones and volcanics. The conglomerates were deposited in a channelized mid-fan (Kekerengu and Ethelton) and a trenchslope basin on top of an accretionary complex (Mount Saul; Bassett & Orlowski 2004). SHRIMP U-Pb zircon ages of igneous clasts from the three conglomerates range from 147 Ma to 135 Ma and from 128 Ma to 123 Ma (Fig. 6). Calc-alkaline and alkaline clasts are indistinguishable in age, chemical composition and petrogenesis from the granitoids of the Tutoko Complex. An Early Jurassic calcalkaline rhyolite clast from Kekerengu (188 ± 3 Ma) correlates with the Bounty Islands granite, Campbell Plateau (Fig. 1). Geochronology, geochemistry and Sr-Nd isotopes of

igneous clasts identify the Tutoko Complex (s.l) as a contributor of detritus to the Pahau depositional basin. Igneous clast Sr and Nd isotopes. Pahau terrane igneous clast Sr-Nd data are presented in detail elsewhere (Wandres et al. 2004£>) but are used here to demonstrate the petrogenesis of these Pahau clasts, which helps, when combined with the tectonostratigraphic terrane concept of regional geology, to put constraints on the emplacement of the clast source. It is now widely accepted that many granitoids consist of at least two isotopically and geochemically distinct components: a depleted mantle component and a relatively radiogenic continental crustal component (e.g. McCulloch & Chappell 1982; Pickett & Wasserburg 1989). The relatively radiogenic isotopic ratios of most of the Pahau terrane igneous clasts are typical of I-type granitoids (McCulloch & Chappell 1982). The igneous clasts plot within the present-day mantle array of DePaolo & Wasserburg (1979), suggesting the involvement of a significant mantle-derived component (Fig. 7).

Fig. 8. Early Cretaceous reconstruction of Gondwana indicating the main features discussed in the text. Amundsen Province (A) and Ross Province (R) of Pankhurst et al. (19980) are juxtaposed next to each other. South Pole and crustal block configuration adapted from Mukasa & Dalziel (2000), but the displacement between the Amundsen and Ross provinces, proposed by DiVenere et al. (1995), has been removed.

NEW ZEALAND TECTONOSTRATIGRAPHY Plutons similar in age and composition to the Darran Suite plutons (including the Darran Complex, Fig. 5) have been identified as the major source for the Pahau terrane igneous clasts, with minor contributions from an Electric Granite type rock (which has a eNd(i) value very similar to the Darran Suite, but an exceptionally high 87Sr/86Sr initial ratio). Figure 7 shows that the isotopic compositions of clasts from Mount Saul and a single clast from Kekerengu are indistinguishable from those of the Darran Suite plutons and the Electric Granite. However, the isotopic compositions of the remaining conglomerate clasts do not allow for derivation from a single mantle source and indicate the involvement of a crustal component during petrogenesis. Waight (1995) demonstrated that mixing between a depleted mantle component and the Buller terrane Greenland Group (crustal endmember) cannot account for the isotopic compositions of the Hohonu batholith, as the amount of Greenland Group sediments (c. 50%) is too high to retain the I-type characteristics of the Hohonu batholith. Waight et al. (1998) instead used the Separation Point-type depleted mantle of Muir et al. (1995) as the lithospheric mantle source. Here, the Darran Suite has been chosen as the mantle source component, which is very similar to that of the Separation Point-type depleted mantle of Muir et al. (1995). There is no direct evidence of a continental crust within the Tutoko Complex that could have been involved during the petrogenesis of magmas similar in composition to the more radiogenic clasts of Ethelton and Kekerengu. The only sedimentary rocks identified within the Tutoko Complex are the five volcano-sedimentary units discussed above (Figs 3, 5) that rest unconformably on plutonic rocks and are Jurassic or earliest Cretaceous in age. However the emplacement of the Crow, Copperstain and Deceit granites into the Western Province justifies the assumption that the Greenland Group of the Buller terrane not only forms the country rock to the plutons but also was assimilated during petrogenesis. Modelling indicates that the isotopic compositions of the Ethelton and Kekerengu clasts are achievable by the mixing of Darran Suite derived melt and 10-25% average Greenland Group composition (Fig. 7). The mixing between the two components is, of course, a simplification, as the lower continental crust of the Buller terrane is more likely to be an extremely complex and heterogeneous mixture of unknown lower crustal and Palaeozoic-Mesozoic igneous components (e.g. Muir

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et al. 19965). The crustal contribution required during the petrogenesis of the more radiogenic igneous clasts, together with the contemporaneous emplacements of the Darran Suite plutons and the Crow, Copperstain and Deceit granites lends further support to an autochthonous model for the Tutoko Complex, as discussed above.

Reconstructing New Zealand Data drawn from the New Zealand tectonostratigraphy summarized above, combined with data from the described conglomerate clasts, play a pivotal part in the reconstruction of the New Zealand sector of the Panthalassan Gondwana margin. This account of the Permian to Cretaceous New Zealand terrane configuration is drawn from the data presented in the time-space diagram (Fig. 6) and a set of palaeotectonic geographical stage maps (Figs 7, 8). Conglomerate geochronology and chemistry indicate that convergent tectonics dominated from at least the Carboniferous to the Early Cretaceous along an active continental margin with the deposition of the New Zealand Eastern Province sediments. The location of the New Zealand microcontinent immediately prior to the separation from Marie Byrd Land (West Antarctica) is well known (Mayes et al. 1990; Lawver & Gahagan 1994; Sutherland 1999; Mukasa & Dalziel 2000). Results summarized here indicate that the South Island Pahau terrane sedimentation continued from the Late Jurassic up to almost the date of fragmentation of this Gondwana sector (c. 87 Ma, Laird & Bradshaw 2004). Given that the approximate location of the New Zealand block at that time is well constrained, it is reasonable to reconstruct the Late Jurassic to Early Cretaceous time first, followed by a Permo-Triassic compilation.

Late Jurassic to Early Cretaceous reconstruction Tectonic constraints implied by the nature and ages of igneous clasts. Calc-alkaline Pahau conglomerate clasts analysed by Wandres et al. (2004b) are indistinguishable in age (Fig. 6), chemical composition and petrogenesis from the calc-alkaline granitoids of the Darran Suite and the Thurston Island granitoids from West Antarctica. The A-type clasts from Mount Saul and Kekerengu are indistinguishable in age, geochemistry and petrogenesis from the Electric Granite, whereas clasts with adakitic

Fig. 9. Late Triassic reconstruction of the Gondwana margin indicating the main features discussed in the text. South Pole, 60°. Meridian and crustal block configuration are from Gahagan et al (1999). LHR, Lord Howe Rise; CHP, Challenger Plateau; CR, Chatham Rise; EMBL, eastern Marie Byrd Land; WMBL, western Marie Byrd Land; AP, Antarctic Peninsula; EWM, Ellsworth-Whitmore Mountains; SA, South America.

NEW ZEALAND TECTONOSTRATIGRAPHY

affinities can be correlated best geochemically with the Separation Point Suite. Wandres et al. (20046) identified the subduction-related Tutoko Complex/Amundsen Province plutons as the major source for the Pahau terrane igneous clasts (Fig. 7). Rounding of the conglomerate clasts indicates relief in the source area and the clast lithologies preserved in the conglomerates indicate that a spectrum of intrusive levels was exposed at the time of the conglomerate deposition. The paucity of volcanic rocks in present-day source areas and the ubiquitous presence of predominantly rhyolitic volcanic and hypersolvus granitoid clasts in the conglomerates are distinctive. They point strongly to erosion of the higher levels of the Tutoko Complex and the Amundsen Province and subsequent transportation to the place of final deposition. The presence of both Darran Suite and Separation Point-type derived clasts at Kekerengu and of the Electric Granite derived clasts at Mount Saul is distinctive. Furthermore, the similar petrogenesis of A-type clasts from Kekerengu and Mount Saul and of I-type clasts from all three conglomerate locations is noteworthy. If the close proximity (present geography) of the three conglomerate locations is considered, then this might indicate that all three conglomerate locations were sourced from a relatively narrow sector of the Tutoko Complex where the Darran Suite, the Separation Point-type rocks and the Electric Granite were in close proximity. Granitoid clasts from the Chatham Island conglomerates show strong similarities with their counterparts in the conglomerates discussed here (Dean 1993 in Wandres et al. 20046). In particular, some Atype granitoid clasts from the Chatham Island conglomerates (Fig. 1) strongly resemble the Atype volcanic clasts from Mount Saul and Kekerengu. Wandres et al. (20046) correlated the Early Jurassic calc-alkaline I-type clast from Kekerengu with the granite exposed on Bounty Islands, which share a similar petrogenetic history with the contemporaneous Pine Island granitoids, suggesting that Campbell Plateau was in close proximity to Thurston Island/Marie Byrd Land in the Early to Middle Jurassic. The calc-alkaline clasts (147-123 Ma) are similar geochemically to the Whitsunday Volcanics (eastern Australia, Fig. 1, Ewart et al. 1992), and Sr- and Nd-isotopes of clasts from Mount Saul and Kekerengu are similar to units of this volcanic province. Volcanic activity in the extension-related Whitsunday Volcanic Province (WVP, 125-95 Ma, Ewart et al. 1992;

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Bryan et al. 2000) was penecontemporaneous with the sedimentation in the Pahau basins and extension-related magmatism recorded in Marie Byrd Land (105-95 Ma, Storey et al. 1999). The occurrence of volcaniclastic sediments derived from the WVP in the Otway-Gippsland Basin (Bryan et al. 1997) indicates that the WVP arc extended into what is now known as the Lord Howe Rise and may be the source for the tuffs in the Kyeburn and Horse Range formations (Fig. 2) and the Stitts Tuff in the Pororari Group (Fig. 3). The possible extension of the WVP arc into the New Zealand region (Challenger Plateau) is shown in Figure 8. Extensional period? The Tutoko Complex experienced extension and/or temporarily waning magmatism in the Tithonian, indicated by the dates obtained from alkaline clasts. Geochemically these clasts show no affinities with magmas usually associated with hotspots or mantle plumes (Wandres et al. 20046). The clasts show characteristics similar to magmas interpreted as being derived from continental crust, or crust that has been through a cycle of continent collision or island-arc magmatism and emplaced in a variety of tectonic settings, including post-collisional, or what may be true anorogenic magmatism. If the alkaline clasts are related to the waning stage of a subduction system, then this might indicate that the Pahau source subduction system experienced a temporary extensional event as suggested by Tulloch & Kimbrough (1995). Crustal extension in New Zealand started before the demise of the subduction system and is evident from core complexes that formed in the Eastern and Western provinces (between c. 140-110 Ma, Gibson et al. 1988; Spell et al. 2000; Forster & Lister 2003) and from an upper Jurassic-Cretaceous rift fill sequence in the Taranaki Basin (Uruski 2003). Regionally, the younger New Zealand extensional events (125-100 Ma) coincide with the development of a core complex in Marie Byrd Land (Luyendyk et al. 1996). The older extension coincides broadly with an extensional event in the West Antarctic rift system (De Santis et al. 1994) and the break-up between Australia and East Antarctica (Muller et al. 2000). Rift-related silicic magmatism occurred in West Antarctica/South America (Chon Aike province, Sarmiento complex) and was emplaced between c. 190 Ma and 140 Ma (Pankhurst et al. 1998b; 2000; Riley et al. 2001; Stern & De Wit 2003), coinciding with the emplacement of the extensive mafic Ferrar magmatism that extended into

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the New Zealand region (Kirwans Dolerite, Mortimer et al. 1995). Silicic magmatism in the WVP (Australia) occurred from c. 125 Ma to 95 Ma (Ewart et al. 1992) and was associated with rifting. All these Jurassic/Cretaceous events appear to be related, progressing from the Antarctic sector at c. 190-140 Ma and 105-95 Ma through the New Zealand sector at c. 140-100 Ma to the eastern Australian sector at c. 135-95 Ma. It is possible that the extensional event associated with the Middle Jurassic Kirwans Dolerite was the cause of the exhumation of part of the Western Province (Longwood Range cooling, Mortimer et al. 19990), which, at that time, experienced a magmatic lull. The deposition of the Rainy River Conglomerate (Tutoko Complex) and the Barretts Formation of the Brook Street terrane (Tulloch et al. 1999; Adams et al. 2002) again may indicate a period of exhumation coinciding their deposition. At the same time as the Middle Jurassic extension took place, part of the Eastern Province experienced strong regional compression, ascribed to the Caples-Torlesse terrane collision (e.g. Mortimer 1993, and references therein) and the formation of the Otago Schist. Recycling of the Rakaia terrane. Several lines of evidence presented by Wandres et al. (2005) and others strongly support the earlier claims (MacKinnon 1983; Roser & Korsch 1999) that Pahau depocentres received detritus recycled from older inboard, mainly Rakaia, terranes. Rapid cooling of the Otago Schist between 140 Ma and 130 Ma has been interpreted by Little et al. (1999) as a signal of exhumation culminating in the Otago metamorphic core complex (Forster & Lister 2003), thereby making Rakaia/Caples terranes available for erosion. Sandstone clasts collected from Pahau conglomerates (Ethelton and Mount Saul) show strong petrographic similarities with Rakaia sandstones and this affinity is confirmed by the geochemistry of these clasts (Wandres et al. 2005). The recycling of Caples/Rakaia sandstone was proposed by Roser & Korsch (1999) to explain the Pahau sandstone composition. Results from work here show that the Tutoko Complex also contributed detritus to the Pahau Basin and that at least three or more sources are required to produce the Pahau sandstones. There are various other source rocks that could have contributed detritus by erosion to the Pahau Basin, including the Brook Street terrane and the Longwood and Holly Burn Intrusives (Mortimer et al. 19990). Detritus contribution by the Maitai and Murihiku terranes is a possi-

bility. That the Murihiku Basin received detritus not only from the Maitai/Caples terranes to the east, but also from the Tutoko Complex to the west is indicated by the presence of Tutoko Complex-like igneous clasts in the Murihiku terrane (Graham & Korsch 1990). Furthermore, the detritus shed from the Tutoko Complex and the Maitai/Caples terranes is broadly similar in £Nd composition to that of the Murihiku terrane (Frost & Coombs 1989; Wandres et al. 20046). If borehole (Waimamaku-2) interpretations by Isaac et al. (1994) are correct (bottom of core interpreted as Murihiku), then the Murihiku terrane is still a depositional basin (sediment trap) at the time of the Pahau sandstone deposition and cannot be a major contributor of detritus. This problem of getting detritus across the Murihiku Basin is resolved by the suggestion of Kear & Mortimer (2003) that subsidence of the Murihiku Basin ceased in the Late Jurassic and an unconformable overlap succession, the Waipa Group, lay across the Murihiku and adjacent terranes. Zircon U-Pb age patterns also point towards a component of recycled Rakaia-type material in the Pahau conglomerates. Triassic sandstones from the Rakaia, Caples and Waipapa terranes contain major detrital zircon peaks at 250-230 Ma, while additional minor age populations at 300-280 Ma and 340-320 Ma are found in Rakaia sandstones (Pickard et al. 2000; Cawood et al. 2002; Wandres et al. 2005). Sandstones from all three terranes also have minor Ordovician and Precambrian components. All of these zircon age components are found in the sand-sized matrix of the Barremian (121 Ma) Pahau conglomerate at Ethelton (Wandres et al. 2005). The only obvious difference between the patterns is a minor Devonian peak in the conglomerate matrix that appears to be absent from the Rakaia/Caples/Waipapa sandstones. Overall the similarities are striking. Other zircon evidence for recycling includes the presence of Early Jurassic age peaks in the Waipapa/Caples sandstones (Pickard et al. 2000) and in both Pahau sandstones (Omaio facies, Cawood et al. 1999) and the Ethelton conglomerate matrix (Wandres et al. 2005). The study by Wandres et al. (20046) also identified an Early Jurassic rhyolitic clast in the conglomerate at Kekerengu (188 ± 3 Ma). The Early Jurassic zircon age peak coincides broadly with the Jurassic age proposed for the Caples/Rakaia terranes' docking, magmatism in the Bounty Islands region and exhumation in the Longwood Ranges (Little et al. 1999; Mortimer et al. 19990). The Aptian reconstruction of the New

NEW ZEALAND TECTONOSTRATIGRAPHY Zealand sector of the Gondwana margin (Fig. 7) is adapted from that proposed by Mukasa & Dalziel (2000) but modified by removing the dextral displacement proposed by DiVenere et al. (1995) between the eastern and western Marie Byrd Land crustal blocks (Ross and Amundsen Province of Pankhurst et al. 19980). This modification is based on a number of tectonostratigraphic constraints. (1) The LeMay Group of Alexander Island, Antarctic Peninsula, is a Mesozoic accretionary prism constructed during subduction of Pacific ocean floor and the Phoenix plate (Holdsworth & Nell 1992; Tranter 1992; Doubleday et al. 1993; McCarron & Larter 1998). Radiolarian biostratigraphy constrains its depositional age range to include the range from latest Jurassic to Albian (Holdsworth & Nell 1992). If sedimentation of the LeMay Group continued at least into the Albian, then the positioning of the Thurston Island block in front of Alexander Island in the Barremian reconstruction of Mukasa & Dalziel (2000, their fig. 9a) is questionable. In the reconstruction here, the Thurston Island, eastern Marie Byrd Land and eastern Campbell Plateau blocks (Mukasa and Dalziel terms) have been moved sinistrally, thereby removing the strike-slip separation between the Marie Byrd Land and Campbell blocks and closing the 'gap' between the Campbell and Challenger plateaux. (2) The strike-slip separation proposed by DiVenere et al. (1995) is based on data from granitoid dykes and volcanics along the Ruppert Coast, including areas adjacent to the core complex structure in the Fosdick Mountains of the Ford Ranges (Luyendyk et al. 1996). Movement associated with the core complex and a zone of continental separation may have caused post-magnetization rotation of the measured poles. This may account for palaeopoles that require strike-slip displacement and the placement of the Thurston Island crustal block in front of the accretionary complex while accretion was still taking place. (3) One of the most striking features on the present-day Campbell Plateau is a northeast-trending zone of high amplitude positive magnetic anomalies, termed the Campbell Magnetic Anomaly System (CMAS, Davey & Christoffel 1978). The gravity and the magnetic anomalies indicate a major geological feature under-

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lying the area, but its possible correlation with known tectonic zones (Tutoko Complex, Amundsen Province) is uncertain (Sutherland 1999). If the Bounty Islands granite is part of the Tutoko Complex (Cook et al. 1999), then the CMAS may represent Tutoko Complex correlatives intruded into the Campbell Plateau. Bradshaw et al. (1997) proposed similar palaeo-relationships. The reconstruction here, therefore, shows the Tutoko Complex extending across the Campbell Plateau to join with Amundsen Province correlatives. (4) The continuity of the CMAS across the entire central Campbell Plateau removes the need for major strike-slip displacement between the Eastern and Western Campbell plateaux as proposed by Mukasa & Dalziel (2000) and lends further support to the crustal block arrangement presented here. (5) It has been suggested that the Torlesse terranes are 'exotic' (DiVenere et al. 1995; Pickard et al. 2000). In the Great South Basin (Fig. 1) non-marine graben-fill successions overlie basement rocks of the Eastern Province (Beggs 1993). Cenomanian ages for the cover sequence are suggested by pollen data from a drill hole (Raine et al. 1993). However, seismic interpretation indicates that c. 1000 m of older sediments lie between the base of the drill hole and the basement and an Albian age for these units is possible. Therefore, the terranes of the Eastern Province were probably in place at c. 105 Ma. There is no such control on the juxtaposition of the Pahau terrane and the inboard Eastern Province terranes. The Pahau terrane was cut by the within-plate Mandamus Suite by c. 100 Ma (Weaver & Pankhurst 1991) and related lamprophyric dykes cut the Esk Head Melange. Therefore, scope for an exotic Cretaceous segment of the Pahau, as proposed by DiVenere et al. (1995), is limited.

Permian to Late Triassic reconstruction A key to understanding the Permo-Triassic southeast Gondwana margin configuration is the provenance of the Rakaia terrane. The Rakaia terrane constitutes (together with the Pahau terrane) approximately 60-70% of the New Zealand land area and a large part of the microcontinent. As discussed previously, different sources have been proposed, resulting

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in contradicting palaeotectonic reconstructions. In this context, it is therefore important to discuss the provenance of this accretionary complex. The Rakaia terrane sediments are coarse grained and show a marked compositional uniformity from the Permian to the Late Triassic, indicating either rapid deposition by large energetic rivers that drained a relatively uniform continental arc source or that there was an efficient mixing system operating during transport and deposition. Clast data from four conglomerates (Figs 2 & 9) have helped to characterize broadly the age, chemical and isotopic composition and petrogenesis of the sources (Wandres et al 2005). Assuming that the clasts in a conglomerate broadly reflect the lithotypes exposed in the source area, then the petrography of the collected clasts reflects the structural levels of the source that were available for erosion. Clast petrography indicates that detritus provided by a magmatic source to the (Permian?) Boundary Creek conglomerate was mainly by volcanic activity and the erosion of subvolcanic levels. A similar conclusion can be drawn for the Kazanian Te Moana and the Dorashamian McKenzie Pass conglomerates, where volcanic and hypersolvus clasts dominate over subsolvus clasts. At Lake Hill (Carnian) mylonitic and gneissic clasts indicate that deeper levels of the arc were exposed and eroded, and that volcanic lithologies were minor. The exposure of gneissic rocks indicates that substantial erosion of the upper levels of the source took place before Carnian times. Adakitic rocks sampled from this conglomerate probably attest to the presence of a mature crust, given that one petrogenetic model involves the melting of newly underplated mafic magma beneath a thickened continental crust (Muir et al. 1995; Petford & Atherton 1996; Wareham etal. 1997). The apparent trend shown by geochronological data and petrographic observations indicates the progressive unroofing of a continental volcanic/plutonic arc source that experienced near-continuous magmatism from (at least) the Carboniferous to the Middle Triassic. The trend observed in the igneous clast source is paralleled by the progressive evolution of the Rakaia accretionary wedge from the Permian to the Late Triassic (e.g. MacKinnon 1983; Roser & Korsch 1999). Petrographic and geochemical data of sandstone clasts from the Rakaia terrane and Rakaia sandstones indicate that clasts in the Te Moana and Lake Hill conglomerates were derived by autocannibalistic reworking of older, consolidated, Rakaia

sediments (Wandres et al. 2005). This lends support to active and rapid accretionary processes along the Panthalassan Gondwana margin in Permian times, contradicting a passive margin model for the Rakaia sandstones (Coombs et al. 1976, their fig. 9). The unroofing trends also coincide with trends observed in the Murihiku terrane where Boles (1974) reported an increase in SiO2 from the Early to the middle Late Triassic and a subsequent reversion to more volcanic detritus from the middle Late Triassic to the Early Jurassic. Boles (1974) ascribed the compositional shift to the change from predominantly andesitic to predominantly dacitic rhyolitic to once again andesitic volcanism in the source. Detrital zircon ages of Permo-Triassic sandstones from the Western Province (Parapara Peak Group, Fig. 6, Wysoczanski et al. 1997) are very similar to those of the Rakaia sandstones (Pickard et al. 2000; Cawood et al. 2002; Wandres et al. 2005), suggesting a common source and a possible proximity of the two New Zealand provinces in Permo-Triassic times (Wysoczanski et al. 1997). Cambrian magmatism was confined to the Western Province and its Australian and Antarctic correlatives, and the presence of Cambrian cobble-sized detritus clasts supports a linkage between the Rakaia terrane and the Gondwana margin, at least in Kazanian times. Although the provenance of the Rakaia terrane clearly indicates derivation from Gondwana, the lateral continuity of rock units along the southeast margin of the supercontinent makes definition of the exact location of the source difficult. Detrital zircon age distributions from Rakaia sandstones identify a Permo-Triassic source as the main contributor of detritus (Ireland 1992; Pickard et al. 2000). Comparison of Permo-Triassic Rakaia igneous clasts with the potential source provinces is ambiguous and inconclusive and both Australia (New England Fold Belt, NEFB) and Antarctica (Amundsen Province, Marie Byrd Land) were identified as likely sources for these igneous clasts (Wandres et al. 20040). In combining geochronological, geochemical and isotopic data of the Rakaia igneous clasts these authors envisaged an igneous source for the clasts as a volcanic/plutonic continental arc that experienced continuous magmatism into the early Middle Triassic, and had Cambrian and Carboniferous plutons and volcanics exposed at the time of erosion. These two potential sources, also proposed by other researchers (e.g. Korsch & Wellman 1988; Adams & Kelley 1998; Pickard et al. 2000; Cawood et al. 2002; Adams 2004), are evaluated here in the light of

NEW ZEALAND TECTONOSTRATIGRAPHY

tectonostratigraphic and regional constraints, and a corresponding palaeotectonic reconstruction is shown in Figure 9. Australian source. The detrital zircon age populations of Rakaia sandstones (Pickard et al. 2000; Cawood et al. 2002; Wandres et al. 2005) are dominated by a distinct 300-200 Ma age group and are similar to 40Ar/39Ar muscovite mineral ages reported by other researchers (Adams & Kelley 1998). These age data are characteristic and point to a NEFB provenance for the Rakaia sandstones according to Adams & Kelley (1998) and Pickard et al. (2000). These authors proposed a model whereby the depocentres were subsequently tectonically transported, by strike-slip motion >2500 km over 60 Ma, southwards around the Panthalassan margin of Gondwana into a high-latitude Cretaceous position. This model is also favoured by Veevers (2000, p. 121) who noted that the detrital zircon age population (and provenance) of the Torlesse sandstones is very similar to that of the Early Triassic Terrigal Formation of the Sydney Basin and modern beach sands in northern New South Wales (Hummock Hill Island, Sircombe 1999), both of which he considers to be derived from the NEFB. Based on this observation Veevers (2000) supports the model whereby the Torlesse sandstones were deposited adjacent to the NEFB (Pickard et al. 2000). However, the Hummock Hill Island beach sand has a pronounced zircon age peak between 400 Ma to 300 Ma (Sircombe 1999, his fig. 2) and the same age group is the dominant age group in the Terrigal Formation (Sircombe 1999, his fig. 4). This 400-300 Ma peak is almost completely absent from the Triassic Torlesse muscovite data of Adams & Kelley (1998) and Adams et al. (1998). 40Ar/39Ar single crystal mineral ages of detrital muscovites from Rakaia sandstones show an age population dominated by a major Permo-Triassic (280-205 Ma) peak and a minor mid-Palaeozoic (460-410 Ma) peak. There is a complete absence of peaks in the range 500-460 Ma and 410-330 Ma. The Devonian to Carboniferous peak differs markedly from the zircon age distribution observed in the Rakaia sandstones, which have this peak poorly developed (or absent). Therefore, unless both the Triassic Terrigal Formation and modern NSW sands had contributions from mid/late Palaeozoic sources unavailable to the Permo-Triassic Rakaia sandstone on the east side of the NEFB, the zircon evidence presented in favour of the NEFB source is, perhaps, not as convincing as assumed by the authors.

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In the Permian, the Bowen Basin formed behind the arc in a back-arc extensional to foreland-basin setting, between the NEFB and the cratonic block to the west (Roser & Korsch 1999, and references therein). Deposition continued through the earliest Late Triassic time, and the Bowen Basin fill thus represents the back-arc deposits of the proposed Rakaia source. Petrographic data from the Bowen Basin differ significantly from data from the Rakaia sandstones and Roser & Korsch (1999) noted that the compositional contrast between the Rakaia and Bowen sediments requires a sorting mechanism that sheds quartzofeldspathic detritus into the forearc basin and volcanolithic to quartzose detritus into the back-arc basin. Roser & Korsch (1999) argued that the Rakaia and the Bowen sandstones could only have been derived from the same arc if a major contrast in lithotype (volcanic versus granitoid) was present across it. The Permo-Triassic igneous clast population of the Rakaia terrane shows a good correlation with New England Supersuite rocks and a NEFB provenance for the penecontemporaneous Rakaia detritus is, therefore, feasible. However, Famennian to Visean magmatism (Boundary Creek clasts) in eastern Australia was restricted to the back-arc basin of the NEFB (both sides along the Sydney Basin) and to small pockets in north Queensland. Cambrian magmatism was, apart from the Upper Bingara Plagiogranite, restricted to Tasmania Delamerian Orogen, South Australia, Foden et al. (2002), and north Queensland (Scheibner & Veevers 2000, their figs 219 and 215), i.e. the NEFB clearly lacks the older igneous rocks required to explain the presence of these older cobble-sized igneous clasts in the Rakaia conglomerates. In addition, tectonostratigraphic and regional constraints, together with the discrepancy between the detrital mineral age data from the Rakaia sandstones and the NEFB igneous data, do not support a NEFB provenance for the Rakaia sandstones. New Zealand/Antarctica source. In stark contrast to eastern Australia, there are very few rocks exposed in West and East Antarctica (<1% of the total area, Stump 1995), and interpretations offered in this summary are based inevitably on data from a few selected outcrops. The Carboniferous to Triassic granitoids from the Western Province and Amundsen Province represent a remnant volcanic/plutonic arc that was active at the time of the Rakaia sandstone deposition (Weaver et al. 1991; Palais et al. 1993; Pankhurst et al. 1993; 19980). The

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close proximity between the Rakaia terrane and the Western Province of New Zealand (Mortimer et al. 19990) indicates that other Carboniferous igneous rocks (Ross and Western Province correlatives) were in close proximity to the Western Province and Amundsen Province plutons at the time of Permo-Triassic Rakaia sedimentation. In addition, Cambrian orthogneisses, now exposed at Mount Murphy and along the Wallgreen Coast and in the Western Province were in the vicinity of Western Province and Amundsen Province plutons. Other Cambrian rocks in Antarctica occur in the Transantarctic Mountains (TAM) that are now separated from Marie Byrd Land (Ross Province) by the West Antarctic rift system (WARS). Hamilton et al. (2001) identified two major Mesozoic to Cenozoic extensional phases in this rift system that post-date the PermoTriassic Rakaia sedimentation, indicating that Marie Byrd Land Antarctica occupied a position closer or adjacent to the TAM at the time of sedimentation. However, in the PermoTriassic, the TAM were part of a depositional basin and were covered by the Devonian to Triassic Beacon Supergroup (e.g. Collinson et al. 1994, and references therein). Detrital contribution of the TAM to the Rakaia basin must, therefore, be considered to be minor. The dacitic clast from the Lake Hill conglomerate correlates with the Liv Volcanics of the central TAM. The presence of a clast with a TAM-like signature in the Rakaia depocentre may indicate that prior to the opening of the WARS, and at the time of the Rakaia sedimentation, the TAM rocks may have been exposed to erosion along the Panthalassan edge of the central Transantarctic sedimentary basin. The edge of the sedimentary basin is now covered by the Ross Ice Shelf and any contribution of Cambrian TAM detritus to the Rakaia basin remains speculative. The Cambrian monzogranite from Te Moana correlates best with basement rocks of West Antarctica (Ross Province), which is favoured as the Cambrian protosource for the Rakaia detritus. The subdued Cambrian detrital zircon age peaks in both Rakaia sandstones are ascribed to the detrital contribution of the Amundsen/Ross Province basement. Magmatism in the source and sedimentation of the Rakaia sandstones was penecontemporaneous and the sedimentary basins received boulder, cobble, pebble and sand detritus from this source. The presence of boulder- and cobble-sized clasts indicates a short travel distance for the coarse detritus (see above). Age

signatures of sandstone detrital zircons (Pickard et al. 2000; Wandres et al. 2005) match well with igneous clast ages, indicating that the depositional basins were positioned adjacent to the volcanic/plutonic arc segment from which the clasts were derived. The Western Province and Amundsen Province (including the older provinces) comprised plutons and volcanics of appropriate ages to explain the igneous clast population sampled in the Rakaia conglomerates and to account for the detrital zircon age signatures in the Rakaia sandstones (Wandres et al. 20040). The Permian detrital zircon age peaks of the Rakaia sandstones support the conclusion that granitoids from the Kohler Range were a major source for the Triassic sediments, and isotope data presented by Wandres et al. (20046, their fig. lib) indicate that erosion of Amundsen Province plutons could produce at least part of the Rakaia terrane detritus. The Carboniferous age component in Torlesse samples resembles that of the Ross Province magmatism. The exposure of a Cambrian protosource next to or within the Western Province and Amundsen Province is also shown in most Torlesse samples by the presence of a subdued Cambrian peak. Geochronology and geochemistry of the New Zealand and West Antarctica plutons correlate broadly with igneous clasts from the four Rakaia conglomerates. The Amundsen Province and Ross Province are postulated as likely sources for the Permo-Triassic, Carboniferous and Cambrian igneous clasts, respectively. Apart from the Early Permian fusulinid fauna within the Akatarawa terrane (Hada & Landis 1995; Leven & Campbell 1998) the inferred cold-water affinities of the Rakaia faunal assemblages support a deposition of the Rakaia detritus in a high-latitude environment (Ito et al. 2000; Shi & Grunt 2000; Cawood et al. 2002). In addition, palaeomagnetic studies have constrained the apparent polar wander path of the Australian continent and its Gondwana neighbours in the Late Palaeozoic, Mesozoic and Cenozoic (Schmidt & Clark 2000, and references therein). The Late CarboniferousPermian (327-250 Ma) pole path indicates a polar latitude up to the end of the Permian for the Australian part of Gondwana. The Triassic to Early Jurassic group of poles indicates mid to high southern latitudes for Australia before its return to higher latitudes (Schmidt & Clark 2000), with Antarctica and the Western Province of New Zealand generally depicted as being in the vicinity of the South Pole (Powell & Li 1994; Mukasa & Dalziel 2000). This further

NEW ZEALAND TECTONOSTRATIGRAPHY

supports an Antarctic source for the Rakaia terrane, as postulated here. Summary The basement geology of New Zealand comprises Early Cambrian to Early Cretaceous rocks that formed along Gondwana's Panthalassan margin. The geological record preserved in these rocks reflects processes of continental growth by tectonic and magmatic addition along a convergent plate margin. The pre-Early Cretaceous basement rocks are described in terms of batholiths and tectonostratigraphic terranes, which have been grouped traditionally into three provinces: the Western Province, the Median Tectonic Zone and the Eastern Province. The Western Province consists of the Buller and Takaka terranes, two distinct Lower Palaeozoic metasedimentary and arc volcanic units that amalgamated in the Middle Devonian and are cut by Devonian, Carboniferous and Early Cretaceous granitoids. The Eastern Province consists of arc, forearc and accretionary complex rocks that relate to Permian to Cretaceous plate convergence. Geochronological research on igneous rocks over the last decade, backed by structural and petrogenetic work, has provided a clearer picture of the plutonic record and indicates that the Median Tectonic Zone is best regarded as a magmatic arc complex that formed within and adjacent to the edge of the Western Province. In this paper it is proposed that the rocks assigned formerly to the Median Tectonic Zone be regarded as part of the Western Province. The Early Jurassic to Early Cretaceous plutonic and volcanic rocks of the former Median Tectonic Zone and their high pressure metamorphic equivalents of Fiordland represent a single magmatic assemblage disguised by thrusting and obduction during Cretaceous convergence. These rocks, which are related to the last major magmatic pulse and crustal thickening, are grouped as the Tutoko Complex. Crustal thickening was succeeded immediately by rapid exhumation manifested in the development of the metamorphic core complexes that was accompanied by minor magmatism. Regional subduction ceased at 105 ± 5 Ma when the Phoenix-Pacific ridge crest and, possibly, the Hikurangi Plateau approached the margin, but continued locally until 100 Ma. The continental extension culminated in the opening of the Tasman Sea and the Southern Ocean followed by further disruption and re-configuration of the New Zealand microcontinent along the presently active plate boundary.

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The timing of accretion of all the Eastern Province terranes and, in particular, the Torlesse terranes to the Gondwana margin is an important requirement for the understanding of the Palaeozoic-Mesozoic evolution of the southwest Pacific Gondwana margin. The Western Province, including the Mesozoic magmatic arc along its eastern margin, can be identified offshore on the continental shelf and extends into Antarctica. Terranes of the Eastern Province have analogues in Caledonia and Antarctica but do not have direct correlatives. Over the last two decades the field of provenance analysis has undergone a revolution with the development of single-crystal isotope dating techniques. Single zircon grains have the potential to travel great distances and so cobble-size conglomerate material is a more certain index of provenance. This paper has summarized results from a detailed rock sampling programme of conglomeratic rocks from New Zealand terranes and combined these data with the most recent research on New Zealand tectonostratigraphy to constrain the tectonic evolution of the late Palaeozoic to Mesozoic New Zealand sector of the Gondwana margin. The conglomerate data have helped to broadly characterize the igneous source for the Pahau and Rakaia terranes. The conglomerate locations were chosen to represent the full stratigraphic range of both terranes, and the geographical distribution of the conglomerates approximates an inboard to outboard transect of the two terranes with respect to the Panthalassan margin of Gondwana. Geochronology and geochemistry of igneous clasts from Pahau terrane conglomerates identify the Tutoko Complex/Amundsen Province plutons as major contributors of detritus to the Pahau depositional basin. The paucity of volcanic rocks in present-day source areas and the ubiquitous presence of predominantly rhyolitic volcanic clasts in the conglomerates point strongly to the stripping and erosion of the higher levels of the source and subsequent transportation to the place of final deposition. Several lines of evidence presented here strongly support earlier claims that Pahau depocentres received detritus recycled from older Torlesse rocks of the inboard Rakaia terrane. The igneous clast data provide the best evidence yet that the Pahau terrane is derived locally. Geochronological and petrological data of Rakaia igneous clasts point towards a progressive unroofing of a continental volcanic/plutonic arc source that experienced continuous magmatism from the Carboniferous to the latest

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sequence, Dansey Pass, South Island, New Middle Triassic. Source magmatism and depoZealand. Journal of the Geological Society, sition of the Rakaia sediments were peneconLondon, 142, 339-349. temporaneous, as indicated by the presence of a late Early Permian volcanic igneous clast in ADAMS, C.J., BARLEY, M.E., FLETCHER, I.R. & PICKARD, A.L. 1998. Evidence from U-Pb zircon the early Late Permian Te Moana conglomerate and 40Ar/39Ar muscovite detrital mineral ages in and further supported by detrital zircon ages of metasandstones for movement of the Torlesse the sandstones. Detrital zircon ages of Rakaia suspect terrane around the eastern margin of sandstones and igneous clast ages correlate Gondwanaland. Terra Nova, 10, 183-189. broadly with crystallization ages of plutons and ADAMS, C.J., BARLEY, M.E., MAAS, R. & DOYLE, M.G. 2002. Provenance of Permian-Triassic volcanivolcanics from the Amundsen Province and clastic sedimentary terranes in New Zealand: Ross Province, and the Antarctic sector of the evidence from their radiogenic isotope characPanthalassan Gondwana margin has to be teristics and detrital mineral age patterns. New (re)considered as the likely source for the Zealand Journal of Geology and Geophysics, 45, Permian to Triassic, Carboniferous, and 221-242. Cambrian igneous clasts and the Rakaia sedi- AITCHISON, J.C, LANDIS, C.A. & TURNBULL, I.M. 1988. ments. Stratigraphy of Stephens Supergroup (Maitai The authors would like to thank R. Maas, N. Mortimer, T. Ireland and G. Gibson who have helped to improve earlier versions of the manuscript. This work was funded by a scholarship for A.W. at the University of Canterbury.

References ADAMS, C.J. 1975. Discovery of Precambrian rocks in New Zealand. Age relations of Greenland Group and the Constant Gneiss, West Coast, South Island. Earth and Planetary Science Letters, 28, 98-104. ADAMS, C.J. 2004. Rb-Sr age and strontium isotopic characteristics of the Greenland Group, Buller Terrane, New Zealand, and correlations at the East Gondwana margin. New Zealand Journal of Geology and Geophysics, 47, 189-200. ADAMS, C.J. & GRAHAM, I.J. 1996. Metamorphic and tectonic geochronology of the Torlesse terrane, Wellington, New Zealand. New Zealand Journal of Geology and Geophysics, 39, 157-180. ADAMS, C.J. & GRAHAM, I.J. 1997. Age of metamorphism of Otago schist in eastern Otago and determination of protoliths from initial strontium isotope characteristics. New Zealand Journal of Geology and Geophysics, 40, 275-286. ADAMS, C.J. & KELLEY, S. 1998. Provenance of Permian-Triassic and Ordovician metagraywacke terranes in New Zealand: Evidence from 40 Ar/39Ar dating of detrital micas. Geological Society of America Bulletin, 110, 422^32. ADAMS, C.J. & MAAS, R. 2004. Rb-Sr age and strontium isotopic characterisation of the Torlesse Supergroup in Canterbury, New Zealand, and implications for the status of the Rakaia terrane. New Zealand Journal of Geology and Geophysics, 47, 201-217. ADAMS, C.J. & RAINE, J.I. 1988. Age of Cretaceous silicic volcanism at Kyeburn, central Otago, and Palmerston, eastern Otago, South Island, New Zealand. New Zealand Journal of Geology and Geophysics, 31, 471-475. ADAMS, C.J., BISHOP, D.G. & GABITES, J.E. 1985. Potassium-argon age studies of a low-grade, progressively metamorphosed greywacke

Group) in the Countess Range-Mararoa River area, northwestern Southland, New Zealand. Journal of the Royal Society of New Zealand, 18, 271-284. AITCHISON, J.C, IRELAND, T.R., CLARKE, G.L., CLUZEL, D. & MEFFRE, S. 1998. Regional implications of U/ Pb SHRIMP age constraints on the tectonic evolution of New Caledonia. Tectonophysics, 299, 333-343. ALLIBONE, A.H. & TULLOCH, A.J. 1997. Metasedimentary, granitoid, and gabbroic rocks from central Stewart Island, New Zealand. New Zealand Journal of Geology and Geophysics, 40, 53-68. ALLIBONE, A.H. & TULLOCH, A.J. 2004. Geology of plutonic basement rocks of Stewart Island, New Zealand. New Zealand Journal of Geology and Geophysics, 47, 233-256. ANDREWS, PB., SPEDEN, I.G. & BRADSHAW, J.D. 1976. Lithological and paleontological content of the Carboniferous-Jurassic Canterbury Suite, South Island, New Zealand. New Zealand Journal of Geology and Geophysics, 19, 791-819. ARONSON, J.L. 1965. Reconnaissance rubidium-strontium geochronology of New Zealand plutonic and metamorphic rocks. New Zealand Journal of Geology and Geophysics, 8, 401-423. BALLANCE, P.P. & CAMPBELL, J.D. 1993. The Murihiku arc-related basin of New Zealand (TriassicJurassic). In: BALLANCE, P.E (ed.) South Pacific Sedimentary Basins. Sedimentary Basins of the World, 2. Elsevier, Amsterdam, 21-33. BALLANCE, P.P., HEMING, R.F. & SAMESHIMA, T. 1981. Petrography of the youngest known Murihiku Supergroup, New Zealand: latest Jurassic arc volcanism on the southern margin of Gondwana. In: CRESSWELL, M. & VELLA, P. (eds) Gondwana Five. Balkema, Rotterdam, 161-166. BARNES, P.M. & KORSCH, R.J. 1991. Melange and related structures in Torlesse accretionary wedge, Wairarapa, New Zealand. New Zealand Journal of Geology and Geophysics, 34, 517-532. BASSETT, K.N. & ORLOWSKI, R. 2004. The Pahau Terrane type locality: fan delta in an accretionary prism trench-slope basin. New Zealand Journal of Geology and Geophysics, 47, 603-623. BEGGS, J.M. 1993. Depositional and tectonic history of

NEW ZEALAND TECTONOSTRATIGRAPHY the Great South Basin. In: BALLANCE, P.P. (ed.) South Pacific Sedimentary Basins. Sedimentary Basins of the World, 2. Elsevier, Amsterdam, 365-373. BERESFORD, S.W., BRADSHAW, J.D., WEAVER, S.D. & Mum, R.J. 1996. Echinus Granite and Pepin Group of Pepin Island, Northeast Nelson, New Zealand - Drumduan terrane basement or exotic fragment in the Median Tectonic Zone. New Zealand Journal of Geology and Geophysics, 39, 265-270. BISHOP, D.G., BRADSHAW, J.D., LANDIS, C.A. & TURNBULL, I.M. 1976. Lithostratigraphy and structure of the Caples terrane of the Humboldt Mountains. New Zealand Journal of Geology and Geophysics, 19, 827-848. BISHOP, D.G., BRADSHAW, J.D. & LANDIS, C.A. 1985. Provisional terrane map of South Island, New Zealand. In: HOWELL, G. (ed.) Tectonostratigraphic Terranes of the Circum-Pacific Region. AAPG Circum-Pacific Counsil for Energy and Mineral Recourses. Earth Science Series 1, 515-521. BLACK, P. 19960. Mesozoic evolution of the Norfolk Ridge system: evidence from New Caledonia and northern New Zealand. In: Mesozoic Geology of the Eastern Australia Plate Conference. Geological Society of Australia, Brisbane, 90. BLACK, P. 19966. Omahuta, Bay of Island and Manaia Hill terranes: Waipapa composite terrane, North Island, New Zealand. Geological Society of New Zealand Miscellaneous Publications, 91A, 29. BLACK, P.M., CLARK, A.S.B. & HAWKE, A.A. 1993. Diagenesis and very low-grade metamorphism of volcaniclastic sandstones from contrasting geodynamic environments, North Island, New Zealand; the Murihiku and Waipapa terranes. Journal of Metamorphic Geology, 11, 429-435. BOLES, J.R. 1974. Structure, stratigraphy and petrology of mainly Triassic rocks, Hokonui Hills, Southland, New Zealand. New Zealand Journal of Geology and Geophysics, 17, 337-374. BORG, S.G. & DE PAOLO, D.J. 1991. Tectonic model of the Antarctic Gondwana margin with implications for southeastern Australia: Isotopic and geochemical evidence. Tectonophysics, 196, 339-358. BRADSHAW, J.D. 1973. Allochthonous Mesozoic fossil localities in melange in the Torlesse Group of North Canterbury. Journal of the Royal Society of New Zealand, 3, 161-167. BRADSHAW, J.D. 1989. Cretaceous geotectonic pattern in the New Zealand region. Tectonics, 8, 803-820. BRADSHAW, J.D. 1993. A review of the Median Tectonic Zone: terrane boundaries and terrane amalgamation near the Median Tectonic Line. New Zealand Journal of Geology and Geophysics, 36, 117-125. BRADSHAW, J.D. 1994. Brook Street and Murihiku terranes of New Zealand in the context of a mobile South Pacific Gondwana margin. South American Journal of Earth Sciences, 7, 325-332. BRADSHAW, J.D., ADAMS, C.J. & ANDREWS, P.B. 1981. Carboniferous to Cretaceous on the Pacific margin of Gondwana; the Rangitata phase of New

209

Zealand. In: CRESSWELL, M. & VELLA, P. (eds) Gondwana Five. Balkema, Rotterdam, 217-221. BRADSHAW, J.D., WEAVER, S.D., PANKHURST, R.J., STOREY, B.C. & MUIR, R.J. 1997. New Zealand superterranes recognized in Marie Byrd Land and Thurston Island. Terra Antartica, 3, 429-436. BRADSHAW, J.Y. & KIMBROUGH, D.L. 1991. Mid-Paleozoic age of granitoids in enclaves within Early Cretaceous granulites, Fiordland, southwest New Zealand. New Zealand Journal of Geology and Geophysics, 34, 455-469. BRADSHAW, M.A. 1995. Stratigraphy and structure of the Lower Devonian rocks of Waitahu and Orlando Outliers, near Reefton, New Zealand, and their relationship to the Inangahua Outlier. New Zealand Journal of Geology and Geophysics, 38, 81-92. BRADSHAW, M.A. 2000. Base of the Devonian Baton Formation and the question of a pre-Baton tectonic event in the Takaka Terrane, New Zealand. New Zealand Journal of Geology and Geophysics, 43, 601-610. BRATHWAITE, R.L., KAMO, S.L. & FAURE, K. 2004. U-Pb geochronology and geochemistry of molybdenum-bearing granodiorite porphyry at Copperstain Creek, west Nelson, New Zealand. New Zealand Journal of Geology and Geophysics, 47, 219-225. BROWN, E.H. 1996. High-pressure metamorphism caused by magma loading in Fiordland, New Zealand. Journal of Metamorphic Geology, 14, 441-452. BRYAN, S.E., CONSTANTINE, A.E., STEPHENS, C.J., EWART, A., SCHON, R.W & PARIANOS, J. 1997. Early Cretaceous volcano-sedimentary successions along the Eastern Australian continental margin: implications for the break-up of Eastern Gondwana. Earth and Planetary Science Letters, 153, 85-102. BRYAN, S.E., EWART, A., STEPHENS, C.J., PARIANOS, J. & DOWNES, P.J. 2000. The Whitsunday volcanic province, central Queensland, Australia: lithological and stratigraphic investigations of a silicicdominated large igneous province. Journal of Volcanology and Geothermal Research, 99,55-78. CAMPBELL, H.J. 20000. The marine Permian of New Zealand. In: YIN, H., DICKINS, J.M., SHI, G.R. & TONG, T. (eds) Permian-Triassic Evolution of Tethys and Western Circum-Pacific. Elsevier, Amsterdam, 111-125. CAMPBELL, H.J. 20006. The Marine Triassic of Australasian and its interregional correlation. In: YIN, H., DICKINS, J.M., SHI, G.R. & TONG, T. (eds) Permian-Triassic Evolution of Tethys and Western Circum-Pacific. Elsevier, Amsterdam, 235-255. CAMPBELL, H.J., MORTIMER, N. & RAINE, J.I. 2001. Geology of the Permian Kuriwao Group, Murihiku Terrane, New Zealand. New Zealand Journal of Geology and Geophysics, 44, 485^98. CAMPBELL, J.D. & WARREN, G. 1965. Fossil localities of the Torlesse group in the South Island. Transactions of the Royal Society of New Zealand, Geology, 3, 99-137. CAWOOD, P.A., NEMCHIN, A.A., LEVERENZ, A., SAEED,

210

A. M. WANDRES & J. D. BRADSHAW

A. & BALLANCE, P.F. 1999. U/Pb dating of detrital zircons: implications for the provenance record of Gondwana margin terranes. Geological Society of America Bulletin, 111, 1107-1119. CAWOOD, P. A., LANDIS, C.A., NEMCHIN, A. A. & HADA, S. 2002. Permian fragmentation, accretion and subsequent translation of a low-latitude Tethyan seamount to the high-latitude East Gondwana margin; evidence from detrital zircon age data. Geological Magazine, 139, 131-144. CLARK, G.L., KLEPEIS, K.A. & DACZKO, N.R. 2000. Cretaceous high-pressure granulites at Milford Sound, New Zealand: metamorphic history and emplacement in a convergent margin setting. Journal of Metamorphic Geology, 18, 359-374. CLUZEL, D., AITCHISON, J.C. & PICARD, C. 2001. Tectonic accretion and underplating of mafic terranes in the late Eocene intraoceanic fore-arc of New Caledonia (Southwest Pacific); geodynamic implications. Tectonophysics, 340, 23-59. COLLINSON, J.W., ISBELL, XL., ELLIOT, D.H., MlLLER,

M.F., MILLER, J.M.G. & VEEVERS, JJ. 1994. Permian-Triassic transantarctic basins. In: VEEVERS, J. & POWELL, M. (eds) Permian-Triassic Pangean Basins and Foldbelts along the Panthalassan Margin of Gondwanaland. Geological Society of America, Memoirs, 184,173-222. CONEY, P.J., EDWARDS, A., HINE, R., MORRISON, F. & WINDRIM, D. 1990. The regional tectonics of the Tasman erogenic system, eastern Australia. Journal of Stuctural Geology, 12, 519-543. COOK, R.A., SUTHERLAND, R. & ZHU, H. 1999. Cretaceous-Cenozoic Geology and Petroleum Systems of the Great South Basin, New Zealand. Institute of Geological and Nuclear Sciences, Monographs 20. Institute of Geological and Nuclear Sciences, Lower Hutt. COOMBS, D.S., LANDIS, C.A., NORRIS, R.J., SINTON, J.M., BORNS, DJ. & CRAW, D. 1976. The Dun Mountain Ophiolite Belt, New Zealand, its tectonic setting, constitution and origin, with special reference to the southern portion. American Journal of Sciences, 276, 561-603. COOMBS, D.S., COOK, N.D.J., KAWACHI, Y., JOHNSTONE, R.D. & GIBSON, I.L. 1996. Park Volcanics, Murihiku Terrane, New Zealand - Petrology, petrochemistry, and tectonic significance. New Zealand Journal of Geology and Geophysics, 39, 469-492. COOMBS, D.S., LANDIS, C.A., HADA, S., Pro, M., ROSER, B.P, SUZUKI, T. & YOSHIKURA, S. 2000. The Chrystalls Beach-Brighton block, southeast Otago, New Zealand. New Zealand Journal of Geology and Geophysics, 43, 355-372. COOPER, R.A. 19890. Early Paleozoic terranes of New Zealand. Journal of the Royal Society of New Zealand, 19, 73-112. COOPER, R.A. 19896. New Zealand tectonostratigraphic terranes and panbiogeograhpy. New Zealand Journal of Zoology, 16, 699-712. COOPER, R.A. & TULLOCH, AJ. 1992. Early Palaeozoic terranes in New Zealand and their relationship to the Lachlan fold belt. Tectonophysics, 214, 129-144. Cox, S.C. 1991. The Caples/Aspiring terrane

boundary - the translation surface of an early nappe structure in the Otago Schist. New Zealand Journal of Geology and Geophysics, 34, 73-82. DACZKO, N.R., KLEPEIS, K.A. & CLARKE, G.L. 2002. Thermomechanical evolution of the crust during covergence and deep crustal pluton emplacement in the Western Province of Fiordland, New Zealand. Journal of Metamorphic Geology, 21, 1-18. DAVEY, FJ. & CHRISTOFFEL, D.A. 1978. Magnetic anomalies across Campbell Plateau, New Zealand. Earth and Planetary Science Letters, 41, 14-20. DEPAOLO, DJ. & WASSERBURG, GJ. 1979. Petrogenetic mixing models and Nd-Sr isotopic patterns. Geochimica et Cosmochimica Acta, 43, 615-627. DE SANTIS, L., BRANCOLINI, G. & BUSETTI, M. 1994. Structural evolution of the Victoria Land Basin south of the Drygalsky ice tongue (western Ross Sea). Terra Antartica, 1, 107-110. DICKIN, A.P. 1995. Radiogenic Isotope Geology. Cambridge University Press, Cambridge. DIVENERE, V., KENT, D.V. & DALZIEL, I.W.D. 1995. Early Cretaceous paleomagnetic results from Marie Byrd Land, West Antarctica: implications for the Weddellia collage of crustal blocks. Journal of Geophysical Research, 100, 8133-8151. DOUBLEDAY, P.A., MACDONALD, D.I.M. & NELL,

P.A.R. 1993. Sedimentology and structure of the trench-slope to forearc basin transition in the Mesozoic of Alexander Island, Antarctica. Geological Magazine, 130, 737-754. BADE, J.V. 1988. The Norfolk Ridge System and its margins. In: NAIRN, A.E., STEHLI, F.G. & UYEDA, S. (eds) The Oceanic Basins and Margins, Volume 7B: The Pacific Ocean. Plenum Press, New York, 303-324. EWART, A., SCHON, R.W. & CHAPPELL, B.W. 1992. The Cretaceous volcanic-plutonic province of the central Queensland (Australia) coast: a rift related "calc-alkaline" province. Geological Society of America, Special Papers, 272, 327-345. FERGUSON, R., HOEY, T, WATHEN, S. & WERRITTY, A. 1996. Field evidence for rapid downstream fining of river gravels through selective transport. Geology, 24,179-182. FISCHER, PJ. 1998. Conodonts in melange, Caples terrane, Nokomai area, northern Southland, New Zealand. Geological Society of New Zealand Miscellaneous Publications, 101A, 94. FLOTTMANN, T, GIBSON, GM. & KLEINSCHMIDT, G. 1993. Structural continuity of the Ross and Delamerian orogens of Antarctica and Australia along the margin of the paleo-Pacific. Geology, 21, 319-322. FODEN, J.D., ELBURG, M.A., TURNER, S.P., SANDIFORD, M., O'CALLAGHAN, J. & MITCHELL, S. 2002.

Granite production in the Delamerian Orogen, South Australia. Journal of the Geological Society of London, 159, 557-575. FORD, P.B., LEE, D.E. & FISCHER, PJ. 1999. Early Permian conodonts from the Torlesse and Caples Terranes, New Zealand. New Zealand Journal of Geology and Geophysics, 42, 79-90. FORSTER, M.A. & LISTER, G.S. 2003. Cretaceous

NEW ZEALAND TECTONOSTRATIGRAPHY metamorphic core complexes in the Otago Schist, New Zealand. Australian Journal of Earth Sciences, 50,181-198. FOSTER, D.A. & GRAY, D.R. 2000. Evolution and stucture of the Lachlan fold belt (orogen) of eastern Australia. Annual Review Earth and Planetary Sciences, 28, 47-80. FROST, C.D. & COOMBS, D.S. 1989. Nd isotope character of New Zealand sediments: implications for terrane concepts and crustal evolution. American Journal of Science, 289, 744-770. GAHAGAN, L.M., LAWYER, L.A. & DALZIEL, I.W.D. 1999. Antarctica Keystone of Gondwana. http://www.ig.utexas.edu/research/overview/. A projects\plates\movies\akog.ppt. GEORGE, A.D. 1992. Deposition and deformation of an Early Cretaceous trench-slope basin deposit, Torlesse terrane, New Zealand. Geological Society of America Bulletin, 104, 570-580. GIBSON, G.M. & IRELAND, T.R. 1995. Granulite formation during continental extension in Fiordland, New Zealand. Nature, 375, 479-482. GIBSON, G.M. & IRELAND, T.R. 1996. Extension of Delamerian (Ross) orogen into western New Zealand - evidence from zircon ages and implications for crustal growth along the Pacific margin of Gondwana. Geology, 24,1087-1090. GIBSON, G.M., MCDOUGALL, I. & IRELAND, T.R. 1988. Age constraints on metamorphism and the development of a metamorphic core complex in Fiordland, southern New Zealand. Geology, 16, 405-408. GRAHAM, IJ. & KORSCH, RJ. 1989. Rb-Sr resetting ages and chemical characterization of turbidites in an accretionary wedge: Torlesse Complex, Otaki Gorge, New Zealand. Geological Society of America Bulletin, 101, 355-363. GRAHAM, IJ. & KORSCH, RJ. 1990. Age and provenance of granitoid clasts in Moeatoa Conglomerate, Kawhia Syncline, New Zealand. Journal of the Royal Society of New Zealand, 20, 25-39. GRAHAM, IJ. & MORTIMER, N. 1992. Terrane characterisation and timing of metamorphism in the Otago Schist, New Zealand, using Rb-Sr and K-Ar geochronology. New Zealand Journal of Geology and Geophysics, 35, 391-401. GRAPES, R.H., ROSER, B.P. & KASHAI, K. 2001. Composition of monocrystalline detrital and authigenic minerals, metamorphic grade, and provenance of Torlesse and Waipapa greywacke, Central North Island, New Zealand. Geological Review, 43, 139-163. GRINDLEY, G.W., OLIVER, P.J. & SUKROO, J.C. 1981. Lower Mesozoic position of southern New Zealand determined from paleomagnetism of the Glenham Porphyry, Murihiku Terrane, eastern Southland. In: CRESSWELL, M. & VELLA, P. (eds) Gondwana Five. Balkema, Rotterdam, 319-326. HADA, S. & LANDIS, C.A. 1995. Te Akatarawa Formation: an exotic oceanic-continental margin terrane within the Torlesse-Haast Schist transition zone. New Zealand Journal of Geology and Geophysics, 38, 349-359. HAMILTON, RJ., LUYENDYK, B.P, SORLIEN, C.C. & BARTEK, L.R. 2001. Cenozoic tectonics of the

211

Cape Roberts rift basin and Transantarctic Mountain front, southwestern Ross Sea, Antarctica. Tectonics, 20, 325-342. HILL, EJ. 1995. The Anita Shear Zone - a major, Middle Cretaceous tectonic boundary in northwestern Fiordland. New Zealand Journal of Geology and Geophysics, 38, 93-103. HITCHING, K.D. 1979. Torlesse geology of Kakahu, South Canterbury. New Zealand Journal of Geology and Geophysics, 22, 191-197. HOLDSWORTH, B.K. & NELL, P.A.R. 1992. Mesozoic radiolarian faunas from the Antarctic Peninsula: age, tectonic and paleo-oceanographic significance. Journal of the Geological Society, London, 149, 1003-1020. HOLLIS, J.A., CLARK, G.L., KLEPEIS, K.A., DACZKO, N.R. & IRELAND, T.R. 2003. Geochronology and geochemistry of the Arthur River Complex, Fiordland, New Zealand: Cretaceous magmatism and metamorphism on the Paleo-Pacific margin. Journal of Metamorphic Geology, 21, 299-313. HOWELL, D.G. 1980. Mesozoic accretion of exotic terranes along the New Zealand segment of Gondwanaland. Geology, 8, 487-491. HURLEY, P.M., HUGHES, H., PINSON, W.H. & FAIRBAIRN, H.W. 1962. Radiogenic argon and strontium diffusion parameters in biotite at low temperatures obtained from Alpine Fault uplift in New Zealand. Geochimica et Cosmochimica Acta, 26, 76-80. IRELAND, T.R. 1992. Crustal evolution of New Zealand: evidence from age distributions of detrital zircons in Western Province paragneisses and Torlesse greywacke. Geochimica et Cosmochimica Acta, 56, 911-920. IRELAND, T.R. & GIBSON, G.M. 1998. SHRIMP monazite and zircon geochronology of highgrade metamorphism in New Zealand. Journal of Metamorphic Geology, 16, 149-167. IRELAND, T.R., FLOTTMANN, T, FANNING, C.M., GIBSON, G.M. & PREISS, W.V. 1998. Development of the early Paleozoic Pacific margin of Gondwana from detrital-zircon ages across the Delamerian Orogen. Geology, 26, 243-246. ISAAC, M.J., HERZER, R.H., BROOK, FJ. & HAYWARD, B.W. 1994. Cretaceous and Cenozoic Sedimentary Basins of Northland, New Zealand. Institute of Geological and Nuclear Sciences, Monographs 8. Institute of Geological and Nuclear Sciences, Lower Hutt. ITO, M., YOSHIAKI, A. & SHIGEKI, H. 2000. New radiolarian age information for the Chrystalls Beach Complex, southwest of Dunedin, New Zealand. New Zealand Journal of Geology and Geophysics, 43, 349-354. JOHNSTON, M.R. 1990. Geology of the St. Arnaud district, southeast Nelson (Sheet N29). New Zealand Geological Survey Bulletin, 99, 13-17. JOHNSTON, M.R., RAINE, J.I. & WAITERS, W.A. 1987. Drumduan Group of East Nelson, New Zealand; plant-bearing Jurassic arc rocks metamorphosed during terrane interaction. Journal of the Royal Society of New Zealand, 17, 275-301. JONGENS, R., BRADSHAW, J.D. & FOWLER, A. 2003. The Ballon Melange, northwest Nelson: origin,

212

A. M. WANDRES & J. D. BRADSHAW

structure, and emplacement. New Zealand Journal of Geology and Geophysics, 46, 437-448. KAMP, P.J. 2000. Thermochronology of the Torlesse accretionary complex, Wellington region, New Zealand. Journal of Geophysical Research, 105, 19 253-19 272. KAMP, PJ. 2001. Jurassic age for part of the Rakaia Terrane in Canterbury: Implications for development of the Torlesse accretionary prism. New Zealand Journal of Geology and Geophysics, 44, 185-203. KEAR, D. & MORTIMER, N. 2003. Waipa Supergroup, New Zealand: a proposal. Journal of the Royal Society of New Zealand, 33,149-163. KIMBROUGH, D.L. & TULLOCH, A.J. 1989. Early Cretaceous age of orthogneiss from the Charleston Metamorphic Group, New Zealand. Earth and Planetary Science Letters, 95,130-140. KIMBROUGH, D.L., MATTINSON, J.M., COOMBS, D.S., LANDIS, CA. & JOHNSTON, M.R. 1992. Uraniumlead ages from the Dun Mountain ophiolite belt and Brook Street Terrane, South Island, New Zealand. Geological Society of America Bulletin, 104, 429-443. KIMBROUGH, D.L., TULLOCH, A.J., GEARY, E., COOMBS, D.S. & LANDIS, C.A. 1993. Isotopic ages from the Nelson region of South Island, New Zealand: crustal structure and definition of the Median Tectonic Zone. Tectonophysics, 225, 433-448. KIMBROUGH, D.L., TULLOCH, A.J. & RATTENBURY, M.S. 1994fl. Late Jurassic-Early Cretaceous metamorphic age of Fraser Complex migmatite, Westland, New Zealand. New Zealand Journal of Geology and Geophysics, 37, 137-142. KIMBROUGH, D.L., TULLOCH, A.J., COOMBS, D.S., LANDIS, C.A., JOHNSTON, M.R. & MATTINSON, J.M. 1994&. Uranium-lead zircon ages from the Median Tectonic Zone, South Island, New Zealand. New Zealand Journal of Geology and Geophysics, 37, 393-419. KLEPEIS, K.A., CLARKE, G.L. & RUSHMER,T.R. 2003. Magma transport and coupling between deformation and magmatism in the continental lithosphere. GSA Today, 13, 4-11. KLEPEIS, K.A., CLARKE, G.L., GEHRELS, D. & VERVOORT, J. 2004. Processes controlling vertical coupling and decoupling between the upper and lower crust of orogens: results from Fiordland, New Zealand. Journal of Structural Geology, 26, 765-791. KODAMA, Y. 19940. Downstream changes in the lithology and grain size of fluvial gravels, the Watarase River, Japan: evidence of the role of abrasion in downstream fining. Journal of Sedimentary Research, 64, 68-75. KODAMA, Y. 1994b. Experimental study of abrasion and its role in producing downstream fining in gravel-bed rivers. Journal of Sedimentary Research, 64, 76-85. KORSCH, R.J. & WELLMAN, H.W. 1988. The geological evolution of New Zealand and the New Zealand region. In: NAIRN, A.E.M., STEHLI, EG. & UYEDA, S. (eds) The Oceanic Basins and Margins. Volume 7B: The Pacific Ocean. Plenum Press, New York, 411-482.

KRULL, E.S., RETALLACK, G.J., CAMPBELL, HJ. & LYON, G.L. 2000. Delta C-13(org) chemostratigraphy of the Permian-Triassic boundary in the Maitai Group, New Zealand: evidence for highlatitudinal methane release. New Zealand Journal of Geology and Geophysics, 43, 21-32. LAIRD, M.G. & BRADSHAW, J.D. 2004. The break-up of a long term relationship: the Cretaceous separation of New Zealand from Gondwana. Gondwana Research, 7, 273-286. LANDIS, C.A. & BISHOP, D.G. 1972. Plate tectonics and regional stratigraphic-metamorphic relations in the southern part of the New Zealand geosyncline. Geological Society of America Bulletin, 83, 2267-2284. LANDIS, C.A. & COOMBS, D.S. 1967. Metamorphic belts and orogenesis in southern New Zealand. Tectonophysics, 4, 501-518. LANDIS, C.A., CAMPBELL, H.J., ASLUND,T, ETAL. 1999. Permian-Jurassic strata at Productus Creek, Southland, New Zealand: implications for terrane dynamics of the eastern Gondwanaland margin. New Zealand Journal of Geology and Geophysics, 42, 255-278. LAWYER, L.A. & GAHAGAN, L.M. 1994. Constraints on timing of extension in the Ross Sea region. Terra Antartica, 1, 545-552. LEBRUN, J.F., LAMARCHE, G. & COLLOT, J.Y. 2003. Subduction initiation at a strike-slip plate boundary: the Cenozoic Pacific-Australian plate boundary, south of New Zealand. Journal of Geophysical Research, 108, 2543. LEVEN, E.J. & CAMPBELL, HJ. 1998. Middle Permian (Murgabian) fusuline faunas, Torlesse terrane, New Zealand. New Zealand Journal of Geology and Geophysics, 41,149-156. LEVERENZ, A. & BALLANCE, P.F. 2001. Terrane affiliation and terrane boundaries of Mesozoic accretionary complexes, northeastern North Island, New Zealand: some implications from recycling. New Zealand Journal of Geology and Geophysics, 44, 589-599. LITTLE, T.A., MORTIMER, N. & MCWILLIAMS, M. 1999. An episodic Cretaceous cooling model for the Otago-Marlborough Schist, New Zealand, based on 40Ar/39Ar white mica ages. New Zealand Journal of Geology and Geophysics, 42, 305-325. LUYENDYK, B.P. 1995. Hypothesis for Cretaceous rifting of east Gondwana caused by subducted slab capture. Geology, 23, 373-376. LUYENDYK, B., CISOWSKI, S., SMITH, C, RICHARD, S. & KIMBROUGH, D. 1996. Paleomagnetic study of the northern Ford Ranges, Western Marie Byrd Land, West Antarctica: motion between West and East Antarctica. Tectonics, 15, 122-141. MACKINNON, T. 1983. Origin of the Torlesse terrane and coeval rocks, South Island, New Zealand. Geological Society of America Bulletin, 94, 967-985. MATTINSON, J.M., KIMBROUGH, D.L. & BRADSHAW, J.Y. 1986. Western Fiordland Orthogneiss; Early Cretaceous arc magmatism and granulite facies metamorphism, New Zealand. Contributions to Mineralogy and Petrology, 92, 383-392. MAYES, C.L., LAWYER, L.A. & SANDWELL, D.T. 1990.

NEW ZEALAND TECTONOSTRATIGRAPHY Tectonic history and new isochron chart of the South Pacific. Journal of Geophysical Research, 95, 8543-8567. MAZENGARB, C. & HARRIS, D.H.M. 1994. Cretaceous stratigraphic and structural relations of the Raukumara Peninsula, New Zealand: stratigraphic patterns associated the migration of a thrust system. Annales Tectonicae, 8, 100-118. MCCARRON, JJ. & LARTER, R.D. 1998. Late Cretaceous to early Tertiary subduction history of the Antarctic Peninsula. Journal of the Geological Society, London, 155, 255-268. McCuLLOCH, M.T. & CHAPPELL, B.W 1982. Nd isotopic characteristics of S- and I-type granites. Earth and Planetary Science Letters, 58, 51-64. MORTIMER, N. 1993. Jurassic tectonic history of the Otago Schist, New Zealand. Tectonics, 12, 237-244. MORTIMER, N. 1995. Origin of the Torlesse terrane and coeval rocks, North Island, New Zealand. International Geology Review, 36, 891-921. MORTIMER, N. 2004. New Zealand's geological foundations. Gondwana Research, 1, 261-272. MORTIMER, N. & PARKINSON, D. 1996. Hikurangi Plateau - a Cretaceous large igneous province in the Southwest Pacific Ocean. Journal of Geophysical Research, 101, 687-696. MORTIMER, N. & ROSER, B.P. 1992. Geochemical evidence for the position of the Caples-Torlesse boundary in the Otago Schist, New Zealand. Journal of the Geological Society, London, 149, 967-977. MORTIMER, N, PARKINSON, D., RAINE, J.I., ADAMS, C.J., GRAHAM, I.J., OLIVER, P.J. & PALMER, K. 1995. Ferrar magmatic province rocks discovered in New Zealand: implications for Mesozoic Gondwana geology. Geology, 23, 185-188. MORTIMER, N, TULLOCH, A.J. & IRELAND, T.R. 1997. Basement geology of Taranaki and Wanganui Basins, New Zealand. New Zealand Journal of Geology and Geophysics, 40, 223-236. MORTIMER, N., HERZER, R.H., GANS, P.B., PARKINSON, D.L. & SEWARD, D. 1998. Basement geology from Three Kings Ridge to West Norfolk Ridge, Southwest Pacific Ocean - evidence from petrology, geochemistry and isotopic dating of dredge samples. Marine Geology, 148, 135-162. MORTIMER, N., GANS, P., CALVERT, A. & WALKER, N. 19990. Geology and thermochronometry of the east edge of the Median Batholith (Median Tectonic Zone): a new perspective on Permian to Cretaceous crustal growth of New Zealand. Island Arc, 8, 404-425. MORTIMER, N, TULLOCH, A.J., SPARK, R.N., WALKER, N.W., LADLEY, E., ALLIBONE, A. & KIMBROUGH, D.L. 1999ft. Overview of the Median Batholith, New Zealand: a new interpretation of the geology of the Median Tectonic Zone and adjacent rocks. Journal of African Earth Sciences, 29, 257-268. MORTIMER, N, DAVEY, F.J., MELHUISH, A., Yu, J. & GODFREY, N.J. 2002. Geological interpretation of a deep seismic profile across the eastern Province and Median Batholith, New Zealand: crustal architecture of an extended Phanerozoic conver-

213

gent orogen. New Zealand Journal of Geology and Geophysics, 45, 349-364. MUIR, R.J., IRELAND, T.R., WEAVER, S.D. & BRADSHAW, J.D. 1994. Ion microprobe U-Pb zircon geochronology of granitic magmatism in the Western Province of the South Island, New Zealand. Chemical Geology, 113, 171-189. MUIR, R.J., WEAVER, S.D., BRADSHAW, J.D., EBY, G.N. & EVANS, J.A. 1995. The Cretaceous Separation Point Batholith, New Zealand: granitoid magmas formed by melting of mafic lithosphere. Journal of the Geological Society, London, 152, 689-701. MUIR, R.J., IRELAND, T.R., WEAVER, S.D. & BRADSHAW, J.D. 19960. Ion microprobe dating of Paleozoic granitoids: Devonian magmatism in New Zealand and correlations with Australia and Antarctica. Chemical Geology, 127, 191-210. MUIR, R.J., WEAVER, S.D., BRADSHAW, J.D., EBY, G.N, EVANS, J.A. & IRELAND, T.R. 19966. Geochemistry of the Karamea Batholith, New Zealand and comparisons with the Lachlan Fold Belt granites of southeast Australia. Lithos, 39, 1-20. MUIR, R.J., IRELAND, T.R., WEAVER, S.D., BRADSHAW, J.D., WAIGHT,T.E., JONGENS, R. & EBY, G.N. 1997. SHRIMP U-Pb geochronology of Cretaceous magmatism in northwest Nelson-Westland, South Island, New Zealand. New Zealand Journal of Geology and Geophysics, 40, 453-463. MUIR, R.J., IRELAND, T.R., WEAVER, S.D., BRADSHAW, J.D., EVANS, J.A., EBY, G.N. & SHELLEY, D. 1998. Geochronology and geochemistry of a Mesozoic magmatic arc system, Fiordland, New Zealand. Journal of the Geological Society, London, 155, 1037-1052. MUKASA, S.B. & DALZIEL, I.W.D. 2000. Marie Byrd Land, West Antarctica: Evolution of Gondwana's Pacific margin constrained by zircon U-Pb geochronology and feldspar common-Pb isotopic compositions. Geological Society of America Bulletin, 112, 611-627. MULLER, R., GAINA, C. & CLARK, S. 2000. Seafloor spreading around Australia. In: VEEVERS, J.J. (ed.) Billion-Year Earth History of Australia and Neighbours in Gondwanaland. GEMOG Press, Department of Earth and Planetary Sciences, Macquarie University, Sydney, 18-28. MUNKER, C. & COOPER, R.A. 1997. The early Paleozoic Takaka terrane, New Zealand, as part of the Australian/Antarctic Gondwana margin: new insights from the trace element and isotope geochemistry of Cambrian volcanics. In: BRADSHAW, J.D. & WEAVER, S.D. (eds) Terrane Dynamics 97, International Conference on Terrane Geology. Canterbury University Press, Christchurch, 124-127. MUNKER, C. & COOPER, R. 1999. The Cambrian arc complex of the Takaka terrane, New Zealand: an integrated stratigraphical, paleontological and geochemical approach. New Zealand Journal of Geology and Geophysics, 42, 415-445. OLIVER, G.J.H. 1980. Geology of the granulite and amphibolite facies gneisses of Doubtful Sound, Fiordland, New Zealand. New Zealand Journal of Geology and Geophysics, 23, 27-41. OLIVER, G.J.H. & COGGON, J.H. 1979. Crustal

214

A. M. WANDRES & J. D. BRADSHAW

structure of Fiordland, New Zealand. Tetonophysics, 54, 253-292. OLIVER, P.J., CAMPBELL, J.D. & SPEDEN, I.G. 1982. The stratigraphy of Torlesse rocks of the Mt. Somers area (S81), mid-Canterbury. Journal of the Royal Society of New Zealand, 12, 343-271. PALAIS, D.G., MUKASA, S.B. & WEAVER, S.D. 1993. U-Pb and 40Ar/39Ar geochronology for plutons along the Ruppert and Hobbs coasts, Marie Byrd Land, West Antarctica: evidence for rapid transition from arc to rift related magmatism. EOS, Transactions, American Geophysical Union, Spring Meeting Supplement, 74(16), 123. PANKHURST, R.J., MILLAR, I.L., GRUNOW, A.M. & STOREY, B.C. 1993. The pre-Cenozoic magmatic history of the Thurston Island crustal block, West Antarctica. Journal of Geophysical Research, 98, 11835-11849. PANKHURST, R.J., WEAVER, S.D., BRADSHAW, J.D., STOREY, B.C. & IRELAND, T.R. 19980. Geochronology and geochemistry of pre-Jurassic superterranes in Marie Byrd Land, Antarctica. Journal of Geophysical Research, 103, 2529-2547. PANKHURST, R.J., LEAT, P.T., SRUOGA, P., RAPELA, C.W., MARQUEZ, M., STOREY, B.C. & RILEY, T.R. 1998/?. The Chon Aike province of Patagonia and related rocks in West Antarctica: a silicic large igneous province. Journal of Volcanology and Geothermal Research, 81, 113-136. PANKHURST, R.J., RILEY, T.R., FANNING, C.M. & KELLEY, S.P. 2000. Episodic silicic volcanism in Patagonia and the Antarctic Peninsula: chronology of magmatism associated with the break-up of Gondwana. Journal of Petrology, 41, 605-625. PELL, S.D., WILLIAMS, I.S. & CHIVAS, A.R. 1997. The use of protolith zircon-age fingerprints in determining the protosource areas for some Australian dune sands. Sedimentary Geology, 109, 233-260. PETFORD, N. & ATHERTON, M. 1996. Na-rich partial melts from newly underplated basaltic crust; the Cordillera Blanca batholith, Peru. Journal of Petrology, 37, 1491-1521. PICKARD, A.L., ADAMS, CJ. & BARLEY, M.E. 2000. Provenance of New Zealand Late Permian to Cretaceous depocentres at the Australian margin of Gondwanaland: evidence from detrital zircon SHRIMP dating. Australian Journal of Earth Sciences, 47, 987-1007. PICKETT, D.A. & WASSERBURG, GJ. 1989. Neodymium and strontium isotopic characteristics of New Zealand granitoids and related rocks. Contributions to Mineralogy and Petrology, 103, 131-142. POWELL, C.M. & Li, Z.X. 1994. Reconstruction of the Panthalassan margin of Gondwanaland. In: VEEVERS, I & POWELL, M. (eds) Permian-Triassic Pangean Basins and Foldbelts along the Panthalassan Margin of Gondwanaland. Geological Society of America, Memoirs, 184, 5-9. RAINE, J.I., STRONG, C.P & WILSON, GJ. 1993. Biostratigraphic Revision of Petroleum Exploration Wells, Great South Basin, New Zealand. Institute of Geological and Nuclear Sciences, Science Reports 93/32. RATTENBURY, M.S. 1991. The Fraser Complex; high-

grade metamorphic, igneous and mylonitic rocks in central Westland, New Zealand. New Zealand Journal of Geology and Geophysics, 34, 23-33. RATTENBURY, M.S., COOPER, R.A. & JOHNSTON, M.R. 1998. Geology of the Nelson Area. Institute of Geological & Nuclear Sciences 1:250000 Geological Map 9. Institute of Geological & Nuclear Sciences, Lower Hutt. RILEY, T.R., LEAT, P.T., PANKHURST, R.J. & HARRIS, C. 2001. Origins of large volume rhyolitic volcanism in the Antarctic Peninsula and Patagonia by crustal melting. Journal of Petrology, 42, 1043-1065. ROSER, B.P. 1986. Chemistry, provenance and tectonic setting of New Zealand greywacke terranes. Geological Society of New Zealand, 16th Annual Conference; Programme and Abstracts. Geological Society of New Zealand, Christchurch, 88. ROSER, B.P. & COOPER, A.F. 1990. Geochemistry and terrane affiliation of Haast Schist from the western Southern Alps, New Zealand. New Zealand Journal of Geology and Geophysics, 33, 1-10. ROSER, B.P. & KORSCH, R.J. 1986. Determination of tectonic setting of sandstone-mudstone suites using SiO2 content and K2O/Na2O ratio. The Journal of Geology, 94, 635-650. ROSER, B.P. & KORSCH, R.J. 1988. Provenance signatures of sandstone-mudstone suites determined using discriminant function analysis of majorelement data. Chemical Geology, 67,119-139. ROSER, B.P. & KORSCH, R.J. 1999. Geochemical characterization, evolution and source of a Mesozoic accretionary wedge: the Torlesse terrane, New Zealand. Geological Magazine, 136, 493-512. ROSER, B.P, MORTIMER, N.,TURNBULL, I.M. & LANDIS, C.A. 1993. Geology and geochemistry of the Caples terrane, Otago, New Zealand: compositional variations near a Permo-Triassic arc margin. In: BALLANCE, P.F. (ed.) South Pacific Sedimentary Basins. Sedimentary Basins of the World, 2. Elsevier, Amsterdam, 3-19. ROSER, B.P, COOPER, R.A., NATHAN, S. & TULLOCH, A.J. 1996. Reconnaissance sandstone geochemistry, provenance, and tectonic setting of the lower Paleozoic terranes of the West Coast and Nelson, New Zealand. New Zealand Journal of Geology and Geophysics, 39, 1-16. SCHEIBNER, E. & VEEVERS, J.J. 2000. Tasman fold belt system. In: VEEVERS, J.J. (ed.) Billion-Year Earth History of Australia and Neighbours in Gondwanaland. GEMOG Press, Department of Earth and Planetary Sciences, Macquarie University, Sydney, 154-234. SCHMIDT, P.W. & CLARK, D.A. 2000. Paleomagnetism, apparent polar-wander path & paleolatitude. In: VEEVERS, J.J. (ed.) Billion-Year Earth History of Australia and Neighbours in Gondwanaland. GEMOG Press, Department of Earth and Planetary Sciences, Macquarie University, Sydney, 12-17. SHI, GR. & GRUNT, T.A. 2000. Permian Gondwanaboreal antitropicality with special reference to

NEW ZEALAND TECTONOSTRATIGRAPHY brachiopod faunas. Palaeogeogmphy, Palaeoclimatology, Palaeoecology, 155, 239-263. SlLBERLING, N.J., NlCHOLS, K.M., BRADSHAW, J.D. &

BLOME, C.D. 1988. Limestone and chert in tectonic blocks from the Esk Head subterrane, South Island, New Zealand. Geological Society of America Bulletin, 100, 1213-1223. SIRCOMBE, K.N. 1999. Tracing provenance through the isotope ages of littoral and sedimentary detrital zircons, Eastern Australia. Sedimentary Geology, 124, 47-67. SIVELL, W.J. & McCuLLOCH, M.T. 2000. Reassessment of the origin of the Dun Mountain Ophiolite, New Zealand: Nd-isotopic and geochemical evolution of magma suites. New Zealand Journal of Geology and Geophysics, 43, 133-146. SIVELL, W.J. & RANKIN, P.C. 1983. Arc-tholeiite and ultramafic cummulate, Brook Street Volcanics, west of D'Uurville Island, New Zealand. New Zealand Journal of Geology and Geophysics, 26, 239-257. SMALE, D. 1997. Heavy minerals in the Torlesse terrane of the Wellington area, New Zealand. New Zealand Journal of Geology and Geophysics, 40, 499-506. SMALE, D., CAMPBELL, H., GRAPES, R., MORTIMER, N. & RAINE, I. 1996. Permian/Triassic sedimentary sequences in northwestern South Island, New Zealand, and correlation with Australian Gondwanaland sequences. Mesozoic Geology of the Eastern Australia Plate Conference. Geological Society of Australia, Brisbane, 510. SPEDEN, I.G. 1974. Additional Fossil Localities of the Torlesse Rocks of the South Island (Including Summations of Faunal Assemblages and Lithological Information). New Zealand Geological Survey, Reports 69. SPEDEN, I.G. 1976. Fossil localities in Torlesse rocks of the North Island, New Zealand. Journal of the Royal Society of New Zealand, 6, 73-91. SPELL, T.L., MCDOUGALL, I. & TULLOCH, A.J. 2000. Thermochronologic constraints on the breakup of the Pacific Gondwana margin: the Paparoa metamorphic core complex, South Island, New Zealand. Tectonics, 19, 433-451. SPORLI, K.B. 1978. Mesozoic tectonics, North Island, New Zealand. Geological Society of America Bulletin, 89, 415-425. SPORLI, K.B. & BALLANCE, P.P. 1989. MesozoicCenozoic ocean floor/continent interaction and terrane configuration, southwest Pacific area around New Zealand. In: BEN-AVRAHAM, Z. (ed.) The Evolution of the Pacific Ocean Margins. Oxford Monographs on Geology and Geophysics, 8. Oxford University Press, New York, 176-190. STERN, C. & DE WIT, M. 2003. Rocas Verdes ophiolites, southernmost South America: remnants of progressive stages of development of oceanictype crust in a continental margin back-arc basin. In: DILEK, Y. & ROBINSON, P. (eds) Ophiolites in Earth History. Geological Society, London, Special Publications, 218, 1-19. STOREY, B.C., LEAT, P.T., WEAVER, S.D., PANKHURST, R.J., BRADSHAW, J.D. & KELLEY, S. 1999. Mantle

215

plumes and Antarctica-New Zealand rifting; evidence from Mid-Cretaceous mafic dykes. Journal of the Geological Society, London, 156, 659-671. STUMP, E. 1995. The Ross Orogen of the Transantarctic Mountains. Cambridge University Press, Cambridge. STUMP, E., WHITE, A J.R. & BORG, S.G. 1986. Reconstruction of Australia and Antarctica: evidence from granites and recent mapping. Earth and Planetary Science Letters, 79, 348-360. SUGGATE, R.P., STEVENS, G.R. & TE PUNGA, M.T. (eds) 1978. The Geology of New Zealand. Government Printer, Wellington. SUTHERLAND, R. 1999. Basement geology and tectonic development of the greater New Zealand region: an interpretation from regional magnetic data. Tectonophysics, 308, 341-362. TAPPENDEN, V.E., HORNLE, K., WEAVER, S. & IRELAND, T.R. 2002. Slab detachment during the initial stages of mid-Cretaceous rifting of New Zealand from Gondwana. In: STOREY, B.C. (ed.) Gondwana 11, Correlations and Connections. University of Canterbury, Christchurch, New Zealand, 87. TRANTER, T.H. 1992. Underplating of an accretionary prism: an example from the LeMay Group of central Alexander Island, Antarctic Peninsula. Journal of South American Earth Sciences, 6, 1-20. TULLOCH, A.J. & KIMBROUGH, D.L. 1989. The Paparoa metamorphic core complex, New Zealand; Cretaceous extension associated with fragmentation of the Pacific margin of Gondwana. Tectonics, 8, 1217-1235. TULLOCH, A.J. & KIMBROUGH, D.L. 1995. Mesozoic plutonism in western New Zealand: a 50 Ma transition from subduction to continental rifting on the Gondwana margin. In: BROWN, M. & PICCOLI, P.M. (eds) The Origin of Granites and Related Rocks. US Geological Survey Circulars, C1129, 153-154. TULLOCH, A.J. & KIMBROUGH, D.L. 2003. Paired plutonic belts in convergent margins and the development of high Sr/Y: Peninsular Ranges batholith of Baja-California and Median batholith of New Zealand. Geological Society of America, Special Papers, 374, 275-295. TULLOCH, A.J., KIMBROUGH, D.L., LANDIS, C.A., MORTIMER, N. & JOHNSTON, M.R. 1999. Relationships between the Brook Street Terrane and Median Tectonic Zone (Median Batholith): evidence from Jurassic conglomerates. New Zealand Journal of Geology and Geophysics, 42, 279-293. TURNBULL, I.M. 19790. Stratigraphy and sedimentology of the Caples terrane of the Thomson Mountains, northern Southland, New Zealand. New Zealand Journal of Geology and Geophysics, 22, 555-574. TURNBULL, I.M. 19796. Petrography of the Caples terrane of the Thomson Mountains, northern Southland, New Zealand. New Zealand Journal of Geology and Geophysics, 22, 709-727. URUSKI, C.I. 2003. Cretaceous paleogeography of the

216

A. M. WANDRES & J. D. BRADSHAW

Taranaki Basin. In: MORTIMER, N. & LEE, D.E. (eds) Geological Society of New Zealand Miscellaneous Publications, 116A. University of Otago, Dunedin, 150. VEEVERS, J.J. 2000. Antarctic Beadmore-Ross and Mirney provenances saturate PaleozoicMesozoic East Gondwanaland with 0.6-0.5 Ga zircons. In: VEEVERS, J.J. (ed.) Billion-Year Earth History of Australia and Neighbours in Gondwanaland. GEMOG Press, Department of Earth and Planetary Sciences, Macquarie University, Sydney, 110-130. WAIGHT,T.E. 1995. The Geology and Geochemistry of the Hohonu Batholith and Adjacent Rocks, North Westland, New Zealand. PhD thesis, University of Canterbury, Christchurch. WAIGHT,T.E., WEAVER, S.D., IRELAND, T.R., MAAS, R., MUIR, RJ. & SHELLEY, D. 1997. Field characteristics, petrography, and geochronology of the Hohonu Batholith and the adjacent Granite Hill Complex, North Westland, New Zealand. New Zealand Journal of Geology and Geophysics, 40, 1-17. WAIGHT, T.E., WEAVER, S.D., MUIR, R.J., MAAS, R. & EBY, G.N. 1998. The Hohonu Batholith of North Westland, New Zealand; granitoid compositions controlled by source H2O contents and generated during tectonic transition. Contributions to Mineralogy and Petrology, 130, 225-239. WANDRES, A.M., WEAVER, S.D., SHELLEY, D. & BRADSHAW, J.D. 1998. Change from calc-alkaline to adakitic magmatism recorded in the Early Cretaceous Darran Complex, Fiordland, New Zealand. New Zealand Journal of Geology and Geophysics, 41, 1-14. WANDRES, A.M., BRADSHAW, J.D., WEAVER, S.D., MAAS, R., IRELAND, T.R. & EBY, G.N. 20040. Provenance of the sedimentary Rakaia terrane, Torlesse Terrane, South Island, New Zealand: the use of igneous clast compositions to define the source. Sedimentary Geology, 168, 193-226. WANDRES, A.M.C., BRADSHAW, J.D., WEAVER, S.D., MAAS, R., IRELAND, T.R. & EBY, G.N. 20046. Provenance analysis using conglomerate clast lithologies: a case study from the Pahau Terrane of New Zealand. Sedimentary Geology, 167, 57-89. WANDRES, A.M., BRADSHAW, J.D. & IRELAND, T. 2005. The Paleozoic-Mesozoic recycling of the Rakaia Terrane, South Island, New Zealand: sandstone clast and sandstone petrology, geochemistry, and geochronology. New Zealand Journal of Geology and Geophysics, 48, in press. WARD, C.M. 1986. The Fanny and Goodyear terranes of southern Fiordland and their relations with West Nelson. In: STEWART, M. (ed.) Abstracts from the Lower Paleozoic Workshop. Geological Society of New Zealand Miscellaneous Publications, 34. Geological Society of New Zealand, Christchurch, 46-47.

WAREHAM, CD., MILLAR, I.L. & VAUGHAN, A.P.M. 1997. The generation of sodic granite magmas, Western Palmer Land, Antarctic Peninsula. Contributions to Mineralogy and Petrology, 128, 81-96. WEAVER, S.D. & PANKHURST, R.J. 1991. A precise Rb-Sr age for the Mandamus igneous complex, North Canterbury, and regional tectonic implications. New Zealand Journal of Geology and Geophysics, 34, 341-345. WEAVER, S.D., BRADSHAW, J.D. & ADAMS, C.J. 1991. Granitoids of the Ford Ranges, Marie Byrd Land, Antarctica. In: THOMSON, M.R., CRAME, J.A. & THOMSON, J.W. (eds) Geological Evolution of Antarctica. Cambridge University Press, Cambridge, 345-351. WEAVER, S.D., STOREY, B.C., PANKHURST, R.J., MUKASA, S.B., DlVENERE, V.J. & BRADSHAW, J.D.

1994. Antarctica-New Zealand rifting and Marie Byrd Land lithospheric magmatism linked to ridge subduction and mantle plume activity. Geology, 22, 811-814. WILLIAMS, J.G. 1978. Eglinton Volcanics - stratigraphy, petrography, and metamorphism. New Zealand Journal of Geology and Geophysics, 21, 713-732. WILLIAMS, J.G. & HARPER, C.T. 1978. Age and status of the Mackay Intrusives in the Eglinton-upper Hollyf ord area. New Zealand Journal of Geology and Geophysics, 21, 733-742. WOMBACHER, F. & MuNKER, K. 2000. Pb, Nd, and Sr Isotopes and REE systematics of Cambrian sediments from New Zealand: implications for the reconstruction of the Early Paleozoic Gondwana margin along Australia and Antarctica. Journal of Geology, 108, 663-686. WOOD, R. 1994. Geophysical aspect of the opening of the Tasman Sea. In: VAN DER LINGEN, G.J., SWANSON, K.M. & MUIR, R.J. (eds) Evolution of the Tansman Sea Basin. Balkema, Rotterdam, 19-34. WOOD, R. & DAVY, B. 1994. The Hikurangi Plateau. Marine Geology, 118, 153-173. WOOD, R.A. & HERZER, R.H. 1993. The Chatham Rise, New Zealand. In: BALLANCE, P.F. (ed.) South Pacific Sedimentary Basins. Sedimentary Basins of the World, 2. Elsevier, Amsterdam, 329-349. WOOD, R., ANDREWS, P.B. & HERZER, R.H. 1989. Cretaceous and Cenozoic Geology of the Chatham Rise Region, South Island, New Zealand. New Zealand Geological Survey Basin Studies, 3. New Zealand Geological Survey, Lower Hutt. WYSOCZANSKI, R.J., GIBSON, GM. & IRELAND, T.R. 1997. Detrital zircon age patterns and provenance in Late Paleozoic Early Mesozoic New Zealand; terranes and development of the PaleoPacific Gondwana margin. Geology, 25, 939-942.

Pacific subduction coeval with the Karoo mantle plume: the Early Jurasssic Subcordilleran belt of northwestern Patagonia C. W. RAPELA1, R. J. PANKHURST2, C. M. FANNING3 & R HERVE4 Centro de Investigaciones Geologicas, Calle 1 No. 644,1900 La Plata, Argentina (e-mail: crapela@cig. museo. unlp. edu. ar) 2 NERC Isotope Geosciences Laboratory, Key worth, Nottingham NG12 5GG, UK (e-mail: r.pankhurst@nigl. nerc. ac. uk) 3 PRISE, Research School of Earth Sciences, The Australian National University, Mills Road, Canberra ACT 0200, Australia (e-mail: [email protected]) 4 Departamento de Geologia, Universidad de Chile, Casilla 13518, Correo 21, Santiago, Chile (e-mail: fherve@cec. uchile. cl) l

Abstract: The Early Mesozoic magmatism of southwestern Gondwana is reviewed in the light of new U-Pb SHRIMP zircon ages (181 ± 2 Ma, 181 ± 3 Ma, 185 ± 2 Ma, and 182 ± 2 Ma) that establish an Early Jurassic age for the granites of the Subcordilleran plutonic belt in northwestern Argentine Patagonia. New geochemical and isotopic data confirm that this belt represents an early subduction-related magmatic arc along the proto-Pacific margin of Gondwana. Thus, subduction was synchronous with the initial phase of Chon Aike rhyolite volcanism ascribed to the thermal effects of the Karoo mantle plume and heralding rifting of this part of the supercontinent. Overall, there is clear evidence that successive episodes of calc-alkaline arc magmatism from Late Triassic times until establishment of the Andean Patagonian batholith in the Late Jurassic involved westerly migration and clockwise rotation of the arc. This indicates a changing geodynamic regime during Gondwana break-up and suggests differential rollback of the subducted slab, with accretion of new crustal material and/or asymmetrical 'scissor-like' opening of back-arc basins. This almost certainly entailed dextral displacement of continental domains in Patagonia.

The causal links between subduction-related magmatism, mantle plumes and Gondwana break-up during Jurassic-Late Cretaceous time is a major issue in Earth history (e.g. Storey 1995). Ultimate answers to many questions on this topic may only come from studies of mantle dynamics, but precise geochronological, geochemical and isotope studies can provide important constraints to models of supercontinent break-up. Patagonia, in southernmost South America, is a key area in the break-up history of SW Gondwana as it preserves the largest acid magmatic province developed during rifting (the Chon Aike Province, Pankhurst et al 1998 and references therein) and, subsequently, protracted subduction-related magmatism on the Pacific margin (Fig. 1). Together, these provide a complete magmatic record of pre-, syn-, and post-break-up episodes. The timing and characteristics of the Early Cretaceous stage of supercontinent break-up that culminated with the eruption of the Parana-

Etendeka flood basalts and the opening of the South Atlantic is well known (Turner et al 1994 and references therein), whereas the geodynamic environment in the proto-Pacific margin of SW Gondwana during initial rifting in the Early Jurassic (c. 180 Ma; Storey 1995) is still poorly understood. This paper presents new U-Pb SHRIMP data for the emplacement of Early Jurassic granitic rocks in the Subcordilleran belt of northwestern Patagonia (Fig. 1). Their crystallization ages are used, together with geochemical and Sr-Nd isotopic data, to investigate the generation, spatial distribution and migration with time of I-type magmatism in Patagonia, from Late Triassic to Cretaceous times. The significance of this magmatic history is reviewed in terms of a possible palaeogeographical reconstruction of southwestern Gondwana and other break-up magmatism including the Jurassic rhyolites of Patagonia and the Karoo-Ferrar basaltic provinces.

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 217-239. 0305-8719/$15.00 © The Geological Society of London 2005.

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Fig. 1. Distribution of Triassic and Jurassic magmatism and location of the main sedimentary basins in Patagonia (modified from Rapela 2001). The boundary of the Subcordilleran belt (cross-hatched area inside the box) is after the palaeogeographical reconstruction of the Liassic basin by Lizuain (1999), while the main areas of the Triassic-Early Jurassic accretionary prism in the Pacific margin are from Thomson & Herve (2002). Isotopic ages for the Chon Aike Province are from Pankhurst et d. (2000): rocks of V1 (c. 188-178 Ma) occur in the eastern North Patagonian Massif; those of V2 (c. 172-162 Ma) and V3 (c. 157-153 Ma and younger) are not distinguished. See text for data sources and discussion of the isotopic ages of the I-type granites: the coastal batholith (and the sedimentary basins outlined) are largely post-break-up.

SUBCORDILLERAN MAGMATISM OF PATAGONIA

Late Triassic to Jurassic igneous episodes in Patagonia Batholith of Central Patagonia and Deseado Monzonite Suite The Late Triassic 'Batholith of Central Patagonia' (Fig. 1), represents the culmination of Itype granite magmatism that started in Permian times across the North Patagonian Massif (Pankhurst et al 1992; Rapela et al 1992). Departing from the typical N-S Andean alignment, the Batholith of Central Patagonia is associated with a NW-SE fault system that extends from Lago Panguipulli in Chile to the Gastre area in central Patagonia. Rapela (2001) reviewed its lithology, geochemistry and isotopic characteristics. Its known chronology is based on Rb-Sr isochrons defined in the Gastre area, indicating two separate episodes in the Late Triassic (Gastre Superunit 220 ± 3 Ma and Lipetren Superunit 208 ± 1 Ma, Rapela et al 1992). The Triassic age of the Gastre Superunit is now supported by the authors' unpublished U-Pb zircon age of 221 ± 2 Ma. Satellite plutons are located to the north of the batholith, in the North Patagonian Massif (Rb-Sr isochrons of 210 ± 2 Ma and 210 ± 9 Ma, Cingolani et al 1991; Rapela et al 1996). The volcanic counterparts of these Late Triassic plutonic rocks are calc-alkaline rhyolitic ignimbrites interbedded with air-fall deposits carrying a rich Dicroidium and associated flora (Rb-Sr isochron of 222 ± 2 Ma, Rapela et al 1996). Finally, the southernmost granitic outcrops correlated with the Batholith of Central Patagonia belong to the Deseado Monzonite Suite, located in the northeastern sector of the Deseado Massif (Fig. 1). These rocks are characterized by the abundance of quartz monzodiorite and quartz monzonite (Godeas 1993; Rapela & Pankhurst 1996) and have been dated as earliest Jurassic by precise Rb-Sr isochrons, at 202 ± 2 Ma and 203 ± 2 Ma (Pankhurst et al 1993).

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their precise definitions are subject to modification as new data become available - for example, Pankhurst et al (20030) have reported U-Pb zircon ages of 138 ± 2 Ma (i.e. Early Cretaceous) for rhyolites in Chile that are correlated traditionally with the V3 group in Argentina. This paper retains these ranges, extended slightly to allow for analytical errors in the published age determinations. The Early and Middle Jurassic pulses do not have coeval cordilleran counterparts in the main body of the Patagonian batholith and have been ascribed to crustal melting due to spreading of the Karoo plume head (Pankhurst et al 2000; Riley et al 2001). The Middle and Late Jurassic pulses show the geochemical influences of an active margin (Riley et al 2001) and the youngest pulse of Chon Aike rhyolite volcanism (V3) starts before, but partially overlaps in age, the earliest units of the southern Patagonian batholith.

Subcordilleran belt

The Subcordilleran belt (SCB) of Rio Negro and Chubut provinces is a discontinuous belt of Early Jurassic igneous rocks and Liassic sediments that extends for more than 250 km immediately to the east of the North Patagonian cordillera (Figs 1, 2). The NNW-SSW elongated basin is delimited by a dominantly deltaic and tidal marine facies of Liassic black shales, siltstones, conglomerates, quartz-feldspar sandstones and limestones, with volcanic and volcaniclastic intercalations (Lizuain 1999 and references therein). Because of the contemporaneous volcanism (and local intrusion by gabbroic bodies), this basin was assigned an intra-arc setting by Page & Page (1999). The littoral deposits of the basin are abnormally thick for a transgressive succession, suggesting development of coastal cliffs and shorelines undergoing tectonic uplift (Gonzalez Bonorino 1990). Black mudstones, siltstones and volcaniclastic rocks are found in deep boreholes through the San Jorge basin at c. 46° S; Chon Aike Province 69°30' W and are the southernmost reported The Chon Aike Province is one of the largest occurrence of the Liassic marine sediments rhyolitic provinces in the world (Fig. 1) and also (Uliana & Legarreta 1999). extends to the Antarctic Peninsula. Silicic The plutonic rocks of the SCB were called eruptions started with an Early Jurassic episode the 'Subcordilleran Patagonian Batholith' by in northeast Patagonia, migrated to southern Gordon & Ort (1993), and renamed the Patagonia in Middle Jurassic time and finally 'Subcordilleran Plutonic Belt' by Haller et al moved to the Andean cordillera during the Late (1999) on the grounds that a single continuous Jurassic (Feraud etal 1999; Pankhurst etal 2000 body is not seen. Subalkaline norite and olivine and references therein). Pankhurst et al (2000) gabbro are abundant in the sierras of Tepuel identified these three phases as Vj (188- and Tecka in the southern part of the belt (Page 178 Ma), V2 (172-162 Ma), and V3 (157-153 Ma); & Page 1999; Fig. 2). In the northern part of the

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Fig. 2. Simplified geological map of the Subcordilleran sector between 41°30'and 44°15' S.

belt, granitic and gabbroic plutons are intruded into foliated granites and metamorphic rocks of Palaeozoic age, while in the southern sector they were intruded or emplaced technically in

Carboniferous sediments of the Tepuel basin or the Liassic sediments of the SCB. The gabbros are intruded, in turn, by Cretaceous satellite plutons of the Patagonian batholith (Page &

SUBCORDILLERAN MAGMATISM OF PATAGONIA

Page 1999). The granitic rocks of the SCB are zoned plutons dominated by I-type biotitehornblende granodiorite and quartz monzodiorite, and biotite monzogranite, with minor diorite and leucogranite (Spikermann et al 1988, 1989; Gordon & Ort 1993; Busteros et al 1993; Haller et al. 1999).

Previous studies and sampling The Alto Rio Chubut area was studied by Gordon & Ort (1993) who recognized two main units: the Arroyo La Tuerta granodiorite dominated by hornblende-biotite granodiorites and tonalites with abundant mafic enclaves and synplutonic dykes; and the La Angostura granite composed of biotite monzogranites and granodiorites with occasional hornblende. Minor gabbro bodies also occur associated with the granites. Based on field relationships the main units were interpreted as coeval and comagmatic, defining a zoned-type intrusion, with Rb-Sr isochron ages of 200 ± 24 Ma for the Arroyo La Tuerta granodiorite and 183 ± 13 Ma for La Angostura granite (Gordon & Ort 1993). The Leleque area was studied by Lizuain (1983) and Gordon & Ort (1993), who recognized two units: a small gabbro pluton and a dominant hornblende-biotite granodioritic to monzogranitic phase cutting the gabbro. The modal compositions of the granites in the Leleque and Alto Rio Chubut areas are indistinguishable (Rapela 2001). A K-Ar age of 141 ± 5 Ma was reported for the Leleque unit by Lizuain (1983), while K-Ar ages of 177 ± 5 Ma and 173 ± 10 Ma were obtained in diorites and tonalites at the northern margin of Lago Puelo near the Leleque area (Lizuain 1981). Sample SER-046 (41°58'04.4"S;71°16'13.7"W) is a weathered biotite monzogranite from La Angostura granite in the southern sector of the Alto Rio Chubut area, whereas LEL-052 (42°21'22.6" S; 71°iri0.7" W) is a hornblende-biotite granodiorite from the dominant suite at Leleque, which carries microgranular mafic enclaves of 6-15 cm. The Aleusco batholith in the Precordillera of Chubut was first described by Turner (1982) and Spikermann et al. (1988; 1989). Spikermann et al. (1988) and Haller etal (1999) reported K-Ar ages of 180 ± 10 Ma, and 177 ± 6 Ma, 179 ± 7 Ma and 184 ± 6 Ma, respectively. The Aleusco body is a composite epizonal pluton emplaced in Liassic sediments. It consists of a main facies of hornblende-biotite granodiorite and subordinate diorite with abundant mafic enclaves, and late plutonic leucogranite and aplite. Sample ALE-055 (43°06'14.6" S; 70°28'39.7" W) is a

221

typical hornblende-biotite granodiorite of the main facies (Fig. 3). Near the locality of Jose de San Martin in Chubut province, a string of NE-SW granitic bodies, emplaced mainly into Carboniferous and Liassic sediments, are the southernmost outcrops of the SCB. The petrology and chemical characteristics of these granitoids have been described by Spikermann (1978), Franchi & Page (1980) and Busteros et al (1993). Franchi & Page (1980) reported K-Ar ages of 167 ± 30 Ma, 197 ± 10 Ma and 207 ± 10 Ma for these rocks, which they referred to as the Jose de San Martin Formation. The dominant lithology is medium-K metaluminous granodiorite varying to monzodiorite. The largest exposure is a 10 km long, 2 km wide pluton with a thermal aureole in Upper Palaeozoic sediments, 6 km to the east of the town of Jose de San Martin. Sample JSM-058 (44W27.5" S; 70°23'39.6" W) is a hornblende-biotite quartz monzodiorite when classified by modal proportions, and a tonalite in the normative Ab-An-Or diagram (Fig. 3). The abundant subalkaline gabbros of the southern sector of the SCB (Fig. 2) are recognized as a separate lithostratigraphic unit, the Tecka Formation, which is dominated by norite and olivine gabbro, with minor anorthosite, peridotite and troctolite (Page & Page 1999).

Fig. 3. Normative Ab-An-Or diagram for the Mesozoic and Cenozoic I-type granites of Patagonia. Data for the Subcordilleran belt are from Gordon & Ort (1993), Busteros et al. (1993), Haller et al. (1999) and this paper; for the Deseado Monzonite Suite from Rapela et al. (1996); and for the Batholith of Central Patagonia from Rapela et al. (1992). The normative composition of the Patagonian batholith is based on 415 chemical analyses reported for different latitudes of the body: 53° S (Bruce 1988), 48° S (Weaver 1988), 45-46° S (Wells 1978), 44-46° S (Pankhurst et al. 1999), 43° S (Haller 1985), 42° (Ghiara etal. 1997) and 40-41° S (Rapela unpublished data).

222

C.W.RAPELA£rAL.

An 40Ar/36Ar age of 182.7 ± 1 Ma has been reported for a gabbro of the Tecka Formation (Feraud et al unpublished data cited by Page & Page 1999), while a U-Pb SHRIMP age of 19 ± 3 Ma was obtained within a complex zircon population with Precambrian inheritance in an Opx-Cpx gabbro on the southern shore of Lago Fontana (Muzzio gabbro, Rolando et al 2002). Two samples were collected from the Tecka Formation, a Cpx-Opx quartz gabbro near Quichaura (QUI-225; 43°31'30.9" S; 70°39'34.4" W) and an Ol-Cpx-Opx gabbro west of Gualjaina (QUI-227; 42°45/01.9"S; 70°48'43.9" W), as well as a sample of Cpx-Opx quartz gabbro from the Muzzio gabbro on the southern shore of Lago Fontana (MUZ-224; 44°58'23.4" S; 71°28'10.0" W).

U-Pb Geochronology U-Pb dating was carried out using sensitive high-resolution ion microprobes (SHRIMP RG and SHRIMP II) at The Australian National University, Canberra, following the procedures of Williams (1998). Cathodo-luminescence images show that all the zircons are relatively simple with concentric zonation of the outer parts of the grains and no obviously inherited cores (Fig. 4). Analysis targets were mostl within the well-zoned rims and tips of euhedral grains and were selected so as to avoid cracks and inclusions. Data for the SHRIMP analyses were processed using SQUID (Ludwig 2001) and Isoplot/Ex (Ludwig 1999) and ages were calculated from the 206p|3_238Tj ratios corrected for the appropriate composition of common Pb based on the measured 207Pb (Table 1). Tera-Wasserburg diagrams (Fig. 5) show results for the analysed samples described

Fig. 4. Reflected light and cathodo-luminescence images of analysed zircon grains from sample SER046 (La Angostura granite). The large grain is c. 250 um in length. SHRIMP ablation spots are marked on the left-hand image, chosen to avoid cracks and inclusions.

Fig. 5. Tera-Wasserburg plots of SHRIMP zircon data for magmatic units of the Subcordilleran plutonism of northwestern Patagonia. Open ellipses are 68% confidence limits for analyses of crystal tips, uncorrected for common Pb. Age errors are reported at 95% confidence limit. Shaded ellipses are for data outside the acceptable range for definition of a single event, probably due to radiogenic Pb-loss; these were excluded from the calculation of crystallization ages. MSWD, Mean Square of Weighted Deviates.

SUBCORDILLERAN MAGMATISM OF PATAGONIA

223

Table 1. SHRIMP U-Pb zircon data Total Grain, spot SER-046 1.1 2.1 3.1 4.1 5.1 6.1 7.1 8.1 9.1 10.1 11.1 12.1 13.1 14.1 15.1 16.1 17.1

Th U Th-U (ppm) (ppm)

206pb*

(ppm)

n.i

2.1 3.1 4.1 5.1 6.1 7.1 8.1 9.1

10.1

ni.i

12.1 13.1 14.1 15.1 t!6.1

n?.i

18.1 19.1

Age (Ma)

±

0.400 0.0501 0.0009 0.0290 0.0003 0.459 0.0535 0.0012 0.0280 0.0004 0.446 0.0501 0.0009 0.0284 0.0004 0.417 0.0503 0.0008 0.0278 0.0003 0.407 0.0501 0.0010 0.0294 0.0004 0.379 0.0500 0.0006 0.0286 0.0003 0.436 0.0492 0.0010 0.0281 0.0003 0.436 0.0512 0.0010 0.0276 0.0003 0.445 0.0513 0.0011 0.0282 0.0004 0.427 0.0496 0.0010 0.0285 0.0004 0.429 0.0513 0.0010 0.0280 0.0003 1.208 0.0518 0.0012 0.0288 0.0010 0.395 0.0502 0.0009 0.0296 0.0003 0.401 0.0498 0.0010 0.0291 0.0003 0.346 0.0500 0.0010 0.0290 0.0003 0.336 0.0497 0.0008 0.0286 0.0003 0.300 0.0504 0.0006 0.0285 0.0002 Weighted mean age and 95% c.l. error

184.5 178.1 180.7 176.9 187.0 181.6 178.6 175.6 179.0 181.1 178.1 183.0 187.8 184.6 184.4 181.8 181.2 181.1

2.1 2.3 2.3 2.0 2.2 2.0 2.2 2.1 2.2 2.2 2.1 6.3 2.2 2.1 1.8 1.7 1.5 1.7

0.52 19.093 0.230 0.0572 0.0019 0.0521 0.0006 0.07 35.974 0.417 0.0502 0.0013 0.0278 0.0003 0.14 35.253 0.399 0.0508 0.0008 0.0283 0.0003 0.19 34.506 0.398 0.0513 0.0009 0.0289 0.0003 0.15 34.361 0.412 0.0510 0.0010 0.0291 0.0004 0.41 34.604 0.430 0.0530 0.0012 0.0288 0.0004 0.71 34.080 0.431 0.0555 0.0013 0.0291 0.0004 0.000056 0.17 35.270 0.452 0.0510 0.0009 0.0283 0.0004 0.000216 0.17 35.028 0.545 0.0510 0.0011 0.0285 0.0004 0.000179 0.30 34.018 0.381 0.0522 0.0012 0.0293 0.0003 0.000128 0.24 36.409 0.433 0.0515 0.0013 0.0274 0.0003 — 0.14 35.360 0.411 0.0508 0.0008 0.0282 0.0003 0.000131 0.51 35.620 0.428 0.0537 0.0010 0.0279 0.0003 0.000199 0.47 35.219 0.470 0.0534 0.0014 0.0283 0.0004 0.000431 0.86 35.495 0.456 0.0565 0.0019 0.0279 0.0004 0.000749 0.92 36.203 0.482 0.0569 0.0022 0.0274 0.0004 0.000397 1.15 36.272 0.512 0.0587 0.0025 0.0273 0.0004 0.000320 0.44 34.934 0.455 0.0532 0.0018 0.0285 0.0004 0.000339 0.38 36.135 0.467 0.0526 0.0011 0.0276 0.0004 Weighted mean age and 95% c.l. error (excluding points markedt)

327.4 176.6 180.1 183.8 184.7 182.9 185.1 179.9 181.2 186.2 174.3 179.5 177.6 179.7 177.6 174.0 173.3 181.2 175.3 181.1

4.0 2.0 2.0 2.1 2.2 2.3 2.3 2.3 2.8 2.1 2.1 2.1 2.1 2.4 2.3 2.4 2.5 2.4 2.3 2.5

0.65 35.849 0.538 0.48 34.360 0.496 0.35 33.295 0.460 0.74 34.595 0.490 0.23 33.753 0.477 0.49 33.414 0.443 0.32 34.118 0.419 0.000233 0.000308 0.28 34.070 0.482 0.000687 17.65 29.331 0.460 0.000068 0.75 34.755 0.457 0.000253 0.39 34.887 0.504 — 0.38 33.851 0.497 0.24 34.451 0.447 0.62 34.975 0.487 Weighted mean age and 95% c.l.

0.0548 0.0012 0.0277 0.0004 0.0536 0.0017 0.0290 0.0004 0.0527 0.0015 0.0299 0.0004 0.0556 0.0016 0.0287 0.0004 0.0517 0.0015 0.0296 0.0004 0.0538 0.0016 0.0298 0.0004 0.0524 0.0017 0.0292 0.0004 0.0520 0.0019 0.0293 0.0004 0.1905 0.0658 0.0281 0.0030 0.0557 0.0014 0.0286 0.0004 0.0528 0.0027 0.0286 0.0004 0.0528 0.0026 0.0294 0.0004 0.0517 0.0012 0.0290 0.0004 0.0547 0.0014 0.0284 0.0004 error (excluding points markedt)

176.2 2.6 184.1 2.7 190.1 2.6 182.4 2.6 187.8 2.7 189.2 2.5 185.6 2.3 2.6 186.0 178.5 18.5 181.5 2.4 181.5 2.7 187.0 2.8 2.4 184.0 180.6 2.5 184.9 2.3

0.67 0.48 0.000055 <0.01 0.000219 0.29 — 0.51 0.000045 0.09

0.0548 0.0533 0.0486 0.0520 0.0539 0.0505

171.5 172.7 187.1 181.2 185.3 182.1

598 301 731 579 466 1234 469 499 360 438 462 321 596 582 409 455 1150

229 106 313 185 185 634 160 206 116 148 189 94 229 231 144 183 562

0.38 0.35 0.43 0.32 0.40 0.51 0.34 0.41 0.32 0.34 0.41 0.29 0.39 0.40 0.35 0.40 0.49

14.9 7.3 17.9 13.9 11.8 30.3 11.3 11.9 8.7 10.7 11.1 8.0 15.1 14.5 10.2 11.2 28.2

0.000318 0.000313 0.000074 0.000144 0.000197 0.000089

148 339 420 361 252 211 195 376 288 458 430 516 274 167 380 335 287 462 492

38 494 356 652 300 244 181 268 326 412 358 646 186 170 312 402 364 388 485

0.26 1.46 0.85 1.80 1.19 1.16 0.93 0.71 1.13 0.90 0.83 1.25 0.68 1.02 0.82 1.20 1.27 0.84 0.99

6.7 8.1 10.2 9.0 6.3 5.2 4.9 9.2 7.1 11.6 10.1 12.5 6.6 4.1 9.2 7.9 6.8 11.3 11.7

— — 0.000127 0.000159 0.000039

185 182 124 100 103 146 195 98 93 134 130 142 238 148

90 165 79 57 64 130 148 55 59 111 117 116 243 81

0.49 0.91 0.63 0.57 0.62 0.89 0.76 0.56 0.64 0.83 0.90 0.82 1.02 0.55

4.4 4.6 3.2 2.5 2.6 3.7 4.9 2.5 2.7 3.3 3.2 3.6 5.9 3.6

95 203 216 160 103 233

66 232 199 170 71 193

0.70 1.14 0.92 1.07 0.69 0.83

2.2 4.7 5.5 3.9 2.6 5.7

LEL-052

Radiogenic 206Fb_ 238U

206pb_ 238U

204pb/ 206pb

0.000091 — 0.000226 0.000181 0.000174 0.000138 0.000000 0.000109 0.000058

%

23*U206pb

0.05 0.49 0.05 0.08 0.03 0.04 <0.01 0.20 0.21 <0.01 0.21 0.26 0.05 0.01 0.03 <0.01 0.08

34.431 35.531 35.157 35.904 33.957 34.983 35.623 36.131 35.442 35.102 35.620 34.643 33.811 34.417 34.461 34.965 35.044

f206

207Fb_ ±

206

Pb

±

±

ALE-055

n.i

2.1 3.1 4.1 5.1 6.1 7.1 8.1 t9.1 10.1 11.1 12.1 13.1 14.1

0.000127 0.000331 0.000228 0.000022

JSM-058

n.i

t2.1 3.1 4.1 5.1 6.1

0.000434

36.833 36.655 34.002 34.983 34.119 34.876

0.532 0.455 0.415 0.446 0.477 0.418

0.0026 0.0012 0.0012 0.0012 0.0015 0.0010

0.0270 0.0272 0.0295 0.0285 0.0292 0.0286

0.0004 0.0003 0.0004 0.0004 0.0004 0.0003

2.5 2.1 2.3 2.3 2.6 2.2

224

CW.RAPELAEr.4L.

Table 1. (continued) Radiogenic

Total Grain, spot 7.1 10.1

m.i 12.1 13.1 14.1 15.1 16.1 17.1

Th U Th-U (ppm) (ppm) 88 68 350 161 129 314 240 467 302

65 38 568 103 130 370 257 573 360

0.74 0.55 1.63 0.64 1.00 1.18 1.07 1.22 1.19

206pb*

(ppm) 2.2 1.6 7.9 3.9 3.2 7.7 5.9 11.7 7.4

204pb/ 206pb

f206 %

238lJ_ 206pb

±

20?Pb_ 206pb

±

206Fb_ 238U

±

0.000199 0.21 35.000 0.512 0.0514 0.0016 0.0285 0.0004 0.001140 1.27 35.664 0.741 0.0597 0.0021 0.0277 0.0006 0.000081 <0.01 38.151 0.545 0.0493 0.0009 0.0262 0.0004 — 0.50 35.905 0.453 0.0536 0.0022 0.0277 0.0004 0.54 35.154 0.474 0.0540 0.0014 0.0283 0.0004 0.10 35.102 0.408 0.0505 0.0009 0.0285 0.0003 0.23 35.173 0.424 0.0515 0.0015 0.0284 0.0003 <0.01 34.308 0.416 0.0495 0.0008 0.0292 0.0004 0.000074 0.19 35.204 0.474 0.0512 0.0010 0.0284 0.0004 Weighted mean age and 95% c.l. error (excluding points marked f )

Age (Ma) 206pb_ 238U

±

181.2 176.0 166.8 176.2 179.8 180.9 180.3 185.3 180.2 181.5

2.7 3.7 2.4 2.3 2.4 2.1 2.2 2.2 2.4 2.3

Notes: Uncertainties are given at the la level, except for the means. The error in FCI reference zircon calibration was 0.61% & 0.30% for the two analytical sessions (included in the errors on the means for comparing data from different mounts and methods). f206% denotes the percentage of 206Pb that is common Pb. Correction for common Pb was made using the measured 238u_206Pb and 207Pb/206Pb ratios following Tera & Wasserburg (1972), as outlined in Williams (1998). c.l., confidence limits. Pb* is radiogenic Pb.

samples of the Subcordilleran belt (including the three granites and three gabbros from Table 2) and equivalent numbers from the other units (see figure captions for sources). The major element compositions of most of the plutonic rocks show general similarities in the normative An-Ab-Or diagram (Fig. 3). The Batholith of Central Patagonia and the Deseado Monzonite Suite lack the conspicuous An-rich facies (typically gabbros and diorites) typical of the Late Jurassic to Pliocene Patagonian batholith, but the association with basic plutonic rocks in the SCB suggests a very close affinity with the Patagonian batholith. The abundance of gabbros and basic rocks in the SCB (assuming the Tecka Formation gabbros to be coeval with the granites), is most probably underestimated in Figure 3, as they form conspicuous outcrops in the southern part of the belt. The sample from La Angostura (SER-046) was considered too weathered for whole-rock analysis, but published data for this area are included in the geochemical plots. New chemical analyses of granitic rocks from the main facies in the Leleque, Aleusco and Jose de Geochemistry San Martin areas, as well as representative The U-Pb results clearly demonstrate that the samples of the abundant gabbros from the granites of the SCB are of Early Jurassic age, southern part of the belt, are shown in Table 2. essentially contemporaneous with the Vj The granitic rocks are all subalkaline metalumiepisode of rhyolite volcanism in the North nous granodiorite/tonalites/ monzogranites with Patagonian Massif (188-178 Ma). Their an ASI = 0.91-0.96 (ASI: Alumina Saturation geochemistry will be considered in the overall Index = molecular Al2O3/CaO+Na2O+K2O), context of Late Triassic to Early Jurassic SiO2 61.46-69.66%, Na2O+K2O 6.01-7.28%, magmatism in the region and the subsequent FeOt+MgO 3.58-8.25%, REE patterns with plutonism of the Andean Patagonian batholith, [La/Yb]N = 4.7-10.5 and a weak negative Eu an archetypical example of a continental margin anomaly (Eu/Eu* - 0.67-0.78) (Table 2, Figs 6, subduction-related granite batholith (Fig. 1). 7c). The overall subalkaline geochemistry of The geochemical database consists of 35 these granites is typically that of I-type granites

above from La Angostura, Leleque, Aleusco and Jose de San Martin. Seventeen grains from the granite at La Angostura were analysed and gave a mean age of 181 ± 2 Ma (MSWD - 2.0). Nineteen grains were analysed from the Leleque granodiorite, of which one is clearly inherited (c. 330 Ma). Even discounting this, the data spread outside the expected analytical error (MSWD = 3.2) and the four youngest 206 Pb-238U ages were disregarded as probably suffering from slight Pb-loss, apparently at about 174 Ma (perhaps as a result of V2 volcanic heating). The remaining 14 define an Early Jurassic age of 181 ± 3 Ma. From the 14 analysed grains from the Aleusco granodiorite, one suffered from an unusually high analytical uncertainty and one appears to have suffered Pb-loss as in LEL-052; the remaining 12 grains define an Early Jurassic age of 185 ± 2 Ma. Finally, 12 grains of a quartz monzodiorite from the Jose de San Martin body also gave a Early Jurassic age of 182 ± 2 Ma, discounting three grains in which Pb-loss is apparent.

225

SUBCORDILLERAN MAGMATISM OF PATAGONIA Table 2. Chemical analyses of representative samples from the Subcordilleran plutonic belt ALE-055 Gd

JSM-058 QM

MUZ-224 QGa

QUI-225 QGa

QUI-227 Ga

Major oxides (wt%) 69.66 Si02 Ti02 0.40 14.63 A1203 1.35 Fe2O3 FeO 1.38 0.09 MnO MgO 0.98 2.76 CaO 4.24 Na2O K2O 2.97 0.11 P205 0.58 H2O+ H200.08 Total 99.23

67.49 0.54 14.86 0.63 2.95 0.07 1.53 3.14 3.59 3.69 0.12 1.21 0.13 99.95

61.46 0.88 16.11 2.11 3.83 0.11 2.52 4.97 3.74 2.27 0.18 1.19 0.20 99.57

54.38 0.63 15.98 0.39 7.82 0.15 6.03 8.82 1.86 1.14 0.10 1.59 0.28 99.17

54.65 0.62 15.92 0.10 8.03 0.15 6.77 9.28 1.79 1.24 0.11 0.19 0.28 99.13

48.35 0.23 19.27 0.72 5.19 0.11 11.45 11.18 1.17 0.35 0.03 1.00 0.27 99.32

Trace elements (ppm) Cs 4.0 Rb 109 Sr 262 Ba 651 La 31.0 Ce 58.5 Pr 6.73 21.4 Nd Sm 3.99 0.94 Eu Gd 3.27 Tb 0.53 Dy 2.93 Ho 0.60 Er 1.86 Tm 0.30 Yb 1.97 0.32 Lu Sc 5.0 U 2.38 Th 13.8 Y 18.4 Nb 8.7 Zr 151 Hf 4.3 Ta 0.80 Ga 17 2.20 Ga/Al 0.72 Rb/Zr 2.65 FeOt/MgO 21.00 Ba/La

4.7 152 203 573 29.9 57.3 6.74 22.7 4.85 1.01 4.26 0.74 4.31 0.87 2.65 0.41 2.56 0.41 9.0 2.83 20.0 25.7 7.4 232 6.8 0.69 17 2.16 0.66 2.30 19.16

3.8 93 297 472 21.6 45.1 5.89 22.1 5.39 1.36 5.03 0.90 5.26 1.09 3.26 0.50 3.08 0.49 14.0 1.62 9.8 31.2 7.0 185 5.4 0.56 19 2.23 0.50 2.27 21.85

2.3 39 115 268 16.5 34.5 4.06 16.4 3.84 0.96 3.76 0.67 4.18 0.88 2.77 0.41 2.63 0.40 34.0 1.57 6.3 23.1 5.9 125 3.7 0.39 15 1.77 0.31 1.35 16.20

1.6 49 114 249 16.3 34.0 3.98 16.4 3.70 0.96 3.76 0.67 4.16 0.89 2.77 0.41 2.59 0.39 34.0 1.55 6.1 23.5 6.1 127 3.5 0.37 16 1.90 0.39 1.20 15.40

1.5 9 101 77 4.3 8.9 1.09 4.5 1.16 0.47 1.28 0.24 1.57 0.34 1.07 0.16 1.05 0.15 21.0 0.30 1.4 8.9 2.1 31 1.0 0.10 12 1.18 0.29 1.51 17.90

Sample no. Lithology

LEL-052 Gd

Gd = granodiorite; QM = quartz monzonite; QGa = quartz gabbro; Ga = gabbro. Major oxides and trace elements were determined respectively by ICP and ICP-MS at ACTLABS (Canada). Fe++ determined volumetrically at Centre de Investigaciones Geologicas, La Plata.

produced by melting of igneous sources (Chappell & White 1992). Furthermore, the Ga/Al (1000*Ga/Al - 2.2), Rb/Zr (0.5-0.7) and FeOt/MgO (2.4-2.9), Zr/Nb (>17) and Ba/La (c. 20) ratios are characteristic of metaluminous

magmas emplaced in convergent margins rather than intraplate settings (Pearce et al. 1984; Whalen et al 1987; Eby 1990). Major and trace element Marker and multi-element diagrams show that data for the granitic rocks of the

226

C.W. RAPELA ETAL.

Fig. 6. Marker plots for Triassic-Jurassic I-type granites and volcanic rocks from Patagonia. The line dividing alkaline and subalkaline fields is from Irvine & Baragar (1971) while the fields for high, medium and low Fe are after Arculus (2003). Data sources for the Triassic and Jurassic volcanic rocks: Pankhurst & Rapela (1995); Rapela (2001) and unpublished data from a British Antarctic Survey-Argentine joint expedition (P. T. Leat, pers. comm. 2003). Data sources for the I-type granites are the same as reported in Figure 3.

SUBCORDILLERAN MAGMATISM OF PATAGONIA

Fig. 7. (a) Trace element abundances in granitic and basic rocks of the Subcordilleran belt normalized to MORE (Pearce 1983). (b) Trace element abundances in granitic rocks of the Subcordilleran belt normalized to ocean ridge granite (ORG, Pearce et al 1984). The field of the Patagonian batholith has been constructed from chemical data for hornblende-biotite granodiorites and tonalites (61-69% SiO2) reported by Ghiara et al (1997), Pankhurst et al (1999) and unpublished data of C. W. Rapela at 41° S. (c) Chondrite-normalized REE patterns of representative samples from the Subcordilleran plutonic belt.

227

SCB all plot within the fields defined by a large geochemical dataset for the Patagonian batholith (Figs 6, 7b). The main fades of the different granitic complexes of the SCB show enrichment of Rb, Ba and Th relative to Nb, Ta, REE, Hf and Zr (Fig. 7b), a characteristic of magmatic arc rocks (Thompson et al 1984). More specifically, they show the same patterns as the hornblende-biotite granodiorites and tonalites of the Patagonian batholith. The arc geochemical signature of the granitic rocks is shared by the abundant subalkaline gabbros of the southern part of SCB, which show arc-tholeiite affinities when normalized against MORB (Fig. 7a) and are inferred to be subduction-related basic rocks (Page & Page 1999). Remarkably, the shape of the MORBnormalized patterns of the basic rocks mimic those of the granites, but the latter are more enriched in highly incompatible elements (Fig. 7a). The plagioclase-rich olivine gabbro QUI227 has a REE pattern with a positive Eu anomaly (Fig. 7c), suggesting that some of the basic rocks are cumulate-rich facies, consistent with the widespread cumulate fabric of the association and the occurrence of minor anorthosites and ultrabasic bodies reported by Page & Page (1999). It is, therefore, possible that the granitic rocks of the SCB were derived by fractional crystallization from the basic coeval magmas, at least in part. The gabbros from the Quichaura and Muzzio bodies (QUI225 and MUZ-224) are indistinguishable in all geochemical respects, demonstrating without doubt that the latter is a member of the Tecka Formation and that the Subcordilleran plutonic belt extends over at least 300 km from La Angostura to Lago Fontana. Compared to the SCB granites, the dominant facies of the Late Triassic Batholith of Central Patagonia and the Deseado Monzonite Suite (62-68% SiO2) are enriched in P2O5 and Sr, with relatively high Sr/Y ratios and low FeOt /MgO ratios (Fig. 6). None of the Triassic and Jurassic plutonic rocks of Patagonia have geochemical characteristics of the TTD (trondhjemite-tonalite-dacite) 'adakite' suite, formed by partial melting of subducted ocean crust, according to Drummond et al (1996). Such compositions are recognized in Andean Miocene-Holocene dacites at 48-54° S (Kay et al 1993; Stern & Kilian 1996) and c. 70 Ma-old trondhjemites at 53° S in the Patagonian batholith (Bruce 1988). The Deseado Monzonite Suite shows some geochemical affinities with the TTD association, such as a tendency to high Sr/Y (Fig. 6) and depletion of HREE, but high LIL (large-ion

C.W.RAPELAETAL.

228

lithophile) element concentrations and enriched Sr and Nd isotopic signature rule out a depleted N-MORB source (Rapela & Pankhurst 1996). The Early Jurassic Vi group of silicic volcanic rocks, which crop out in Patagonia east of the Andes and at the same latitudes as the SCB (Fig. 1), show geochemical signatures very different from the majority of the I-type granites. These high-K rhyolites and ignimbrites are enriched in alkalis, P2O5 and, notably, in Zr and Nb, particularly in the intermediate SiO2 range (Fig. 6). In some cases, such as on the Atlantic coast at 45° S (Peninsula Camarones), they reach peralkaline compositions, with high Zr (425-600 ppm, Fig. 6) and TiO2 (0.45-0.95%) (Pankhurst & Rapela 1995). Coeval Vi rhyolites in the Antarctic Peninsula show similar geochemical characteristics to those of Patagonia (Riley et al 2001). However, in northern Patagonia, the Vx group was preceded by the Late Triassic hornblende-biotite ignimbrite of Los Menucos, which shows geochemical trends akin to those of the contemporaneous I-type units of the Batholith of Central Patagonia (Fig. 6). The easternmost rhyolites and ignimbrites of the Middle Jurassic V2 group, located near the Atlantic coast at 47°45' S (Puerto Deseado), are also enriched in Nb and Zr (Fig. 6). In contrast, the remainder of the V2 group and the younger Middle-Late Jurassic V3 group, located in Andean and pre-Andean regions, exhibit major and trace element abundances and trends similar to those observed in the I-type granites (Fig. 6). It is worth noting that Riley et al (2001) concluded that all three rhyolite groups

inherited some subduction-related trace element characteristics from their source, although less so in Y! than in V2 and V3.

Nd and Sr isotope data Nd and Sr isotope ratios were determined on four SCB granite samples (Table 3). In Figure 8 their initial isotope compositions are compared with those of other Mesozoic-Cenozoic I-type granites, and the extensive Triassic and Jurassic volcanic rocks of Patagonia for which precise U-Pb, Ar/Ar and Rb-Sr dating have been reported (Pankhurst & Rapela 1995; Feraud et al 1999; Pankhurst et al 2000). Data for the SCB granites plot mostly within the field of the Late Jurassic-Miocene I-type granites of the Patagonian batholith, as do those of the Deseado Monzonite Suite and the Batholith of Central Patagonia (Fig. 8a). Samples LEL-052 and ALE-055 have E Ndt values indistinguishable from those of the Deseado suite, which range from -0.3 to -2.5. The quartz monzodiorite of Jose de San Martin (JSM-058) is the most primitive of all the Jurassic and Triassic I-type granites, with a depleted-mantle signature comparable to that of Mid-Cretaceous granites in the Patagonian batholith (Pankhurst et al 1999). A microdiorite enclave in the monzodiorite (JSM-059) gives the same Nd and Sr isotopic signature, confirming the same primitive source. Late Jurassic tonalites, granodiorites and quartz monzodiorites of c. 150 Ma that can be considered the earliest event of the Patagonian batholith (Martin et al 2001; Pankhurst et al

Table 3. Isotope data for samples from the Subcordilleran plutonic belt Sample Rb (ppm) Sr (ppm) Rb/86Sr 87 Sr/86Sr (87Sr/86Sr)0 Sm (ppm) Nd (ppm) 87

147Sm_144Nd 143Nd/144Nd (143Nd/144Nd)()

e

Ndt

TCHUR TDM*

LEL-052 Gd

ALE-055 Gd

JSM-058 QM

JSM-059 X

115.5 253.5 1.3183 0.708605 0.705193 3.77 21.63 0.1052 0.512445 0.512320 -1.6 323 1029

154.0 175.5 2.5398 0.711854 0.705282 4.94 25.73 0.1160 0.512478 0.512340 -1.2 303 995

86.0 243.0 1.0239 0.707387 0.704737 5.80 25.79 0.1360 0.512639 0.512477 1.4 -3 770

55.0 277.5 0.5734 0.706431 0.704947 4.65 21.18 0.1329 0.512631 0.512473 1.3 17 777

Gd = granodiorite; QM = quartz monzonite; X = microdiorite enclave. Analytical methods as in Pankhurst & Rapela (1995): Rb and Sr by XRF at British Geological Survey (± 0.5%, lo); 87Sr/86Sr measured on MAT 262 (± 0.01%, la); Sm and Nd by MS Isotope dilution at NIGL (± 0.1%, lo); 143Nd/144Nd measured on MAT 262 (± 0.005%, la); TDM* for crustal source after DePaolo et al (1991).

SUBCORDILLERAN MAGMATISM OF PATAGONIA

Fig.8. Variation of Nd( versus initial 87Sr/86Sr for the Late Triassic and Jurassic magmatic units of Patagonia. Data sources for I-type granites: Rapela et al (1992), Rapela & Pankhurst (1996), Pankhurst et al (1999) and Pankhurst & Herve, unpublished data. For volcanic rocks: Pankhurst & Rapela (1995) and Pankhurst, unpublished data. The fields for the upper and lower crust in (c) are from the Chilean accretionary prism and the North Patagonian Massif Vj rhyolite group, respectively (Pankhurst et al. 1999). The MORE (mid-ocean ridge basalts) field is from Zindler & Hart (1986), while SVZ (mafic Southern Volcanic Zone) is from Gorring & Kay (2001 and references therein).

2000; Suarez & De la Cruz 2001) plot in the enriched-source sector of the Nd-Sr isotopic diagram. Data for Late Jurassic satellite bodies located east of the batholith, such as the Sobral

229

tonalite (Pankhurst et al 2000; Fig. 8a), confirm that the earliest magma batches of I-type magmas are among the most isotopicallyevolved of the batholith, regardless of geographical location. In contrast, the Sm-Nd relationships of granitoids from the Patagonian batholith at 44-46° S indicate source compositions that change from slightly LIL-enriched for the Early Cretaceous rocks to significantly depleted for the Early Miocene rocks, the latter in turn very similar to those of the Tertiary to Recent mafic strato-volcanoes of the Southern Volcanic Zone of the Andes (Pankhurst et al 1999). Bruce et al (1991) also showed a broad regional trend of decreasing initial 87Sr/86Sr with time, with some values as high as 0.707 for Late Jurassic-Early Cretaceous granites at 48° S. None of the extensive Triassic and Jurassic volcanic rocks of Patagonia seem to be derived from long-term LIL-depleted sources such as those of the SCB granites and the Mid Cretaceous-Tertiary granites of the Patagonian batholith (Fig. 8b). In particular, many of the earliest rhyolitic rocks of northern Patagonia have £Ndt values that fall significantly below the field of I-type batholiths, with low £ Ndt at relatively low 87Sr/86Sr (Fig. 8b). In the case of the Vi group rhyolites, this signature has been ascribed to a hypothetical Grenville-age lower crust (Pankhurst & Rapela 1995; Riley et al 2001), implying a completely different genesis to that of the coevally emplaced granites of the SCB. However, some Early Cretaceous granites of the Patagonian batholith plot on a steep trend extending towards the compositions of these V! rhyolites, suggesting contamination of mantlederived magmas with lower crustal melts (Pankhurst et al 1999). Simple mixing arrays between melts derived from the lithospheric mantle wedge and either upper or lower crustal contaminants could explain the isotopic trends observed in I-type granites (Fig. 8c). The change towards depletedsource melts with time in the Patagonian batholith has been attributed to a progressive decrease in crustal contamination due to 'magmatic inflation' during the growth of the body, whereby younger magmas emplaced along the axis of the batholith were physically isolated from contamination by radiogenic wall rocks by earlier intrusions (Bruce et al 1991). However, if the isotopic composition of the main source of all I-type magmas were always similar to (or more primitive than) that of the Tertiary granites (eNdt values between -4 and -6, Pankhurst et al 1999), it is difficult to understand why depleted signatures are never observed in Triassic and Jurassic granites nor

230

C.W.RAPELAETAL.

any of the Andean volcanic rocks (V3 group, Fig. 8b). Instead, it is suggested here that melting occurred in progressively more LILdepleted sources with time and that the relatively primitive signature of the abundant Mid-Cretaceous-Tertiary granites is indicative of melting in the lithospheric mantle underlying the Patagonian batholith (Fig. 8c). The evolve composition of Late Triassic to Early Cretaceous granites could then be explained by variable amounts of upper and lower crustal contamination of melts derived from an effectively undepleted lithospheric source, such as that represented by the Deseado Monzonite Suite or the granites of the SCB (Figs 8a,c).

The Subcordilleran belt: remnant of an Early Jurassic magmatic arc The U-Pb SHRIMP crystallization ages of typical metaluminous granodiorites and quartz monzodiorites from the SCB between 42° S and 44° S are remarkably consistent, much more precise than the Rb-Sr and K-Ar ages reported

for similar rocks by Gordon & Ort (1993) and Haller et al. (1999), and compatible with the Ar/Ar age for the subalkaline gabbros given by Page & Page (1999). The age interval of 187-178 Ma, allowing for analytical error, falls in the Pliensbachian and Toarcian stages of the GSA and IUGS-ICS time-scales (Palmer & Geissman 1999; Remane etal. 2000). Altogether the data presented here for four separate granite bodies in the Subcordilleran belt indicate that I-type magmatism of short apparent duration was emplaced into the 300 km long (at least) Liassic sedimentary basin within a very few million years of its deposition. It cannot be traced south of 44°30' S, although Early Jurassic arc rocks could be buried by Cretaceous sediments in the San Jorge Basin or by thick Middle Jurassic ignimbrites and Tertiary plateau basalts in the Deseado Massif. The 700 m thick sequence of fossil-bearing Liassic marine sediments and volcaniclastic rocks found in deep boreholes of the San Jorge Basin at 46° S (Uliana & Legarreta 1999) suggest that this possibility would be worth testing in future research. Nevertheless, the

Fig. 9. Histograms of ages for the Mesozoic and Cenozoic I-type granites of Patagonia. Data sources: Patagonian batholith between 45° and 53° S (Bruce et al. 1991; Pankhurst et al 1999; Suarez & De la Cruz 2001); Subcordilleran belt (Lizuain 1981; Gordon & Ort 1993; Haller et al. 1999; this paper); Deseado Monzonite Suite (Pankhurst et al 1993); Batholith of Central Patagonia (Rapela et al 1992). Cumulative probability curve and age groups for the Jurassic rhyolite province of Patagonia are from Pankhurst et al. (2000). Inferred ages for the Karoo and Tristan mantle plumes are from Duncan et al (1997) and Turner et al (1994), respectively.

SUBCORDILLERAN MAGMATISM OF PATAGONIA

timing and disposition of SCB magmatism show that this was a distinct event from both the Triassic 'oblique' batholiths and the Patagonian batholith (Fig. 9). Early Jurassic ages have not been reported from any part of the Patagonian batholith. On the other hand, the geochemistry and isotopic signature of the SCB I-type granites and coeval basic rocks are typical of subductionrelated cordilleran-type belts in convergent continental margins and cannot be distinguished from those of the Andean batholiths emplaced in the Pacific margin of South America after Gondwana break-up (Figs 3, 6-8). Furthermore, emplacement in a continental-margin marine basin, filled with littoral sediments and coeval volcanic and volcaniclastic rocks, suggests an intra-arc environment (Lizuain 1999; Page & Page 1999).

I-type magmatism, rhyolite provinces and Gondwana break-up The I-type batholiths and plutons of Patagonia record a complex Mesozoic history of convergence episodes that preceded, were coeval with, and outlasted the major stages of supercontinent rifting. The relative abundance of mafic igneous rocks in the SCB and the Patagonian batholith, associated the final stage of rifting and the active opening of the South Atlantic Ocean, is perhaps the most important lithological difference from the Triassic I-type suites emplaced during early supercontinental convergence and rifting. The geometrical relationships between the Karoo-Ferrar provinces, the proto-Pacific subduction margin and the supercontinent break-up lines (Fig. 10) suggest that these are all interrelated and that Karoo magmatism initiated separation between Eastern and Western Gondwana (Cox 1992, and references therein). The data presented here for the SCB granites demonstrate that initiation of Andeantype subduction in Patagonia was coincident with intraplate magmatism, including the mafic magmatism of the Ferrar and Karoo mafic igneous provinces at 184-179 Ma, with a climax at 183 Ma (Encarnacion et al. 1996; Duncan et al 1997;), and the contemporaneous V\ episode of rhyolitic magmatism in northeastern Patagonia (Fig. 9). Recently, Riley et al. (2004) have shown, also using SHRIMP U-Pb zircon dating, that Karoo rhyolites in the Lebombo area of eastern South Africa were erupted between 182 ± 3 Ma and 180 ± 2 Ma, indistinguishable from the age of the SCB granites. Thus, although

231

Elliot & Fleming (2000) argued for a single source for the mafic lavas, centred on a triple junction in the proto-Weddell Sea, felsic magmatism associated with this event extended far away from the site of the Karoo plume, reaching the proto-Pacific margin of the supercontinent, as suggested by Pankhurst et al. (2000). Middle Jurassic magmatic activity shifted towards southern Patagonia and the Antarctic Peninsula, developing the extensive rhyolite sequences of the Chon Aike Formation (V2 event, Pankhurst et al. 2000) (Figs 1, 9). In Tierra del Fuego, the peraluminous Darwin granite suite is thought to be a correlative of the V2 rhyolites (164.1 ± 1.7 Ma, Mukasa & Dalziel 1996). This Middle Jurassic acid event (c. 172-162 Ma: e Ndt = -1.8 to -3.9; initial 87 Sr/86Sr - 0.7064-0.7169) is now the only one for which no I-type cordilleran counterpart has been reported (Fig. 9), although further systematic and precise geochronological work will be necessary to confirm such a temporal hiatus in the subduction-related cordilleran magmatism along the southern and austral Andes. From Late Jurassic to Tertiary times, the main locus of I-type magmatic activity defines the axis of the modern Andean Cordillera (Rapela & Kay 1988). The plutonic component of this activity is represented by the Patagonian batholith, which ranges in age from c. 155 Ma to 5 Ma. The oldest granites (155-143 Ma, Late Jurassic) are restricted to the eastern side of the batholith and satellite plutons east of the Patagonian batholith (Bruce et al 1991; Martin et al 2001; Suarez & De la Cruz 2001, Rolando et al 2002). Late Jurassic K-Ar ages have also been reported from the northeastern sector of the Patagonian batholith (39°42'-42° S) (Gonzalez Diaz 1982; Rapela & Kay 1988), whereas ages in this range have not been identified along the western side of the batholith. Late Jurassic volcanic rocks also occur along the length of the Patagonian Andes, as well as in Argentine sectors of the North Patagonian and Deseado massifs (Pankhurst et al 2000; Suarez & De la Cruz 2001 and references therein). Thick rhyolite sequences of the V3 group (c. 157-153 Ma and younger) in the Andean sector of southern Patagonia are considered to be active-margin volcanic products associated with the Late Jurassic granites (Pankhurst et al 2000). On the oceanic side of the Andes, this volcanism continued locally into Cretaceous times (e.g. Suarez & De la Cruz 19970, b\ Pankhurst et al 20030).

232

C.W.RAPELAETAL.

Fig. 10. Schematic Early Jurassic palaeogeographical reconstruction of southwestern Gondwana (modified after the Early Jurassic base map of Storey et al. 1992). A tight fit is obtained by a palinspastic restoration of southern South America and the Antarctic Peninsula as proposed by MacDonald et al (1998), involving 50% closing of the largest southernmost Mesozoic basins (Austral, San Jorge and Colorado, Fig. 1), reducing by 20% the size of the Antarctic Peninsula to remove approximately Late Mesozoic and Tertiary crustal growth, and allowing for moderate dextral displacement of about 140 km of large crustal blocks along major fault systems bounding the North Patagonian and the Deseado massifs. Partial restoration of the Cretaceous-Cenozoic oroclinal curvature in the southern tip of South America and the Antarctic Peninsula has also been performed (e.g. Kraemer 2003). Triassic and Early Jurassic accretionary prism after Thomson & Herve (2002). Location and timing of the Jurassic magmatic provinces are after Pankhurst et al. (2000), while heavy broken lines that show the sites of the initial areas of the Indian Ocean (1) and the future South Atlantic Ocean (2) are from Cox (1992). Crustal blocks: EWM, Ellsworth-Whitmore mountains; SG, South Georgia; F/M, Falkland/Malvinas microplate; TI, Thurston Island; MBL, Marie Byrd Land; NZ, New Zealand.

From the above-described space-time relationships of the I-type belts it is clear that during the Early Jurassic to Early Cretaceous time interval there was a significant ocean-ward migration of the magmatic arc (of at least 150 km at the latitude of Jose de San Martin, Fig. 1). In the northern Antarctic Peninsula and Thurston Island (Fig. 10), Jurassic subductionrelated magmatism also migrated from the eastern margin towards the western margin, with emplacement of granitic plutons closer to the trench (Pankhurst 1982). Thus, ocean-ward migration of the arc during the Jurassic was

probably a continental-scale characteristic of the proto-Pacific margin of southwestern Gondwana. This westward shift of the cordilleran I-type belts ceased after the opening of the South Atlantic, the distribution of the Early Cretaceous to Miocene plutons of the Patagonian batholith suggesting a quasi-stationary position. However, the NNW-SSE orientation of the SCB compared with the N-S trend of the North Patagonian cordillera (Fig. 1) suggests that the ocean-ward shift of the arc could be described better as a clockwise migration of the axis in

SUBCORDILLERAN MAGMATISM OF PATAGONIA

northern Patagonia. The regular decrease of Ar/Ar ages in the rhyolites of Patagonia from the ENE (c. 187 Ma) to the WSW (c. 144 Ma) along about 650 km (Feraud et al 1999) might be explained also by back-arc crustal extension behind a clockwise-migrating arc. Palinspastic restoration of the Patagonian orocline at 56° S shows the opening of a 230 km wide oceanic back-arc basin at this latitude, also indicating Early-Mid-Jurassic to Early Cretaceous clockwise rotation of the magmatic arc (Dalziel 1981; Kraemer 2003 and references therein), a conclusion remarkably coherent with the results of this study. The NW-SE strike of the Batholith of Central Patagonia (Fig. 1) could also be taken as suggesting that migration of the magmatic arc may have started during the Late Triassic (Fig. 10), although the present spatial relationships could be, at least in part, an artefact resulting from crustal extension and block displacements away from the Karoo mantle plume (Fig. 10). In any case, the development of the SCB during the Early Jurassic involved an important rearrangement in the geometry of plate convergence in Patagonia, that evolved from oblique belts dominated by intermediate and acid units that characterized the later Gondwana stage, towards the north-south Andean style, with abundant basic rocks.

2.

3.

Early Jurassic palaeogeography and tectonics The apparent clockwise rotation with time of the magmatic axis represented by the progression of the elongated Batholith of Central Patagonia, the Subcordilleran plutonic belt and the Andean Patagonian batholith, from Late Triassic to Mid-Jurassic, is a major geological feature whose causes are not yet well understood. Assuming that they were all formed parallel to, and at a certain distance from, the trench in an active subducting margin, they suggest that the margin of southwest Gondwana was displaced in the same way during this period. There are a number of major tectonic factors that could be related to such a rearrangement of the margin. 1.

Differential rollback of the subduction zone, greater in E-W amplitude towards the south. This could have occurred episodically, as there are lulls in plutonic activity, or at least periods without magma-generating subduction, between that of the three

4.

233

batholiths. The 'space' left by the trench stepping out in this way would be filled by accreted or displaced material or by intracontinental extension. Asymmetric ('scissor-like') extension of the Patagonia continental crust increasing to the south. Generation of the Jurassic acid volcanic province in Patagonia suggests a generally extensional environment. Immediately east of the cordillera, the Jurassic Rocas Verdes Basin developed as a marginal basin with quasi-oceanic floor and a parautochthonous terrane to the west (with continental crust topped by a magmatic arc) (Dalziel 1981 and references therein), and closed back in a more westerly position in Cretaceous times (Kraemer 2003 and references therein). The existence of a second marginal basin, today represented by the middle Jurassic mafic blueschists of the Diego de Almagro Complex (Herve & Fanning 2003) could enhance this mechanism further. Accretion of tectonostratigraphic terranes along the southern part of the margin, resulting in the subduction zone jumping progressively westwards. The Madre de Dios accretionary complex (Fig. 10) has been considered as exotic to the Patagonian margin (Forsythe 1982), but palaeontological evidence (Ling et al 1987) and detrital zircon dating (Herve et al 2003) show that the rocks of this complex are mostly of Permian age and that sandstones in the upper part (Duque de York complex) are probably derived from the same cratonic sources as the Eastern Andes Metamorphic Complex, so that accretion must have occurred before rotation of the magmatic axis. The Chonos accretionary complex, like the Trinity Peninsula and LeMay groups of the Antarctic Peninsula (Fig. 10), contain both Permian and Triassic detrital zircons (Herve et al 2003; Millar & Pankhurst, unpublished data), but their times of accretion are not well constrained. It should be noted that continuous development of an accretionary prism, growing from east to west across Patagonia (e.g. Forsythe 1981), would conflict with the proposition that Patagonia has a Grenvilleage deep crust (Pankhurst et al 1994). Displacement of terranes along the continental margin during the periods without magmatic activity could also occur by strike-slip continental margin conditions. Left-lateral strike-slip movement of blocks along the southwestern margin of

234

C.W. RAPELA ETAL.

Gondwana-South America has been well identified for post-Early Cretaceous times (Rapalini et al 2001; Olivares et al 2003), is still active along the Magellan Fault and may also have been active before. In the Chonos region, the northward late Cenozoic movement of the Chiloe block allowed the formation of the extensional Traiguen Basin in a period of trench parallel subduction with no contemporaneous arc magmatism (Herve et al 1993). Subduction-related magmatism resumed when subduction again became orthogonal (Pankhurst et al. 1999). Prior to development of the Cretaceous arc, the spatial distribution of the highly oblique Late Triassic Batholith of Central Patagonia and the rather less oblique SCB (Fig. 1) is difficult to explain by convergence of the Pacific and the South American plates, differing radically from that prevailing after western Gondwana break-up in Early Cretaceous times. Early Triassic arc magmatism does not continue to the south of the North Patagonian Massif and explanation of this absence is a major challenge to understanding the evolution of Patagonia. Several factors suggest that the southern sector of the South American plate is a collage of continental blocks, some of them with strong geological and palaeontological affinities with southern Africa. Thus, palaeomagnetic and geological studies in the Falkland/Malvinas islands support a 180° tectonic rotation from a pre-Gondwana break-up position adjacent to the southeast coast of South Africa (Mitchell et al. 1986; Taylor & Shaw 1989; Marshall 1994; Curtis & Hyam 1998). An original position of part of Patagonia close to southern Africa is also consistent with the remarkable affinities of the Permian and Triassic flora of the Deseado Massif to that of the Karoo Basin (Archangelsky 1990; Artabe et al 2003). In addition, the recovery of Cambrian granodioritic orthogneiss from borehole drilling in Tierra del Fuego (Sollner et al 2000; Pankhurst et al 20036) suggests that the basement of the southern tip of South America may have also been a displaced terrane originally associated with the Cambrian proto-Pacific margin, and also represented by the Cape Granite suite in South Africa and the Ellsworth Mountain in Antarctica (Rapela et al. 2003 and references therein) (Fig. 10). In order to explain the displacement of the Falkland/Malvinas microplate, the oblique spatial arrangement of the Late Triassic magma-

tism and the 'excess space' problem in the reconstruction of the South Atlantic, Rapela & Pankhurst (1992) proposed the dextral displacement of a southern Patagonian block along the transcontinental Gastre Fault System (Fig. 1). Mafic dykes in the Falkland/Malvinas islands are coeval with (K-Ar and Ar/Ar ages of 180-192 Ma), and compositionally similar to, the Karoo dyke swarms (Cingolani & Varela 1976; Mussett & Taylor 1994), suggesting that this continental block was still adjacent to the eastern end of the Cape Fold Belt of southern Africa in Early Jurassic times, although it was most probably attached to Patagonia by the Early Cretaceous (Mussett & Taylor 1994). Palaeomagnetic results in central Patagonia provide evidence for 25-30° clockwise rotation of tens-of-kilometre-sized crustal blocks during the Late Jurassic-Early Cretaceous, similar in timing to other clockwise microplate rotations in southwestern Gondwana (Geuna et al 2000 and references therein). Although the tighter fit produced with the model of Rapela & Pankhurst (1992) may alleviate some reconstruction space constraints, the amount of dextral displacement needed along a single fault system (c. 500 km) is unrealistic (Rapela 1997). Many of the problems in pre-Cretaceous reconstructions of the South Atlantic arise because fits have been carried out using the present-day geographical boundaries of the continental fragments. However, these boundaries are actually the result of large displacement of microplates that occurred during western Gondwana breakup, the formation of large sedimentary basins and the crustal extension during CretaceousTertiary times (see also Jacques 2003). Relatively simple palinspastic reconstructions such as those recommended by MacDonald et al (1998), produce a 'tight fit' that takes into account these distorting factors, and can explain the pre-Late Jurassic geological arrangement better (Fig. 10).

Conclusions Subduction-related magmatism occurred along the southwestern (proto-Pacific) margin of Gondwana episodically over the period of more than 50 Ma prior to establishment of the Andean geotectonic cycle. The Late Triassic Batholith of Central Patagonia and the coeval metasedimentary rocks of the accretionary prism represent the last supercontinent convergence episode in the proto-Pacific margin of Patagonia. In contrast, the Early Jurassic SCB plutonic belt appears to indicate the

SUBCORDILLERAN MAGMATISM OF PATAGONIA re-establishment of subduction-related marginal magmatism further west, simultaneously with the earliest phase of rhyolite volcanism in Patagonia and the Antarctic Peninsula, as the impingement of the Karoo mantle plume and Ferrar magmatism heralded the break-up of Gondwana. An extension interval, without obvious marginal magmatism, occurred during the Mid-Jurassic stage of Chon Aike volcanism, but subduction resumed finally with the arc in the fixed position occupied by the Andean Patagonian batholith about 150 Ma, some 10-12 Ma prior to formation of the South Atlantic Ocean. Clockwise rotation through about 30° of the axis of Pacific marginal magmatism occurred during the interval 240-150 Ma, together with significant westward migration as each new episode took place. The explanation for this is not understood in detail, but has significance for the dynamics of the break-up process and must involve extensive rearrangement of fragments and segments of the preexisting continental crust. Funding for this work was provided by a Leverhulme Trust Emeritus Fellowship award to RJP and by CONICET grant PIP 02082 to CWR. The authors are indebted to Alfredo Benialgo for preparing the digital basis of Figure 2.

References ARCHANGELSKY, S. 1990. Plant distribution in Gondwana during the Late Paleozoic. In: TAYLOR, T.N. & TAYLOR, E. (eds) Antarctic Paleobiology: its role in the reconstruction of Gondwana. Springer-Verlag, Berlin, 102-117. ARCULUS, R.J. 2003. Use and abuse of the terms calcalkaline and calcalkalic. Journal of Petrology, 44, 929-935. ARTABE, A.E., MOREL, E.M. & SPALLETTI, L.A. 2003. Caracterizacion de las provincias fitogeograficas triasicas del Gondwana extratropical. Ameghiniana, 40, 387-405. BRUCE, R.M. 1988. Petrochemical evolution and physical construction of an Andean arc: evidence for the southern Patagonian batholith at 53 °S. PhD thesis, Colorado School of Mines, USA. BRUCE, R.M., NELSON, E.P., WEAVER, S.G. & Lux, D.R. 1991. Temporal and spatial variations in the southern Patagonian batholith; Constraints on magmatic arc development. In: HARMON, R.S. & RAPELA, C.W. (eds) Plutonism from Antarctica to Alaska. Geological Society of America Special Papers, 265, 1-12. BUSTEROS, A., FRANCHI, M. & LEMA, H.A. 1993. El magmatismo calcoalcalino del area de Jose de San Martin, provincia del Chubut. XII Congreso Geologico Argentino, Mendoza, Adas IV, 128-133. CHAPPELL, B.W. & WHITE, A.J.R. 1992. I- and S-type

235

granites in the Lachlan Fold Belt. Transactions of the Royal Society of Edinburgh, 83,1-26. CINGOLANI, C.A. & VARELA, R. 1976. Investigaciones geologicas y geocronologicas en el extreme sur de la Isla Gran Malvina, sector de Cabo Belgrano (Cabo Meredith). VI Congreso Geologico Argentino, Bahia Blanca, Adas I. Asociacion Geologica Argentina, Buenos Aires, 457-480. CINGOLANI, C., DALLA SALDA, L., HERVE, F, MUNIZAGA, F, PANKHURST, R.J., PARADA, M.A. & RAPELA, C.W. 1991. The magmatic evolution of northern Patagonia; new impressions of preAndean and Andean tectonics. In: HARMON, R.S. & RAPELA, C.W. (eds) Plutonism from Antarctica to Alaska. Geological Society of America Special Papers, 265, 29-44. Cox, K.G. 1992. Karoo igneous activity, and the early stages of the break-up of Gondwanaland. In: STOREY, B.C., ALABASTER, T. & PANKHURST, RJ. (eds) Magmatism and the Causes of Continental Break-up. Geological Society, London, Special Publications, 68,137-148. CURTIS, M.L. & HYAM, D.M. 1998. Late Palaeozoic to Mesozoic structural evolution of the Falkland Islands: a displaced segment of the Cape Fold Belt. Journal of the Geological Society, London, 155, 115-129. DALZIEL, I.W.D. 1981. Back-arc extension in the southern Andes, a review and critical reappraisal. Philosophical Transactions of the Royal Society of London, Series A, 300, 319-335. DEPAOLO, D.J., LINN, A.M. & SCHUBERT, G. 1991. The continental crustal age distribution; methods of determining mantle separation ages from Sm-Nd isotopic data and application to the Southwestern United States. Journal of Geophysical Research, 96, 2071-2088. DRUMMOND, M.S., DEFANT, M.J. & KEPZHINSKAS, P.K. 1996. Petrogenesis of slab-derived trondhjemitetonalite-dacite/adakite magmas. Transactions of the Royal Society of Edinburgh, 87, 205-215. DUNCAN, R.A., HOOPER, PR., REHACEK, I, MARSH, J.S. & DUNCAN, A.R. 1997. The timing and duration of the Karoo igneous event, southern Gondwana. Journal of Geophysical Research, 102, 18 127-18 138. EBY, G.N. 1990. The A-type granites: A review of their occurrence and chemical characteristics and speculations on their petrogenesis. Lithos, 26, 115-134. ELLIOT, D.H. & FLEMING, T.H. 2000. Weddell triple junction: the principal focus of Ferrar and Karoo magmatism during initial breakup of Gondwana. Geology, 28, 539-542. ENCARNACION, I, FLEMING, T.H., ELLIOT, D.H. & EALES, H.V. 1996. Synchronous emplacement of Ferrar and Karoo dolerites and the breakup of Gondwana. Geology, 24, 535-538. FERAUD, G., ALRIC, V., FORNARI, M., BERTRAND, H. & HALLER, M. 1999.40Ar/39Ar dating of the Jurassic volcanic province of Patagonia: migrating magmatism related to Gondwana break-up and subduction. Earth and Planetary Science Letters, 172, 83-96.

236

C. W. RAPELA ETAL.

FORSYTHE, R. 1981. Geological investigations of preLate Jurassic terranes in the southernmost Andes. PhD thesis, Columbia University, USA. FORSYTHE, R. 1982. The late Palaeozoic and Early Mesozoic evolution of southern South America: a plate tectonic interpretation. Journal of the Geological Society, London, 139, 671-682. FRANCHI, M.R. & PAGE, R.FN. 1980. Los basaltos cretacicos y la evolution magmatica del Chubut occidental. Revista de la Asociacion Geologica Argentina, 35, 208-229. GEUNA, S.E., SOMOZA, R., VizAN, H., FIGARI, E.G. & RINALDI, C.A. 2000. Paleomagnetism of Jurassic and Cretaceous rocks in central Patagonia: a key to constraint the timing of rotations during the breakup of Gondwana? Earth and Planetary Science Letters, 181, 145-160. GHIARA, M.R., BARBIERI, M., STANZIONE, D., HALLER, M.J., CASTORINA, E, MENDITTI, I. & TRUDU, C. 1997. Inference from basalt-rhyolite dyke swarms on the petrogenesis of a granodiorite to leucogranite suite from Patagonian Batholith (42.40°-42.50° Lat. S). Mineralogica Petrografica Acta, 15, 117-136. GODEAS, M. 1993. Geoquimica y marco tectonico de los granitoides en el Bajo de La Leona (Formation la Leona), provincia de Santa Cruz. Revista de la Asociacion Geologica Argentina, 47, 347-347. GONZALEZ BONORINO, G. 1990. Cambios relatives en el nivel del mar y su posible relation con magmatismo en el Jurasico temprano, Formation Lepa, Chubut noroccidental, Argentina. Revista de la Asociacion Geologica Argentina, 45, 129-135. GONZALEZ DIAZ, E.F. 1982. Chronological zonation of granitic plutonism in the northern Patagonian Andes of Argentina: the migration of intrusive cycles. Earth Science Reviews, 18, 365-393. GORDON, A. & ORT, M.H. 1993. Edad y correlation del plutonismo subcordillerano en las provincias de Rio Negro y Chubut (41°-42°30'L.S.). XII Congreso Geologico Argentino, Mendoza, Actas IV, 120-127. GORRING, M.L. & KAY, S.M. 2001. Mantle processes and sources of Neogene slab window magmas from southern Patagonia, Argentina. Journal of Petrology, 42, 1067-1094. HALLER, MJ. 1985. El magmatismo Mesozoico en Trevelin, Cordillera Patagonica, Argentina. 4° Congreso Geologico Chileno, Antofagasta, Actas 3, Universidad del Norte, Antofagasta, 215-234. HALLER, M.J., LINARES, M., OSTERA, H.A. & PAGE, S.M. 1999. Petrology and geochronology of the Subcordilleran Plutonic Belt of Patagonia, Argentina. // South American Symposium on Isotope Geology, Carlos Paz, Argentina, Actas, SEGEMAR, Buenos Aires, 210-214. HERVE, F. & FANNING, C.M. 2003. Early Cretaceous subduction of continental crust at the Diego de Almagro archipelago, southern Chile. Episodes, 26, 285-289. HERVE, F, PANKHURST, R.J., SUAREZ, M. & DE LA CRUZ, R. 1993. Basic magmatism in a mid-

Tertiary transtensional basin, Isla Magdalena, Aysen, Chile. Second Symposium International, Geodynamique Andine, Oxford (21-23 September). Editions de 1'ORSTOM, Colloques et Seminaires, Paris, 367-369. HERVE, R, FANNING, C.M. & PANKHURST, R.J. 2003. Detrital Zircon Age Patterns and Provenance of the metamorphic complexes of Southern Chile. Journal of South American Earth Sciences, 16, 107-123. IRVINE, T.N. & BARAGAR, W.R. 1971. A guide to the chemical classification of the common igneous rocks. Canadian Journal of Earth Sciences, 8, 523-548. JACQUES, J.M. 2003. A tectonostratigraphic synthesis of the sub-Andean basins; inferences on the position of South American intraplate accommodation zones and their control on South Atlantic opening. Journal of the Geological Society, London, 160, 703-717. KAY, S.M., RAMOS, V.A. & MARQUEZ, M. 1993. Evidence in Cerro Pampa volcanic rocks for slab-melting prior to ridge-trench collision in southern South America. Journal of Geology, 101, 703-714. KRAEMER, P. 2003. Orogenic shortening and the origin of the Patagonian Orocline (56 °S Lat.). Journal of South American Earth Sciences, 15, 731-748. LING, H.Y., FORSYTHE, R.D. & DOUGLAS, C.R. 1987. Late Paleozoic microfaunas from southernmost Chile and their relation to Gondwanaland forearc development. Geology, 13, 357-360. LIZUAIN, A. 1981. Caracteristicas y edad del plutonismo en los alrededores del lago Puelo-Provincia del Chubut. VIII Congreso Geologico Argentino, San Luis, Actas 3, 607-616. LIZUAIN, A. 1983. Geologia de la Cordillera Patagonica entre las localidades de Lago Puelo y Leleque. PhD thesis, Universidad de Buenos Aires. LIZUAIN, A. 1999. El Jurasico y Cretacico de la Patagonia y Antartida. 1. Estratigrafia y evolucion geologica del Jurasico y Cretacico de la Cordillera Patagonica Septentrional. In: CAMINOS, R. (ed.) Geologia Argentina. Subsecretaria de Mineria de la Nation, Servicio Geologico Minero Argentino, Institute de Geologia y Recursos Minerales, Anales N° 29, 433-443. LUDWIG, K.R. 1999. Isoplot/Ex, a geochronological toolkit for Microsoft Excel. Berkeley Geochronological Center Special Publication No. 1. version 2.31. 2455 Ridge Road, Berkeley, CA 94709. LUDWIG, K.R. 2001. Squid 1.00. A User's Manual. Berkeley Geochronology Center Special Publication No. 2. 2455 Ridge Road, Berkeley CA 94709. MACDONALD, D, GOMEZ-PEREZ, I., FRANZESE, J., SPALLETTI, L., PATON, D., GILPIN, R. & ANDERSEN, L. 1998. Mesozoic evolution of the South Atlantic and its influence on the Tertiary development of the region. South Atlantic Project, Report No. 10, Cambridge Artie Shelf Programme (CASP)Centro de Investigaciones Geologicas (CIG). MARSHALL, J.E.A. 1994. The Falkland Islands: a key

SUBCORDILLERAN MAGMATISM OF PATAGONIA element in Gondwana paleogeography. Tectonics, 13, 499-514. MARTIN, M, PANKHURST, R.J., FANNING, C.M., THOMSON, S.N., CALDERON, M. & HERVE, F. 2001. Age distribution of plutons across the southern Patagonian batholith: new U-Pb age data on zircons. Third South American Symposium of Isotope Geology, Pucon, Chile. CD-ROM, SERNAGEOMIN, Santiago, 585-588. MITCHELL, C, TAYLOR, G.K., Cox, K.G & SHAW, J. 1986. Are the Falkland Islands a rotated microplate? Nature, 319, 131-134. MUKASA, S.B. & DALZIEL, I.W.D. 1996. Southernmost Andes and South Georgia Island, north Scotia Ridge: zircon U-Pb and muscovite 40Ar/39Ar age constraints on tectonic evolution of southwestern Gondwanaland. Journal of South American Earth Sciences, 9, 349-365. MUSSETT, A.E. & TAYLOR, G.K. 1994. 40Ar-39Ar ages of dykes from the Falkland Islands with implications for the break-up of southern Gondwanaland. Journal of the Geological Society, London, 151, 79-81. OLIVARES, B., CEMBRANO, J., HERVE, E, LOPEZ, G. & PRIOR, D. 2003. Analisis estructural de rocas de una zona de cizalle ductil en isla Diego de Almagro, sur de Chile. Revista Geologica de Chile, 309, 39-52. PAGE, S. & PAGE, R. 1999. Las diabasas y gabros del Jurasico de la Precordillera del Chubut. In: CAMINOS, R. (ed.) Geologia Argentina. Subsecretaria de Mineria de la Nacion, Servicio Geologico Minero Argentine, Institute de Geologia y Recursos Minerales, Anales No. 29, 489-495. PALMER, A.R. & GEISSMAN, J. (compilers) 1999. The 1999 Geologic Time Scale. Geological Society of America, Boulder, CO. PANKHURST, RJ. 1982. Rb-Sr geochronology of Graham Land, Antarctica. Journal of the Geological Society, London, 139, 701-711. PANKHURST, RJ. & RAPELA, C.W. 1995. Production of Jurassic rhyolite by anatexis in the lower crust of Patagonia. Earth and Planetary Science Letters, 134, 23-36. PANKHURST, R.J., RAPELA, C.W., CAMINOS, R., LLAMBIAS, EJ. & PARICA, C. 1992. A revised age for the granites of the central Somuncura batholith, North Patagonian Massif. Journal of South American Earth Sciences, 5, 321-325. PANKHURST, R.J., RAPELA, C.W. & MARQUEZ, M. 1993. Geocronologia y petrogenesis de los granitoides jurasicos del noreste del Macizo del Deseado. XII Congreso Geologico Argentino, Mendoza, Adas IV, 134-141. PANKHURST, R.J., HERVE, F. & RAPELA, C.W. 1994. Sm-Nd evidence for the Grenvillian provenance of the metasedimentary basement of Southern Chile and West Antarctica. In: CAMPOS, E. & CECIONI, A. (eds) Actas del VII Congreso Geologico Chileno, Vol. II. Universidad del Norte Chile, Departamento de Geociencias, Facultad de Ciencias. Antofagasta, Chile, 1414-1418. PANKHURST, RJ., LEAT, P.T., SRUOGA, P. RAPELA, C.W, MARQUEZ, M., STOREY, B.C. & RILEY,T.R. 1998.

237

The Chon-Aike silicic igneous province of Patagonia and related rocks in West Antarctica: a silicic LIP. Journal of Volcanology and Geothermal Research, 81, 113-136. PANKHURST, R.J., WEAVER, S.D., HERVE, F. & LARRONDO, P. 1999. Mesozoic-Cenozoic evolution of the North Patagonian Batholith in Aysen, southern Chile. Journal of the Geological Society, London, 156, 673-694. PANKHURST, R.J., RILEY, T.R., KELLEY, S. & FANNING, CM. 2000. Episodic silicic volcanism in Patagonia and the Antarctic Peninsula: chronology of magmatism associated with the break-up of Gondwana. Journal of Petrology, 41, 605-625. PANKHURST, RJ., HERVE, F, FANNING, C.M. & SUAREZ, M. 20030. Coeval plutonic and volcanic activity in the Patagonian Andes: the Patagonian Batholith and the Ibanez and Divisadero Formations, Aysen, southern Chile. X Congreso Geologico Chileno, Actas, SERNAGEOMIN, Santiago (CD-ROM). PANKHURST, R.J., RAPELA, C.W, LOSKE, W.P., FANNING, C.M. & MARQUEZ, M. 2003/?. Chronological study of the pre-Permian basement rocks of southern Patagonia. Journal of South American Earth Sciences, 16, 27^44. PEARCE, J.A. 1983. Role of the sub-continental lithosphere in magma genesis at active continental margins. In: HAWKESWORTH, CJ. & NORRY, MJ. (eds) Continental Basalts and Mantle Xenoliths. Shiva, Nantwich, 230-249. PEARCE, J.A., HARRIS, N.B.W & TINDLE, A.G. 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. Journal of Petrology, 25, 956-983. RAPALINI, A.E., HERVE, F, RAMOS, V.A. & SINGER, S. 2001. Paleomagnetic evidence of a very large counterclockwise rotation of the Madre de Dios archipelago, southern Chile. Earth and Planetary Science Letters, 184, 471^87. RAPELA, C.W. 1997. El sistemas de fallas de Gastre: e pur si muove. Revista de la Asociacion Geologica Argentina, 52, 219-222. RAPELA, C.W. 2001. El magmatismo Triasico-Jurasico de la Patagonia y su ambiente geodinamico. In: ARTABE, A.E., MOREL, E.M. & ZAMUNER, A.B. (eds) El Sistema Triasico en la Argentina. Fundacion Museo de La Plata 'Francisco Pascasio Moreno', La Plata, 69-80. RAPELA, C.W. & KAY, S. 1988. The Late Paleozoic to Recent magmatic evolution of northern Patagonia. Episodes, 11, 175-182. RAPELA, C.W & PANKHURST, RJ. 1992. The granites of northern Patagonia and the Gastre Fault System in relation to the break-up of Gondwana. In: STOREY, B.C., ALABASTER, T. & PANKHURST, R J. (eds) Magmatism and the Causes of Continental Break-up. Geological Society, London, Special Publications, 68, 209-220. RAPELA, C.W. & PANKHURST, RJ. 1996. Monzonite Suites: the innermost Cordilleran plutonism of Patagonia. Transactions of the Royal Society of Edinburgh, Earth Sciences, 87, 193-203. RAPELA, C.W, PANKHURST, RJ. & HARRISON, S.M.

238

C.W. RAPELA ETAL.

1992. Triassic 'Gondwana' granites of the Gastre district, North Patagonian Massif. Transactions of the Royal Society of Edinburgh, Earth Sciences, 83, 291-304. RAPELA, C.W., PANKHURST, R.J., LLAMBIAS, E.J., LABUDIA, C. & ARTABE, A. 1996. 'Gondwana' magmatism of Patagonia: Inner Cordilleran calcalkaline batholiths and bimodal volcanic provinces. 3rd International Symposium on 'Andean Geodynamics', Saint Malo, Extended Abstracts, ORSTOM, Paris, 791-794. RAPELA, C.W., PANKHURST, R.J., FANNING, C.M. & GRECCO, L.E. 2003. Basement evolution of the Sierra de la Ventana Fold Belt: new evidence for Cambrian continental rifting along the southern margin of Gondwana. Journal of the Geological Society, London, 160, 613-628. REMANE, J., FAURE-MURET, A. & ODIN, G.S. (compilers) 2000. International Stratigraphic Chart. International Commission on Stratigraphy. RILEY, T.R., LEAT, P.T., PANKHURST, R.J. & HARRIS, C. 2001. Origins of large-volume rhyolite volcanism in the Antarctic Peninsula and Patagonia by crustal melting. Journal of Petrology, 42, 1043-1065. RILEY, T.R., MILLAR, I.L., WATKEYS, M.K., CURTIS, M.L., LEAT, P.T., KLAUSEN, M.B. & FANNING, C.M. 2004. U-Pb zircon (SHRIMP) ages for the Lebombo rhyloites, South Africa: refining the duration of Karoo volcanism. Journal of the Geological Society, London, 161, 547-550. ROLANDO, A.P., HARTMANN, L.A., SANTOS, J.O.S., FERNANDEZ, R.R., ETCHEVERRY, R.O., SCHALAMUK, LA. & MCNAUGHTON, NJ. 2002. SHRIMP zircon U-Pb evidence for extended Mesozoic magmatism in the Patagonian batholith and assimilation of Archaean crustal components. Journal of South American Earth Sciences, 15, 267-283. SOLLNER, R, MILLER, H. & HERVE, M. 2000. An Early Cambrian granodiorite age from the pre-Andean basement of Tierra del Fuego (Chile): the missing link between South America and Antarctica? Journal of South American Earth Sciences, 13, 163-177. SPIKERMANN, J.P. 1978. Contribution al conocimiento de la intrusividad en el paleozoico de la region extrandina de Chubut. Revista Asociacion Geologica Argentina, 23, 17-35. SPIKERMANN, IP, STRELIN, I, MARSHALL, P. ETAL. 1988. Geologia del area El Batolito Aleusco, Departamento de Languineo, provincia de Chubut. Revista de la Asociacion Argentina de Miner alogia, Petrologia y Sedimentologia, 19, 39-48. SPIKERMANN, J.P, STRELIN, X, MARSHALL, P. ETAL. 1989. Caracterizacion geologica del Batolito Aleusco, Departamento de Languineo, provincia, provincia de Chubut. Revista de la Asociacion Argentina de Mineralogia, Petrologia y Sedimentologia, 20, 33-42. STERN, Cn.R. & KILIAN, R. 1996. Role of the subducted slab, mantle wedge and continental crust in the generation of adakites from the Andean Austral Volcanic Zone. Contributions to Mineralogy and Petrology, 123, 263-281.

STOREY, B.C. 1995. The role of mantle plumes in continental break-up: case histories from Gondwanaland. Nature, 377, 301-308. STOREY, B.C., ALABASTER, T., HOLE, M J., PANKHURST, R.J. & WEVER, H.E. 1992. Role of subductionplate boundary forces during the initial stages of Gondwana breakup: evidence from the protoPacific margin of Gondwana. In: STOREY, B.C., ALABASTER, T. & PANKHURST, R.J. (eds) Magmatism and the Causes of Continental Break-up. Geological Society, London, Special Publications, 68,149-163. SUAREZ, M. & DE LA CRUZ, R. 19970. Edades K-Ar del grupo Ibanez en la parte oriental del Lago General Carrera (46°-47° LS), Aysen, Chile. VIII Congreso Geologico Chileno, Antofagasta, Universidad del Norte, Antofagasta, Chile. Actas 2. 1548-1551. SUAREZ, M. & DE LA CRUZ, R. 19976. Cronologia magmatica de Aysen sur (45°-48°30' L.S.), Aysen, Chile. VIII Congreso Geologico Chileno, Antofagasta, Universidad del Norte, Antofagasta, Chile. Actas 2. 1543-1547. SUAREZ, M. & DE LA CRUZ, R. 2001. Jurassic to Miocene K-Ar dates from eastern central Patagonian Cordillera plutons, Chile (45°-48 °S). Geological Magazine, 138, 53-66. TAYLOR, GK. & SHAW, J. 1989. The Falkland Islands: New Palaeomagnetic Data and their Origin as a Displaced Terrane from Southern Africa. American Geophysical Union, Geophysical Monographs, 50, 59-72. TERA, F. & WASSERBURG, G. 1972. U-Th-Pb systematics in three Apollo 14 basalts and the problem of initial Pb in lunar rocks. Earth and Planetary Science Letters, 14, 281-304. THOMPSON, R.N., MORRISON, M.A., HENDRY, G.L. & PARRY, S.J. 1984. An assessment of the relative roles of the crust and mantle in magma genesis: an elemental approach. Philosophical Transactions of the Royal Society, London, Series A, 310, 549-590. THOMSON, S.N. & HERVE, F. 2002. New time constraints for the age of metamorphism at the ancestral Pacific Gondwana margin of southern Chile (42-52 °S). Revista Geologica de Chile, 29, 255-271. TURNER, J.C. 1982. Descripcion geologica de la Hoja 44c, Tecka, provincia del Chubut. Servicio Geologico Nacional, Boletin 180. TURNER, S., REGELOUS, M., KELLEY, S., HAWKESWORTH, C. & MANTOVANI, M. 1994. Magmatism and continental break-up in the South Atlantic: high precision 40Ar-39Ar geochronology. Earth and Planetary Science Letters, 121, 333-348. ULIANA, M.A. & LEGARRETA, L. 1999. Jurasico y Cretacico de la cuenca del Golfo de San Jorge. In: CAMINOS, R. (ed.) Geologia Argentina. Subsecretaria de Mineria de la Nation, Servicio Geologico Minero Argentine, Institute de Geologia y Recursos Minerales, Anales N° 29, 496-510. WEAVER, S.G. 1988. The Patagonian batholith at 48 °S latitude, Chile: implication for the petrochemical and geochemical evolution of calc-alkaline

SUBCORDILLERAN MAGMATISM OF PATAGONIA batholiths. PhD thesis, Colorado School of Mines, USA. WELLS, P. 1978. The geochemistry of the Patagonian batholith between 45 °S and 46 °S latitude. MSc thesis, University of Birmingham, UK. WHALEN, J.B., CURRIE, K.L. & CHAPPELL, B.W. 1987. A-type granites: geochemical characteristics, discrimination and petrogenesis. Contributions to Mineralogy and Petrology, 95, 407-419.

239

WILLIAMS, I.S. 1998. U-Th-Pb Geochronology by Ion Microprobe. In: McKiBBEN, M.A., SHANKS, WC. Ill & RIDLEY, W.I. (eds) Applications of microanalytical techniques to understanding mineralizing processes. Reviews in Economic Geology, 7, 1-35. ZINDLER, A. & HART, S.R. 1986. Chemical geodynamics. Annual Review of Earth and Planetary Sciences, 14, 493-571.

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The Famatina complex (NW Argentina): back-docking of an island arc or terrane accretion? Early Palaeozoic geodynamics at the western Gondwana margin HUBERT MILLER & FRANK SOLLNER Department fur Geo- und Umweltwissenschaften der LMU, Munchen, Sektion Geologic, Luisenstr. 37, D-80333 Munchen (e-mail: [email protected]; [email protected]) Abstract: This paper concerns sedimentary, volcanic, metamorphic and igneous rocks of the Sierra de Famatina and adjacent rock series in NW Argentina, for which the name 'Famatina system' has been used widely in the literature. It is suggested that this is renamed the 'Famatina complex', since 'system' is an internationally defined stratigraphical term. The Famatina complex is considered to have formed as an island arc on continental crust, represented in its diversity by metasediments and meta-volcanic rocks of the Sierra de Famatina, with a corresponding back-arc basin exposed in the medium- to high-grade metamorphic rocks of the adjacent Las Termas belt to the east. The Famatina complex was deposited in contact with the eastern Pampean complex ('Pampia terrane') at the western active continental margin of Gondwana in latest Proterozoic and Ordovician times. Deposition was, in part, accompanied by voluminous Ordovician magmatism of calc-alkaline composition. NNW-SSE-striking shear zones, dated previously at 402 ± 2 Ma, are interpreted as marking the final stage of collision of the island-arc/back-arc/continent complex. The dynamics of crustal block movements along the most prominent (TIPA) shear zone indicate overthrusting of the eastern Pampean series onto the western Famatina series and, hence, uplift and cooling of the eastern block must have occurred earlier than cooling of the western one. The composition of inherited components in Famatina metasediments and meta-granitoids argues for autochthonous arc-continent convergence rather than accretion of an exotic terrane.

In compiling the geological history of NW Argentina, Acenolaza & Toselli (1976) introduced the term 'Pampean cycle' for deposition, metamorphism, deformation and igneous activity in late Precambrian to Mid-Cambrian time. The Pampean cycle obtains its name from the Pampean Ranges (Fig. 1), the main morphostructural mountain ranges of NW Argentina, which rise from the great plains that extend to the east (the 'Pampa'). The Pampean orogenic cycle has been proposed as encompassing the time span 550-515 Ma (Sollner et al 20006). The area affected by this orogenic cycle reaches far beyond the point of its definition (the MidCambrian angular unconformity in the Quebrada del Toro, Salta province), from the Argentine/Bolivian border (22° S) to the central Argentine Sierra de Cordoba (Rapela et al 19980, b\ Schwartz et al. 2003; Miro 2004) and even to Patagonia and Antarctica (Acenolaza & Miller 1982; Sollner et al. 20006). Acenolaza & Toselli (1976) used the term 'Famatinian cycle' for comparable events in the Ordovician, which extended locally to Silurian times. It derives its name from the Sierra de Famatina, at 67°20' to 68°20' W and 27-30° S in

the provinces of La Rioja and Catamarca, immediately west of the eastern Pampean Ranges. In its narrow sense, the Sierra de Famatina is characterized by widespread Ordovician plutonism and volcanism. This paper suggests changing the term 'Sistema de Famatina', used by Argentinian (Acenolaza et al 1996) and other authors for the Sierra de Famatina and the immediately surrounding ranges composed mainly of granitoid rocks, to 'Famatina complex', in order to avoid confusion with the general meaning of 'system' as a welldefined stratigraphical term. The Famatinian orogenic cycle as a whole is recognized throughout a much wider region, along the western border of Gondwana from Bolivia at least as far as the Sierras de Cordoba and San Luis (Pankhurst & Rapela 1998; Pankhurst et al 1998; 2000; Saavedra et al 1998). In comparison with other regions in South America, the 'Pampean cycle' is more or less contemporaneous with the Pan-African/ Brasiliano episode of Earth history in the late Precambrian. The 'Famatinian cycle', however, currently unique in the geological history of South America, is comparable with

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 241-256. 0305-8719/$15.00 © The Geological Society of London 2005.

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Fig. 1. Morpho-structural units of NW Argentina.

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the Caledonian/Taconian/Appalachian system of Laurentia and Baltica. The zone where rock units of both complexes are directly in contact with each other is of particular interest. This zone, approximately 100 km wide, is characterized by a two-phase metamorphic overprint and overwhelming igneous activity. East of the Sierra de Famatina, and best exposed in the southwestern part of the Sierra de Fiambala (Fig. 2), medium- to high-grade metamorphic rocks crop out and were designated recently the 'Las Termas belt' (Sollner et al 2001). The lithology of this belt differs significantly from adjacent rock units (see below). The most outstanding features of the Las Termas belt are major shear zones (at least four) striking NNW-SSE, more or less parallel to the trend of the belt (Fig. 2). Similar shear zones occur in a strip up to 100 km wide, which comprises the western part of the Sierra de Fiambala and, to the east, the whole Sierra de Velasco (Fig. 2). Caminos (1979) and Acenolaza & Toselli (1981) demonstrated the regional extent of these mylonite zones. They are now considered to stretch from south of Bolivia to Cordoba and San Luis province (e.g. Whitmeyer & Simpson 2003) and are thought to affect the entire preCarboniferous basement. A summary of the Early and Mid-Palaeozoic orogenic history is given, for example, by Pankhurst & Rapela (1998). The Precordillera of Mendoza and San Juan and the westernmost parts of the Andean basement of Argentina ('western Pampean Ranges') have been assigned jointly to the apparently exotic 'Cuyania' or Precordillera terrane (e.g. Ramos et al 1986; 1998; Astini et al 1995; Acenolaza et al 2002). The central aim of the present paper is to focus on the significance of the Las Termas belt as part of the Famatina complex, its boundaries to the adjacent units to the east and west and its geochronological history, affected mostly by the huge attendant ductile shear zones.

The Pampean cycle (eastern Pampean Ranges) The eastern Pampean Ranges are mainly composed of metasedimentary rocks with a very low- to high-grade metamorphic overprint and corresponding polyphase deformation (Miller et al 1994). Within the very low-grade metamorphic rocks, a Late Proterozoic to Early Cambrian age has been indicated by trace fossils (Durand & Acenolaza 1990). In the Sierra de San Luis, Early Cambrian sedimentation has been supported by age determinations on a

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syn-sedimentary meta-volcanic rock (529 ± 12 Ma, U-Pb on zircons; Sollner et al 20000). The most widespread lithology of the eastern Pampean Ranges consists of sandstone/claystone series and their metamorphic equivalents, accompanied sporadically by marbles and metavolcanic rocks. The provenance of the sediments has been determined by light mineral composition studies (quartz-feldspar-lithics) as predominantly recycled orogenic components, prevalently from the cratonic area to the northeast (Jezek & Miller 1987). They were transported and deposited mostly by turbidity currents. Schwartz & Gromet (2004) also showed that sources of detrital zircons in Cambrian or older metasedimentary rocks of the Sierra de Cordoba were in Gondwana, rather than any Laurentian region. Folding is polyphase, with up to three or four stages (Willner et al 1987; Willner 1990). Sedimentation of the basement rocks ceased before the Mid-Cambrian. Post-orogenic siliciclastic sedimentation of Late Cambrian or Ordovician age is of minor importance, but is found in the Meson Group of the eastern Cordillera (Salta and Jujuy provinces of northwesternmost Argentina) and the La Cebila Formation in the Sierra de Ambato, Fig. 2). In the study area, the metamorphic overprint and tectonic style of the intensively deformed Pampean rock series are best exposed east of Fiambala, in the northeastern part of the Sierra de Fiambala (Fig. 2). In general, post-orogenic sediments postdating both Pampean and Famatinian rock series are younger than Early Carboniferous (Paganzo Group). Granitoid rocks, intruding the metasediments, are widespread. The intrusion ages show three main phases of magmatic activity in the Mid-Cambrian and Ordovician and, less voluminously, in the Carboniferous (e.g. Knuver & Miller 1982; Lork et al 1991; Rapela et al 19980, b\ 2001; Schwartz et al 2003). The first two phases can be correlated with the Pampean and Famatinian tectonometamorphic events.

The superposition of Pampean and Famatinian magmatic and metamorphic events When defining the Pampean and Famatinian cycles, Acenolaza & Toselli (1976) stated that Famatinian tectonomagmatic and metamorphic events were superimposed widely on the Pampean Orogen. This overprinting was demonstrated first in detail by Bachmann et al (1986), using Rb-Sr thin slab and mineral/whole-rock

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Fig. 2. Generalized geological map of the area studied in detail. Slightly modified from Hockenreiner et al (2003, fig. 1), 'Dating the TIPA shear zone: an early Devonian terrane boundary between the Famatinian and Pampean systems (NW Argentina)', Journal of South American Earth Sciences, 16, 2003, pp. 45-66, © 2003, with permission from Elsevier.

isochrons on banded schists from the Sierra de Ancasti and Sierra de Aconquija. Pampean metamorphism was dated at 570-540 Ma and the Famatinian overprint at 470-435 Ma. Lucassen & Becchio (2003) obtained similar ages of superimposed metamorphic events at c. 530-510 Ma and c. 470-420 Ma based on U-Pb titanite dating of high-grade metamorphic rocks from several sites in the Puna and western Pampean Ranges. Lucassen et al (2000) showed that events of 'Pampean' and 'Famatinian' age

occur jointly in wide areas of northwest Argentina and northern Chile. Metamorphic and magmatic events of the Pampean and Famatinian cycle are superimposed also in the Sierra de Cordoba, the southeasternmost part of the Pampean Ranges (e.g. Rapela et al 19980; Sims et al 1998). On the other hand, Whitmeyer & Simpson (2004) argue for significant separation between the Sierra de Cordoba and the Sierra de San Luis before the Late Ordovician, due to a difference in the time of

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peak metamorphism in these sierras. In North Chile and South Peru, Loewy et al (2004) found Pampean and Famatinian orogenic events superimposed on Proterozoic orogenies of the Arequipa terrane.

The Famatina complex (an island arc on continental crust) Geographically, the Famatina complex lies between the eastern Pampean Ranges to the east and the exotic terrane of the Precordillera of Mendoza and San Juan to the west (Fig. 1). The 'Famatinian arc', as an orogenic cycle in the sense of Acenolaza & Toselli (1976), is represented by intense magmatism within continental crust, from southernmost Salta province to the Sierra de San Luis (Rapela et al 19980). Alternatively, it has been interpreted as an Ordovician collisional belt between the 'Occidentalia terrane' and the Gondwana craton, from Bolivia to Patagonia (Dalla Salda et al 1992). The Famatina complex, the specific subject of this paper, is that part of the Famatinian arc located in the central Sierra de Famatina and several surrounding mountain ranges of minor importance; it extends for about 300 km (Fig. 3). Sedimentation started in the Vendian, with siliciclastic sediments similar to those of the Puncoviscana Formation and its equivalents in the Pampean Ranges and the eastern Cordillera to the north (Salta and Jujuy provinces, Fig. 1). After a Mid-Cambrian hiatus, volcaniclastic sedimentation started in the uppermost Cambrian and lasted to the MidOrdovician (Saavedra et al 1998). These sedimentary and volcaniclastic series are absent within the eastern Sierras Pampeanas. Tortello et al (1996) presented a palaeontological documentation of the fossiliferous rocks, while a table summarizing the stratigraphy was presented by Acenolaza et al (1996). Fanning et al (2004) have reported recently a U-Pb zircon SHRIMP age of 468 ± 3 Ma for an interbedded porphyritic rhyolite from Sierra de las Planchadas. East of Villa Union, in the Quebrada Chuschm, the volcano-sedimentary sequence is exposed best in a 2000 m thick sequence with alternating slates, cherts, rhyolites and olistostromes, whose components are mostly rhyodacites and rhyolites (Mannheim & Miller 1996). The palaeogeographical situation (Fig. 4) has been summarized by Mangano & Buatois (1996). Siliciclastic sediments are of craton interior and recycled orogenic origin (Fig. 5) (Clemens & Miller 1996), similar to those deposited in the northern Puna and the eastern

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Cordillera regions (Bahlburg 1990). Heavy mineral fractions of these sediments always contain high percentages of zircon, are partly rich in tourmaline and slightly enriched in sphene, but are very poor in garnet and other typical metamorphic minerals. This may point to a provenance area prevalently composed of granitoid rocks. Zircon morphology studies document sedimentary transport from afar, since zircons with purely magmatic shape are scarce (Clemens & Miller 1996). This is valid for the Cambrian Negro Peinado Formation as well as for rocks of Ordovician age. Euhedral zircons of volcanogenic origin are clearly more abundant in Ordovician than in Cambrian sediments of the Famatina complex. Granitoids intruding Famatinian sediments are common, but limited to the Early and MidOrdovician. The Ordovician magmatism in the Sierra de Famatina is of calc-alkaline composition (Fig. 6) and, thus, can be related to a subduction origin (Mannheim & Miller 1996; Toselli et al 1996; Saavedra et al 1998). The initial 87Sr/86Sr ratios of these magmatic rocks are about 0.708 (Mannheim & Miller 1996; Pankhurst et al 1998). This clearly means that melt composition was influenced by crustal rock assimilation. In consequence, the Famatinian arc developed on continental crust as argued by Pankhurst et al (1998) and Rapela (2000) and accepted widely. The local volcanic arc of the Sierra de Famatina was separated from the Gondwana hinterland (Tampia terrane') by spreading in the Late Cambrian and thus took on an individual development, independent of that of the eastern Pampean Ranges.

The Las Termas belt (a back-arc basin) If there was an island arc, there should also have been a back-arc basin. The Las Termas belt is suggested as representing remnants of this backarc basin. The varied and unusual rock series around the thermal springs of Fiambala north of Tinogasta (Fig. 2) and the 'Fiambala metagabbronorite (FMGN)', a meta-igneous body of about 13 X 0.4 km, elongated in the strike direction, are considered to be typical nonsubducted remains of such a back-arc basin (Mannheim & Miller 1996; Neugebauer & Miller 1996). These rock series are characterized by sillimanite-garnet-biotite gneisses and migmatites, which indicate a medium- to highgrade metamorphic overprint (Grissom 1991), calc-silicate rocks, rare marbles, meta-pelites and meta-gabbronorites. The fill of the back-arc basin was intruded by numerous basic and ultrabasic dykes, with several stretched

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Fig. 3. The distribution of granitoids and stratified (sedimentary and volcanic) rocks in the Sierra de Famatina (Famatina complex). Figure modified after Clemens & Miller (1996).

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Fig. 4. Generalized sketch of sedimentation and volcanism of the Vuelta de Las Tolas member (Suri Formation, Arenigian), a situation typical of rock-forming processes during the Ordovician in the Sierra de Famatina (modified after Mangano & Buatois 1996).

bodies aligned along the principal foliation. The intrusion of these basic and ultrabasic rocks was probably of Mid- to Late Cambrian age (Grissom et al 1998). The varied lithological composition characterizes the series as a small, discontinuous and short-lived back-arc basin, which opened in the Mid-Cambrian. Island-arc

volcanic activity ceased in the Mid-Ordovician, as shown by the palaeontological evidence (Acenolaza et al 1996). The basin closed in the Mid-Ordovician, contemporaneously with the cessation of sedimentation and volcanism in the Sierra de Famatina. Exact sedimentation ages of the varied lithologies cannot be determined

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Fig. 5. Modal provenance plot for sedimentary rocks of the Sierra de Famatina (all observed localities and ages). Analysed samples define the enclosed field, representing homogeneous continental provenance sites (modified after Clemens & Miller 1996). Q, quartzose grains; F, feldspar grains; L, lithic fragments.

Fig. 6. Subvolcanic rocks of the Sierra de Famatina display sub-alkaline, bimodal characteristics. Crosses: Post-mid-Ordovician dykes. Modified after Mannheim & Miller (1996).

due to the metamorphic overprint. This unit, developed between the rather monotonous metasedimentary sequence of the Pampean complex to the east (Los Ratones Formation, Neugebauer & Miller 1996) and the Famatina volcanic arc to the west (a chiefly volcanosedimentary sequence), was named the 'Las

Termas belt' by Sollner et al (2001). Recently, the La Aguadita Formation, Cambrian according to Acenolaza et al. (1996), situated at the eastern edge of the Sierra de Famatina, has been dated as post-Early Ordovician by SHRIMP analysis of detrital zircons (Astini et al 2003). If the 480 Ma ages are not due merely to thermal alteration by the surrounding Ordovician granitoids, these sediments may correspond to a younger phase of the Las Termas belt. Astini et al (2003) consider them as having developed in a back-arc setting environment. Granitoid magmas spread throughout the Famatina complex, including the Las Termas belt. In general, the meta-granodiorites of the Las Termas belt are coarse grained and porphyritic and display K-feldspar crystals of up to 8 cm in length and clusters of biotite. Colours vary from pale grey to reddish. Inclusions of metasedimentary xenoliths are common and the magmatic texture is preserved largely. Zircon fractions, analysed to obtain the intrusion age of the meta-granodiorites of the Las Termas belt yielded results of 467 ± 4 Ma and 488 ± 4 Ma (Sollner et al 2001; Hockenreiner 2003; Hockenreiner et al 2003). The further geochronological evolution of the Las Termas belt is shown in Figure 7. Cooling ages of muscovites give 425 Ma, so that the metamorphic overprint of the 'La Puntilla orthogneiss' occurred probably in the late Ordovician. Upper intercept ages of zircon discordias in the concordia diagram of about 1268 Ma and 1828 Ma (Hockenreiner et al 2003) may represent different portions of detrital zircon cores from different sources of the crustal material involved in generation of the granite magmas. In addition, Transamazonian (c. 2.0-2.2 Ga) and Sunsas ('Grenville age') crustal detritus, and younger late Precambrian to Palaeozoic components, may be involved; only detailed SHRIMP analyses on zircon grains could validate these inferences precisely. Formation of the meta-granodiorites predominantly by crustal melting is well documented by £ Nd (470 Ma) values of approximately -6, and calculated two-stage Nd model ages of the meta-granitoids (and the mylonites) indicate very uniform crustal residence ages of approximately 1.6 Ga (for details see Hockenreiner et al 2003). Because the Nd model ages are in general accordance with the mean ages of inherited cores in zircons, detritus can be explained best by a mixture of an older Proterozoic crustal source and a younger upper mantle component (e.g. Late Proterozoic, Brasiliano cycle). Nevertheless, a simple Mid-Proterozoic crustal source for the granitoids cannot be discarded currently.

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Fig. 7. Timetable, representing ages of tectono-thermal events from the transition zone of the Famatina and Pampean complexes (Las Termas belt). All data, except those marked by 1) and 2) are from Hockenreiner (2003) and Hockenreiner et al (2003); 1) data from Grissom et al. (1998); 2) data from Sollner et al (2001). Nd two-stage model ages of granitoids and mylonites indicate very homogeneous average crustal residence ages of 1.59-1.64 Ga (T2 = 470 Ma in model calculations).

Several deformation events (at least two major isoclinal folding phases) can be distinguished in the metasediments of the Las Termas belt (Neugebauer 1996; Neugebauer & Miller

1996; Fig. 8). The Ordovician plutons normally show only weak deformation and seem to have been emplaced during the final stage of regional deformation of the metasediments. Therefore,

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Fig. 8. Typical polyphase deformation of metasediments of the Las Termas belt (copied from a photograph made on the road to the Termas de Fiambala). Modified after Neugebauer & Miller (1996).

Famatinian igneous activity post-dates an early profound regional tectono-metamorphic phase, but pre-dates a less intense final overprint. This agrees with the results of Grissom et al. (1998), who argued for Famatinian peak metamorphic conditions between 505 Ma and 493 Ma on the basis of U-Pb monazite ages. Muscovite cooling ages of 425 Ma (Fig. 7) indicate that the peak of metamorphic overprint in the Las Termas belt was passed by the Late Silurian.

Geodynamics of back-arc closure: the TIPA (Tinogasta-Pituil-Antinaco) shear zone The back-arc basin (Las Termas belt) behind the Famatina arc was an area of pronounced weakness and thinning in this crustal segment and became the site of tectonic displacement. Movements took place mainly on NNW-striking reverse faults with a dextral transcurrent component (Hockenreiner et al 2003). Several, sub-parallel ductile mylonite zones striking NNW-SSE can be traced in the Sierras de Fiambala, Copacabana and Velasco (Fig. 2), the TIPA (Tinogasta-Pituil-Antinaco; Lopez & Toselli 1993) shear zone being the most prominent one (up to 2 km wide).

Hockenreiner (2003) and Hockenreiner et al. (2003) investigated the tectono-thermal history of the TIPA shear zone by dating mineral growth and closure in the involved granitoids. The results are included in Figure 7, where it may be seen that the cooling history differs substantially on either side of the TIPA shear zone. The Las Termas belt series east of the TIPA shear zone underwent rapid uplift, whereas the western block remained constant or was subjected to increasing thermal conditions during mylonitization, possibly caused by subsidence of the subducting plate (Fig. 9). A simultaneous increase in temperature up to 587 ± 50 °C is documented by garnet growth (indicative of amphibolite facies conditions; garnet-biotite thermobarometry, for details see Hockenreiner et al 2003). Field observations and geochemical analyses reveal that extreme ductile deformation resulted in an intensive mixture of injected material and host rocks. This was demonstrated particularly by mixing models considering Zr and REE contents of the protoliths and the mylonites. The infiltration of melts and fluids is estimated to account for 75% of the mylonite composition (Hockenreiner 2003; Hockenreiner et al 2003). Whereas garnets of up to 5 mm diameter occur abundantly in the mylonitic rocks, the precursor rocks outside the shear zones are devoid of garnet. This allowed dating of the mylonitization process using the Sm-Nd isotopic system (Hockenreiner et al 2003). A four-point garnet isochron yielded an age of 402 ± 2 Ma, confirming that all garnets were cogenetic and crystallized simultaneously during mylonitization. This fundamental deformation and thermal climax is correlated with the final collision between the Famatina magmatic arc and the Pampean hinterland.

Post-shear crustal heating and igneous activity Rb-Sr age determinations of minerals from these rocks provide information on the cooling

Fig. 9. Schematic model of the erogenic evolution of the Pampean and Famatina systems in a cross-section north of Tinogasta (not to scale) for (a) Ordovician and (b) Early Devonian to Carboniferous times. In (a), back-arc spreading led to formation of the Las Termas belt sediment series on continental crust, accompanied by intrusion of partly mantle-derived basic rocks and of granitoids prevalently of crustal origin. In (b), compression of the orogenic belt led to back-thrusting of the Pampean complex over the Famatina complex. The rock series of the Las Termas belt, squeezed between the two complexes, suffered intensive deformation focused in ductile shear zones, with the TIPA shear zone as the most prominent one. The influx of melt and fluids during high-grade ductile deformation induced syn-kinematic garnet growth in the granitoid-mylonites at 402 ± 2 Ma. Taken from Hockenreiner et al (2003, fig. 11): 'Dating the TIPA shear zone: an early Devonian terrane boundary between the Famatinian and Pampean systems (NW Argentina)', Journal of South American Earth Sciences, 16, 2003, pp. 45-66, © 2003, with permission from Elsevier.

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history down to 300 °C (see Fig. 7). A significant difference in Rb-Sr muscovite cooling ages of rocks east (425 ± 17 Ma) and west (390 ± 10 Ma) of the TIPA shear zone, together with published 40 Ar/39Ar cooling ages of amphiboles from the Sierra de Fiambala (Grissom et al 1998), provides additional support for the hypothesis of block movements along this zone (Fig. 7). In consequence of these movements, the eastern part of the Las Termas belt and the metamorphic series of the eastern Pampean complex were thrust over the Famatinian rock series. The impetus for this motion may have derived from the subducting Famatina microplate (Fig. 9). Increased descent of the plate could have resulted in the formation of a slab-window configuration, with an attendant increase in heat flow to the crust. The re-heating process, which affected the crust in the Carboniferous, is documented by U-Pb apatite ages of Famatinian meta-granodiorites (342 ± 2 Ma) and the TIPA shear zone mylonite (328 ± 3 Ma; Hockenreiner et al. 2003). Crustal heating probably initiated the formation of large amounts of granitic melts, which invaded the higher crustal levels. This hypothetical scenario is recorded in Figure 9. Final cooling and uplift in late Carboniferous times is well documented by Rb-Sr biotite ages averaging 300 Ma ± 4 Ma (Fig. 7). The cooling history from c. 340 Ma to 300 Ma suggests a cooling rate of 3-4 °C Ma"1, which corresponds to a total uplift in the Carboniferous of approximately 4 km. The area was at the erosion level from Late Carboniferous onwards, as shown by deposition of the continental siliciclastic Paganzo Group in angular unconformity to the underlying basement (e.g. Buatois & Mangano 1996; Durand et al 1996).

Discussion and conclusions (back-docking of an island arc v. terrane accretion) The crust on which the sediments and volcanic rocks of the Famatina complex developed, including the Las Termas belt, is composed of Proterozoic to Early Cambrian siliciclastic detritus, prevalently of continental origin. Formation of this sequence is considered to have occurred during a time of quiescence at a passive continental margin (Jezek & Miller 1987). The material was supplied from a distant hinterland. During the Pampean Orogeny (about 550-515 Ma) the western passive continental margin of Gondwana turned to an active one (e.g. Pankhurst & Rapela 1998). From this point on, geological development of the western Gondwana margin can be discussed according

to two alternatives. On the one hand, the Famatina complex could have been amalgamated as an allochthonous terrane to the Gondwana continental margin; on the other hand, it may have formed at the active western continental margin of the eastern Sierras Pampeanas complex in the Palaeozoic by intracontinental rifting and back-arc docking of a parautochthonous island arc. Detailed structural, geochemical and geochronological data argue for the second possibility, particularly for the Famatina complex (Mannheim & Miller 1996; Rapela 2000). This model may be applicable to the entire southern Central Andes (Bock et al 2000; Lucassen et al 2000), but back-arc basin material of Famatinian age has not been recognized specifically in the sierras of San Luis or Cordoba (Whitmeyer & Simpson 2004). Steeply dipping subduction zones like those of the recent western Pacific Ocean are accompanied by intense arc volcanism and back-arc basins. Flat subduction zones like the recent one at 27-33° S in South America are characterized by translation of magmatism to the interior of the continent (e.g. Ramos et al 2002; the word 'Pampean' in that paper refers to the Neogene Sierras Pampeanas). If characteristics of recent subduction zones with different dipping angles are assigned to the Early Palaeozoic scenario, a number of conclusions can be drawn. From the Late Cambrian, steepening of the Proto-Pacific/Gondwana margin subduction zone led to marine transgression accompanied by simultaneous calc-alkaline intrusive and effusive magmatism in the Famatina magmatic arc. Increasing dip of the subducting Pacific plate initiated the formation of an extensional belt (Las Termas belt) between the eastern Sierras Pampeanas complex and the incipient volcanic island chain of the Famatina and Puna complexes. Crustal thinning, due to the extensional process, induced initial intrusion of mantle-derived basic to ultrabasic melts at about 515-510 Ma (meta-gabbronorite of Fiambala: DeBari 1994; Grissom et al 1998; Rapela et al 1999), followed by intermediate to acid igneous activity. Rocks comparable to the Fiambala meta-gabbronorite may be the source rocks for the Arenigian 'juvenile detritus' described from the Puna Basin (Bock et al 2000). Formation of the Famatina volcanic arc and the Las Termas back-arc basin are shortlived events compared to the continuous subduction of the Pacific crust beneath the eastern Sierras Pampeanas complex (Fig. 9). Accretion of the Famatina complex as an exotic (?) terrane would imply that almost all of the

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rocks forming the NW Argentine Gondwana margin in Palaeozoic times were part of such an allochthonous terrane. At the present state of investigation, evidence is lacking for distinct exotic (as opposed to Gondwanan) elements in the Famatina complex or other areas to the north that were affected by the Famatinian Orogeny (Lucassen et al 2000; Lucassen & Becchio 2003), though Whitmeyer & Simpson (2004) suggest the influence of Pacific microcontinent collisions (Precordillera and 'Chilenia') for the tectono-metamorphic evolution of the central Argentine sierras of San Luis and Cordoba. Moreover, at the southern Central Andes Pacific margin, the Precordillera of Mendoza and San Juan, possibly including the western Sierras Pampeanas (Pankhurst & Rapela 1998), is considered to be exotic. Individual analyses of detrital zircon cores will help to possibly verify or discredit such exotic components (Finney et al. 2003). Recent comparison of ages and geochemistry of MidOrdovician tuffs from the Precordillera and Sierra de Famatina may argue in favour of a close relationship of both terranes, at least since Mid-Ordovician times (Fanning et al. 2004). Currently, two-stage Nd model ages of the Famatinian meta-granitoids at 1.61 ± 0.02 Ga (la, n - 5, Hockenreiner et al. 2003) are constant and do not differ significantly from those of supposed basement xenoliths in the Precordillera, which range from 1.34 Ga to 1.69 Ga (Kay et al. 1996). The upper discordia intercept ages of 1.83 Ga and 1.27 Ga of Famatinian meta-granitoids, which refer to inherited zircons, may reflect varied provenance from orogens of the Brazilian shield or from the Arequipa block (Loewy et al 2004). The collision of an exotic terrane such as Chilenia and/or Cuyania with the eastern Sierras Pampeanas is not a requirement to explain the magmatic and deformation features. Reduction of velocity and dip of the subducting slab and/or acceleration of horizontal plate movements would be sufficient to close the back-arc basin. All these processes can be understood easily as local phenomena. This is underlined by the opinion of Sims et al (1998) and Lucassen et al (2000) who are doubtful of a great hiatus in the orogenic activity of the Pampean and Famatinian cycles, beyond the classic angular unconformity of the eastern Cordillera in Salta and Jujuy. On the contrary, they argue in favour of a more or less continuous and long-lasting process of orogeny compared with terrane accretion. The Famatinian tectono-thermal cycle ceased with the closure of the back-arc basin and backdocking of the Famatina complex to the eastern

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Sierras Pampeanas. Garnet growth in the TIPA shear zone at 402 ± 2 Ma is indicative of the age of the final deformation. These results point to final closure of the back-arc basin in the Early Devonian. This work is, in great part, based on the Dr.rer.nat. theses of K. Clemens, R. Mannheim, H. Neugebauer and M. Hockenreiner, carried out at the LudwigMaximilians-Universitat Miinchen, generally cited by corresponding publications. It was promoted by the co-operation programme between the LudwigMaximilians-Universitat Miinchen and the Universidad Nacional de Tucuman, especially with F. G. Acenolaza, A. J. Toselli and J. P. Lopez. Grants provided by the Deutsche Forschungsgemeinschaft and the scientific co-operation programme between the governments of Argentina and Germany are acknowledged gratefully. The authors are indebted to the reviewers G. Franz, R. J. Pankhurst and S. J. Whitmeyer for a great number of helpful comments. The paper forms part of IGCP programme 436 (Gondwana Pacific Margin).

References ACENOLAZA, F.G. & MILLER, H. 1982. Early Palaeozoic orogeny in southern South America. Precambrian Research, 17, 133-146. ACENOLAZA, F.G. & TOSELLI, A.J. 1976. Consideraciones estratigraficas y tectonicas sobre el Paleozoico inferior del Noroeste Argentine. Memorial del H. Congreso Latinoamericano de Geologia, Caracas, 1973, 2, Ministerio de Minas e Hidrocarburos, Direction de Geologia, Caracas, 755-764. ACENOLAZA, F.G. & TOSELLI, A.J. 1981. Geologia del Noroeste Argentino. Universidad Nacional de Tucuman, Facultad de Ciencias Naturales, Publication Especial 1287. ACENOLAZA, EG, MILLER, H. & TOSELLI, A.J. 1996. Cuadro general. Munchner Geologische Hefte, 19A, 5-8. ACENOLAZA, EG., MILLER, H. & TOSELLI, A.J. 2002. Proterozoic-Early Paleozoic evolution in western South America - a discussion. Tectonophysics, 354, 121-137. ASTINI, R.A., BENEDETTO, J.L. & VACCARI, N.E. 1995. The early Paleozoic evolution of the Argentine Precordillera as a Laurentia rifted, drifted, and collided terrane: A geodynamic model. Geological Society of America Bulletin, 107, 253-273. ASTINI, R.A., DAVILA, EM., RAPELA, C.W., PANKHURST, R.J. & FANNING, C.M. 2003. Ordovician back-arc clastic wedge in the Famatina Ranges: new ages and implications for reconstruction of the proto-Andean Gondwana margin. In: ALBANESE, G.L., BERESI, M.S. & PERALTA, S.H. (eds) Ordovician from the Andes. Proceedings of the 9th International Symposium on the Ordovician System. INSUGEO Serie Correlation Geologica 17, 375-380. Universidad Nacional de Tucuman, Argentina.

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BACHMANN, G, GRAUERT, B. & MILLER, H. 1986. Isotopic dating of polymetamorphic metasediments from Northwest Argentina. Zentralblatt fur Geologic und Palaontologie. Teil 1,1985,9/10, 1257-1268. BAHLBURG, H. 1990. The Ordovician basin in the Puna of NW Argentina and N Chile: geodynamic evolution from back-arc to foreland basin. Geotektonische Forschungen, 75,1-107. BOCK, B., BAHLBURG, H., WORNER, G. & ZIMMERMANN, U. 2000. Tracing crustal evolution in the Southern Central Andes from Late Precambrian to Permian with geochemical and Nd and Pb isotope data. Journal of Geology, 108, 515-535. BUATOIS, L.A. & MANGANO, M.G 1996. Sedimentacion lacustre postglacial en la Formacion Agua Colorada, Carbomfero Superior, del Sistema de Famatina. Munchner Geologische Hefte, 19A, 103-124. CAMINOS, R. 1979. Sierras Pampeanas Noroccidentales, Salta, Tucuman, Catamarca, La Rioja, San Juan. // Simposio de Geologia Regional Argentina, 1, Academia Nacional de Ciencias de Cordoba, 225-291. CLEMENS, K. & MILLER, H. 1996. Sedimentologia, proveniencia y position geotectonica de las sedimentitas del Precambrico y Paleozoico inferior del Sistema de Famatina. Munchner Geologische Hefte, 19A, 31-50. DALLA SALDA, L., CINGOLANI, C. & VARELA, R. 1992. Early Paleozoic orogenic belt of the Andes in southwestern South America: Result of Laurentia-Gondwana collision? Geology, 20, 617-620. DEBARI, S.M. 1994. Petrogenesis of the Fiambala intrusion, north-western Argentina, a deep crustal syntectonic pluton in a continental magmatic arc. Journal of Petrology, 35, 679-713. DURAND, F.R. & ACENOLAZA, F.G 1990. Caracteres biofaunisticos, paleoecologicos y paleogeograficos de la Formacion Puncoviscana (Precambrico Superior-Cambrico inferior) del noroeste Argentine. In: ACENOLAZA, F.G, MILLER, H. & TOSELLI, AJ. (eds) El Ciclo Pampeano en el Noroeste Argentino. Serie Correlation Geologica, 4, Tucuman, 71-112. DURAND, F.R., VERGEL M. & LECH, R.R. 1996. Las sedimentitas neopaleozoicas del Sistema de Famatina. Munchner Geologische Hefte, 19A, 77-95. FANNING, CM., PANKHURST, R.J., RAPELA, C.W., BALDO, E.G., CASQUET, C. & GALINDO, C. 2004. K-bentonites in the Argentine Precordillera contemporaneous with volcanism in the Famatinian arc. Journal of the Geological Society, London, 161, 747-756. FINNEY, S., GLEASON, J., GEHRELS, G, PERALTA, S. & ACENOLAZA, G. 2003. Early Gondwana connection for the Argentine Precordillera terrane. Earth and Planetary Science Letters, 205,349-359. GRISSOM, G.C. 1991. Empirical constraints on thermal processes in the deep crust of magmatic arcs: Sierra de Fiambala. Northwestern Argentina. PhD thesis, Stanford University, USA.

GRISSOM, G, DEBARI, S.M. & SNEE. L. 1998. Geology of the Sierra de Fiambala, northwestern Argentina: implications for Early Palaeozoic Andean tectonics. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 297-323. HOCKENREINER, M. 2003. Die TIPA-Scherzone (Unterdevon, NW-Argentinien): Geochronologie, Geochemie und Strukturgeologie. Munchner Geologische Hefte, A34, 1-98. HOCKENREINER, M., SOLLNER, M. & MILLER, H. 2003. Dating the TIPA shear zone: an early Devonian terrane boundary between the Famatinian and Pampean systems (NW Argentina). Journal of South American Earth Sciences, 16, 45-66. JEZEK, P. & MILLER, H. 1987. Petrology and facies analysis of turbiditic sedimentary rocks of the Puncoviscana trough (upper Precambrian-lower Cambrian) in the basement of the NW Argentine Andes. In: McKENZiE, G.D. (ed.) Gondwana Six: Structure, Tectonics and Geophysics. American Geophysical Union, Geophysical Monographs, 40, 287-293. KAY, S., ORELL, S. & ABRUZZI, J.M. 1996. Zircon and whole rock Nd-Pb isotopic evidence for a Grenville age and Laurentian origin for the basement of the Precordillera in Argentina. Journal of Geology, 104, 637-648. KNUVER, M. & MILLER, H. 1982. Rb-Sr-geochronology of the Sierra de Ancasti (Pampean Ranges, NW-Argentina). Actas 5. Congreso Latinoamericano de Geologia, 3. Buenos Aires, 457^71. LOEWY, S.L., CONNELLY, IN. & DALZIEL, LW.D. 2004. An orphaned basement block: The ArequipaAntofalla Basement of the central Andean margin of South America. Geological Society of America Bulletin, 116, 171-187. LOPEZ, IP. & TOSELLI, A.I 1993. La faja milonitica TIPA: Faldeo oriental del Sistema de Famatina. XII Congreso Geologico Argentino y II Congreso de Exploracion de Hidrocarburos, Mendoza, Actas 3, 39-42. LORK, A., GRAUERT, B., KRAMM, U & MILLER, H. 1991. U-Pb investigations of monazite and polyphase zircon: Implications for age and petrogenesis of trondhjemites of the southern Cordillera Oriental, NW-Argentina. 6. Congreso Geologico Chileno, resumenes expandidos, Santiago, Servicio Nacional de Geologia, Mineria, 398-402. LUCASSEN, F. & BECCHIO, R. 2003. Timing of highgrade metamorphism: Early Palaeozoic U-Pb formation ages of titanite indicate long-standing high-T conditions at the western margin of Gondwana (Argentina, 26-29°S). Journal of Metamorphic Geology, 21, 649-662. LUCASSEN, R., BECCHIO, R., WILKE, H.G., FRANZ, E, THIRLWALL, M.F., VIRAMONTE, I & WEMMER, K. 2000. Proterozoic-Paleozoic development of the basement of the Central Andes (18-26°S) - a mobile belt of the South American craton. Journal of South American Earth Sciences, 13, 697-715.

FAMATINA COMPLEX GEOTECTONIC SETTING MANGANO, F.G. & BUATOIS, L.A. 1996. Estratigrafia, sedimentologia y evolution paleoambiental de la Formation Suri en la subcuenca de Chaschuil, Ordovicico del Sistema de Famatina. Munchner Geologische Hefte, 19A, 51-76. MANNHEIM, R. & MILLER, H. 1996. Las rocas volcanicas y subvolcanicas eopaleozoicas del Sistema de Famatina. Munchner Geologische Hefte, 19A, 159-186. MILLER, H., TOSELLI, A.J., Rossi DE TOSELLI, J. & ACENOLAZA, F.G. 1994. Regional and geochronological development of the metamorphic basement in Northwest Argentina. Zentralblatt far Geologic und Palaontologie, 1993 (1/2), 263-273. MIRO, R.C. 2004. Magmatismo calco-alcalino en la Sierra Norte de Cordoba: su extension temporal. In: ACENOLAZA, F.G, ACENOLAZA, G.F., HUNICKEN, M, TOSELLI. AJ. (eds) Simposio Bodenbender. INSUGEO, Miscellanea, 13, 43-44. NEUGEBAUER, H. 1996. Die Mylonite von Fiambald Strukturgeologische und petrographische Untersuchungen an der Ostgrenze des Famatina-Systems, Sierra de Fiambald, NW-Argentinien. Dr.rer.nat. thesis, Ludwig-Maximilians-Universitat Miinchen, Germany. NEUGEBAUER, H. & MILLER, H. 1996. La naturaleza tectonica del limite oriental del Sistema de Famatina en la Sierra de Fiambala. Munchner Geologische Hefte, 19A, 325-341. PANKHURST, RJ. & RAPELA, C.W. 1998. The protoAndean margin of Gondwana: an introduction. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 1-9. PANKHURST, R.J., RAPELA, C.W., SAAVEDRA, J., BALDO, E., DAHLQUIST, I, PASCUA, I. & FANNING, CM. 1998. The Famatinian magmatic arc in the central Sierras Pampeanas: an Early to Mid-Ordovician continental arc on the Gondwana margin. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The ProtoAndean Margin of Gondwana. Geological Society, London, Special Publications, 142, 343-367. PANKHURST, R.J., RAPELA, C.W. & FANNING, C.M. 2000. Age and origin of coeval TTG, I- and S-type granites in the Famatinian belt of NW Argentina. Transactions of the Royal Society of Edinburgh, Earth Sciences, 91, 151-168. RAMOS, V.A., JORDAN, T.E., ALLMENDINGER, R.W., MPODOZIS, C, KAY, S.M., CORTES, J.M. & PALMA, M. 1986. Paleozoic terranes of the central Argentine-Chilean Andes. Tectonics, 5, 855-880. RAMOS, V.A., DALLMEYER, R.D. & VUJOVICH, G.I. 1998. Time constraints on the Early Palaeozoic docking of the Precordillera, central Argentina. In: PANKHURST, R J. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 143-158. RAMOS, V.A , CRISTALLINI, E.O. & PEREZ, DJ. 2002. The Pampean flat-slab of the Central Andes. Journal of South American Earth Sciences, 15, 59-78. RAPELA, C.W. 2000. El ambiente geotectonico del

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Ordovicico de la region del Famatina. Revista de la Asociacion Geologica Argentina, 55, 134-136. RAPELA, C.W., PANKHURST, R J., CASQUET, C., BALDO, E., SAAVEDRA, J. & GALINDO, C. 19980. Early evolution of the Proto-Andean margin of South America. Geology, 26, 707-710. RAPELA, C.W, PANKHURST, R J., CASQUET, C., BALDO, E., SAAVEDRA, J., GALINDO, C. & FANNING, M. 1998ft. The Pampean orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de Cordoba. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142,181-217. RAPELA, C.W, COIRA, B., TOSELLI, A J. & LLAMBIAS, EJ. 1999. Sistema Famatiniano de las Sierras Pampeanas y magmatismo eopaleozoico de las Sierras Pampeanas, de la Cordillera Oriental y Puna. In: CAMINOS, R. (ed.) Geologia Argentina, SEGEMAR, Anales, 29. Buenos Aires, 145-158. RAPELA, C.W, PANKHURST, R.J., BALDO, E., CASQUET, C., GALINDO, C., FANNING, C.M. & SAAVEDRA, J. 2001. Ordovician metamorphism in the Sierras Pampeanas: New U-Pb shrimp ages in CentralEast Valle Fertil and the Velasco Batholith. /// South American Symposium on Isotope Geology, Extended Abstracts Volume (CD). Sociedad Geologica de Chile, Santiago, Chile, 616-619. SAAVEDRA, X, TOSELLI, A., Rossi, I, PELLITERO, E. & DURAND, F. 1998. The Early Palaeozoic magmatic record of the Famatina System: a review. In: PANKHURST, R J. & RAPELA, C.W. (eds) The ProtoAndean Margin of Gondwana. Geological Society, London, Special Publications, 142, 283-295. SCHWARTZ, J.-J. & GROMET, L.P. 2004. Provenance of a late Proterozoic-early Cambrian basin, Sierras de Cordoba, Argentina. Precambrian Research, 129, 1-21. SCHWARTZ, J.J., GROMET, PL. & MIRO, R. 2003. Neoproterozoic-Early Cambrian calc-alkaline magmatism in the eastern Sierras Pampeanas, Argentina: U-Pb zircon and isotopic constraints. Geological Society of America Abstracts with Program, Geoscience Horizons (Seattle), 35(6), 345. SIMS, J.P., IRELAND, T.R., CAMACHO, A., ET AL. 1998. U-Pb, Th-Pb and Ar-Ar geochronology from the southern Sierras Pampeanas, Argentina: implications for the Palaeozoic tectonic evolution of the western Gondawana margin. In: PANKHURST, RJ. & RAPELA, C.W (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 259-281. SOLLNER, F, BRODTKORB, M.K. DE, MILLER, H., PEZZUTTI, N.E. & FERNANDEZ, R.R. 20000. U-Pb zircon ages of metavolcanic rocks from the Sierra de San Luis, Argentina. Revista de la Asociacion Geologica Argentina, 55, 15-22. SOLLNER, F, MILLER, H. & HERVE, M. 20006. An Early Cambrian granodiorite age from the preAndean basement of Tierra del Fuego (Chile): the missing link between South America and Antarctica? Journal of South American Earth Sciences, 13, 163-177.

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SOLLNER, K, HOCKENREINER, M. & MlLLER, H. 2001.

Constraints on the ages of Famatinian igneous intrusions and subsequent deformation in the Sierra de Fiambala (Catamarca/NW-Argentina). - ///. South American Symposium on Isotope Geology, abbreviated abstracts, Comunicaciones. Departamento de Geologia, Universidad de Chile, Santiago, 52,167. TORTELLO, F., ESTEBAN, S.B. & LECH, R.R. 1996. Paleontologia del Sistema de Famatina. Miinchner Geologische Hefte, 19A, 125-136. TOSELLI, A.J., SAAVEDRA, X & Rossi DE TOSELLI, XN. 1996. Interpretation geotectonica del magmatismo del Sistema de Famatina. Miinchner Geologische Hefte, 19A, 283-291. WHITMEYER, S.X & SIMPSON, C. 2003. High strain-rate deformation fabrics characterize a kilometersthick Paleozoic fault zone in the Eastern Sierras Pampeanas, central Argentina. Journal of Structural Geology, 25, 909-922.

WHITMEYER, S.X & SIMPSON, C. 2004. Regional deformation of the Sierra de San Luis, Argentina: implications for the Paleozoic development of western Gondwana. Tectonics, 23, TC1005, doi: 10.1029/2003TC001642. WILLNER, A.P. 1990. Division tectonometamorfica del basamento del noroeste Argentine. In: ACENOLAZA, EG., MILLER, H. & TOSELLI, AJ. (eds) El Ciclo Pampeano en el noroeste Argentino. Serie Correlation Geologica, Tucuman, 4,113-159. WILLNER, A.P, LOTTNER, U.S. & MILLER, H. 1987. Early Paleozoic structural development in the NW Argentine basement of the Andes and its implication for geodynamic reconstructions. In: McKENZiE, G.D. (ed.) Gondwana Six: Structure, Tectonics, and Geophysics, American Geophysical Union, Geophysical Monographs, 40, 229-239.

The early Palaeozoic Orogen in the Central Andes: a non-collisional orogen comparable to the Cenozoic high plateau? FRIEDRICH LUCASSEN1 & GERHARD FRANZ2 GeoForschungsZentrum Potsdam, Telegrafenberg, D-14473 Potsdam, Germany and Freie Universitat Berlin, FB Geowissenschaften, Malteserstr. 74-100, D-12249 Berlin, Germany 2 Technische Universitat Berlin, Angewandte Geowissenschaften, D-10623 Berlin, Germany

l

Abstract: The subduction orogeny of the Central Andes, which created the Cenozoic Altiplano-Puna high plateau, shares many geological features with the early Palaeozoic Orogen at the western margin of South America. The presently available datasets for both orogens are compared. The similarities are a large-scale high temperature metamorphism, which was active in the Palaeozoic Orogen over a geological long period of time in the order of 100 Ma and which is active now in the crust of the Cenozoic plateau. It produced abundant granitoid melts from the crust during the Palaeozoic as well as during the Andean Orogen. The main contribution to granitoid magmatism is recycling of felsic crustal material with only minor additions from the mantle. Transport of deep parts of the crust into the erosion level did not occur in both orogens and, in both orogens, large-scale nappe tectonics typical for collision orogens are absent. Based on the similarities of the two orogens it is argued that the early Palaeozoic Orogen is a non-collisional orogen. Indications for terrane accretion are absent in the development of the high-grade metamorphic and igneous basement. The early Palaeozoic Orogen is an analogue for the presently active continental margin and, thus, allows the extrapolation of features which cannot be observed in the Andean Orogen.

The Cenozoic Altiplano and Puna of the Central Andes at the Pacific margin of South America form the world's second largest high plateau (Fig. 1). In contrast to the larger Tibetan plateau, which evolved in Central Asia during continent-continent collision, Altiplano and Puna formed at a non-collisional active continental margin. The long history of subduction and magmatism in this area began in the early Palaeozoic or earlier, with phases of enhanced magmatic activity and widespread high-T metamorphism during the early Palaeozoic. The socalled Andean cycle' started in the Jurassic and was separated from the early Palaeozoic evolution by a passive margin episode in the Devonian and early Carboniferous (Bahlburg & Herve 1997; Augustsson & Bahlburg 2003). The building of orogens and large-scale crustal thickening occurred twice in the Phanerozoic evolution of the Central Andes: during the early Palaeozoic and since the Cenozoic. The formation of the early Palaeozoic Orogen was discussed controversially as a mosaic of locally derived and allochthonous terranes (Fig. 2; e.g. Ramos et al 1986; Ramos 1995; Ramos & Keppie 1999), which accreted during various collisional events, or as 'ensialic' (Damm et al 1990) reworking Proterozoic crust of the

western leading edge of the South American craton (e.g. Lucassen et al. 2000; Acenolaza et al 2002). The general outline of the tectonic and magmatic evolution of the Cenozoic Andes and the physical properties of the crust beneath the plateau, such as seismic velocities, electrical conductivity, density and the thermal regime in the crust, are well constrained (see references below). This paper compares the thermal structure of the non-collisional Cenozoic Orogen derived from geophysical data and type of magmatism with the thermal evolution of the exhumed early Palaeozoic Orogen derived from type and age of metamorphism and magmatism. The radiogenic isotope signatures of sedimentary, metamorphic and magmatic rocks from the early Palaeozoic onwards give additional constraints on the process of crust formation in the Central Andes and the possible accretion of compositionally exotic crustal fragments (terranes).

The Palaeozoic Orogen of the Central Andes In NW Argentina, in the Puna plateau and its southern continuation, metamorphic, magmatic

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 257-273. 0305-8719/$15.00 © The Geological Society of London 2005.

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and sedimentary rocks of the early Palaeozoic Orogen comprise large areas at surface (Fig. 3). In the plateau area of northern Chile erosion is low due to the arid to semi-arid climate and the pre-Cenozoic surface is preserved but covered in many areas by Cenozoic sedimentary and volcanic rocks. The uniform age structure, peak-

metamorphic conditions, chemical and isotopic composition from the basement rocks in Argentina and the scattered outcrops in northern Chile allow the interpolation between isolated outcrops of Palaeozoic rocks in areas with younger cover-rocks. The early Palaeozoic Orogen was characterized by a number of features. The orogen formation was preceded by continental extension and the formation of a large north-southtrending marine basin in the Neoproterozoic and early Cambrian centred mainly east of what later formed the orogen and magmatic arc. The basin with its deposits of the Puncoviscana Formation and equivalents deepened towards the south (Jezek et al 1985; Durand & Acenolaza 1990; Omarini et al 1999). It led possibly to the formation of ocean floor in the southern part (Baldo et al 1996; Rapela et al 1998), but its main parts were still deposited on continental crust. The extensional tectonic setting in latest Proterozoic-early Cambrian occurred during fragmentation of the cratonic areas by PanAfrican-Braziliano movements (Cordani et al 2003). The metamorphic rocks comprise mainly compositionally uniform metasediments, which are commonly equivalents of Neoproterozoic to Cambrian siliciclastic sedimentary rocks and subordinate carbonate in NW Argentina (Willner et al 1985; Lucassen et al 2001; Schwartz & Gromet 2004). Rocks with mafic composition are subordinate and comprise «5% of the outcrop area. The peak metamorphic conditions of regional metamorphism in the deepest exposed sections of the Palaeozoic Orogen between 18° S and 34° S are characterized by uniform felsic migmatite and gneisses, the latter commonly representing deformed migmatite (T c. 650-800 °C, P c. 0.4-0.7 GPa; e.g. Baldo et al 1996; Pankhurst & Rapela 1998; Becchio etal 1999; Lucassen etal 2000; Worner et al 2000; Buttner et al 2005). The metamorphic pathway is simple (see the above references). Indicators of the prograde metamorphic evolution are absent generally in the mineral Fig. 1. (a) The digital elevation model of the Puna-Altiplano plateau that hosts the Recent magmatic arc. The Jurassic-lower Cretaceous magmatic arc is located in the Chilean Coastal Cordillera, the mainly late Cretaceous intraplate rift-related magmatism in the present back-arc and eastern parts of the plateau. The white rectangle shows the location of (b). (b) Distribution of Cenozoic ignimbrites, related calderas and andesite volcanoes on the plateau between 20° S and 27° S. Figure modified from Babeyko et al (2002).

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Fig. 2. Distribution of early Palaeozoic and Proterozoic metamorphic rocks between 16° S and 34° S. Proterozoic ages of metamorphism are restricted to Peru and NW Bolivia but inherited Proterozoic ages are common in zircons from early Palaeozoic (and younger) metamorphic and magmatic rocks. Terrane boundaries and names are from Ramos (1985). Figure modified from Lucassen et al (2000). Am, Arequipa Massif; Br, Berenguela; U, Uyarani.

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Fig. 3. Distribution of early Palaeozoic magmatic and metamorphic rocks, late Proterozoic to early Cambrian sedimentary rocks in a section of NW Argentina (22-33° S map modified after SEGEMAR 1997). The eastern extension of the early Palaeozoic Mobile Belt (c. 18-34° S; Lucassen et al 2000) is speculative. The boundary of the Argentine Precordillera (Astini et al 1995) separates the possible exotic terrane from the western Sierras Pampeanas. The late Proterozoic-Cambrian sedimentary rocks of the Puncoviscana Formation belong to the marine basin which formed before the Cambrian-Silurian Orogeny. They are the protolith of many low- to high-grade metasedimentary rocks at least in the eastern parts of the basement (Salta-Cordoba). Inset shows the approximate position and continuation of the area in South America. paragenesis or mineral zoning of the high-T rocks. Deformation at high temperature is common and transition from migmatic textures

to gneisses occurs at various scales (millimetre to kilometre), leaving the mineral paragenesis unchanged. Changes in the mineral paragenesis

EARLY PALAEOZOIC ANDEAN OROGENESIS or core-rim chemical zoning of minerals occurred only during recrystallization and retrogression from peak temperature. Metamorphic conditions indicate a rather uniform deepest exposure level of the orogen with c. 15-25 km crust removed over a large area of >1000 km parallel to the margin, with considerable east-west width of more than 300-400 km (Figs 2,3). Crustal thickness in the Palaeozoic Orogen must have been 50 km or more considering a post-erosion crustal thickness of c. 35 km. Exposures of high-P rocks such as blueschists and eclogites typical of accretionary wedges and collision zones or high-pressure granulites from the root of the orogen are unknown. To the authors' knowledge there are only two descriptions of high-pressure rocks, one from the Chilean Precordillera at Limon Verde (Lucassen et al 1999«), which is of Permian age, and another one from the Argentine Precordillera (Casquet et al 2001), who determined for the Ordovician peak metamorphic conditions a pressure range of 1.0-1.25 GPa at 550-650 °C (taken from their fig. 3; 1.3 ± 0.1 GPa in text). These conditions (as well as those of Limon Verde) are still in the plagioclase stability field and not typical eclogite facies. They can be explained also as the conditions at the base of a crust of normal thickness. Magmatic rocks occur in the same areas as the high-T metamorphism (Fig. 3). The magmatic rocks are prevailing granitoid intrusions, and mafic intrusions are rare (e.g. Damm et al, 1994; Coira et al 1999; Pankurst & Rapela 1998; Pankhurst et al 2000). Scarce early Palaeozoic mafic intrusions were found in the plateau area and west of it and their compositions resemble rocks from arc magmatism (Damm et al 1990; Coira et al 1999; Zimmermann et al 2003; Kleine et al 2004). The chemical composition of the granites indicates various depths in the lower to mid crust of melt generation and hybridization (e.g. Pankhurst et al 2000), but all melts show considerable contribution of crustal material (see below) or are crustal melts related to the metasediments compositionally. Isotopic ages of high-T metamorphism and crust-derived magmatism indicate longstanding high-T conditions in the crust in the Cambrian and Ordovician. High-T metamorphism between c. 530 Ma and 500 Ma occurred throughout the whole orogen at distant locations at 23° S, 26° S, 29° S and 32° S (U-Pb data and Sm-Nd internal isochrons and additional references in: Damm et al 1990; Pankhurst & Rapela 1998; Lucassen et al 2000; Lucassen & Becchio 2003; Hoeckenreiner et al 2003; Biittner et al 2005). The same type of

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metamorphism is widespread (18-34° S) also between c. 470 Ma and 440 Ma. The Ordovician ages overlap spatially with the Cambrian ages. High-T metamorphism persisted locally until Silurian/early Devonian (above references; Varela et al 2003). Intrusive rocks show U-Pb and Rb-Sr ages between c. 560 Ma and 400 Ma (data and additional references in: Damm et al 1994; Pankhurst & Rapela 1998; Stuart-Smith et al 1999; Pankhurst et al 2000). The most voluminous magmatism was in the lower to midOrdovician between c. 490 Ma and 460 Ma, but older (Cambrian) and younger (Silurian/lower Devonian) intrusions are important locally (above references; Varela et al 2003). The present dataset shows no clear regional trend in the age distribution of Cambrian and Ordovician magmatism and high-T metamorphism. High-T conditions in the crust at c. 29° S could have persisted in the Silurian (Lucassen & Becchio 2003), whereas the onset of erosion (Moya 1988; Bock et al 2000; Egenhoff & Lucassen 2003) and cooling in the metamorphic rocks (Lucassen et al 2000) in northern Chile and the Argentine Puna had occurred already by the Ordovician. Early Palaeozoic high-T metamorphic rocks are overlain by Devonian, Carboniferous or lower Permian sedimentary rocks at various locations, indicating the minimum ages of final uplift and erosion (Becchio et al 1999; Lucassen et al 2000; Worner et al 2000). The age of the protoliths of the early Palaeozoic rocks is constrained by Nd model ages of the Palaeozoic rocks (Fig. 4; most samples plot between 2.0 Ga and 1.6 Ga) and inherited U-Pb ages from zircons in metamorphic and magmatic rocks (Fig. 4). The Nd model ages are in good agreement with the age of a major growth cycle of the Brazilian Shield mainly west of the palaeo-Proterozoic and Archaean cores (Trans-Amazonian Orogeny, c. 2.2-1.8 Ga, Nd model ages; Cordani et al 2000; Sato & Siga 2002). Inherited U-Pb ages show two clusters (Fig. 4). The cluster between 1.2 Ga and 1.0 Ga has the age of the Sunsas Orogeny of the Brazilian Shield (e.g. Cordani etal 2000), which is also documented in Sm-Nd isochrons and K-Ar ages (Worner et al 2000; Varela et al 2003); the second, between 2.0 Ga and 1.8 Ga, indicates the Trans-Amazonian crustal growth. Both events are prominent in zircon age patterns of other locations inside the craton (Valeriano et al 2004) and Amazon (Rino et al 2004) and Orinoco (Goldstein et al 1997) river sands. Inherited palaeo-Proterozoic and Archaean ages (>2.4 Ga), which contribute to the age spectra of the latter sample sites, are absent or rare in the Palaeozoic Orogen.

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Fig. 4. Nd-model ages and ranges of inherited U-Pb ages on zircons. Data sources: Early Palaeozoic magmatic and metamorphic rocks 18-33° S (Damm et al 1994; Pankhurst et al 1998; Rapela et al 1998; Lucassen et al 2001; Loewy et al 2004; authors' unpublished data); Neoproterozoic to Ordovician sedimentary rocks 20-26° S (Bock et al 2000; Lucassen et al 2001; Egenhoff & Lucassen 2003; Zimmermann & Bahlburg 2003; Schwartz & Gromet 2004); Cenozoic ignimbrite (Francis et al 1989; Ort et al 1996; Lindsay et al 2001). All Sm-Nd data are calculated using a single-stage evolution model (Goldstein et al 1984).

The Nd-Sr-Pb isotopic signatures of the Palaeozoic metamorphic (Pankhurst et al 1998; Rapela et al 1998; Lucassen et al 2001; 20020), magmatic (e.g. Damm etal 1994; Pankhurst etal 1998; 2000; Rapela et al 1998; Lucassen et al 2001; 20020) and pre-, syn- and post-orogenic sedimentary rocks (Bock et al 2000; Lucassen et al 2001; Egenhoff & Lucassen 2003; Zimmermann & Bahlburg 2003) are indicative for

variations in the crustal make up. Nd and Sr isotope ratios (Fig. 5a) are consistent with recycling of a rather uniform Proterozoic (Fig. 4) crust or mixing material from a depleted sub-arc mantle source with such crust (Damm etal 1990; Lucassen et al 2001; Zimmermann & Bahlburg 2003; Kleine et al. 2004). Juvenile or weakly crust-contaminated rocks are rare and of minor volume. Uranogenic Pb isotope compositions

Fig. 5. Sr-Nd isotope systematics of early Palaeozoic rocks and magmatic rocks of the Andean cycle, (a) The initial isotope compositions of the various felsic early Palaeozoic metamorphic, magmatic and sedimentary rocks (c. 27-20° S) cover similar compositional ranges. In the Ordovician granitoids, the influence of the depleted sub-arc mantle source becomes prominent in some intrusions (e.g. Damm et al 1990; 1994; Kleine et al 2004). The early Palaeozoic mafic magmatic rocks and amphibolites represent compositionally such a depleted mantle source. In comparison to the mafic magmatic rocks, the amphibolites show considerable scatter in their Sr isotope ratios due to metasomatic effects during metamorphism. Data sources as in Figure 4 and Kleine et al (2004). (b) The isotopic composition of Andean magmatic rocks. 'Lower crustal xenolith' represents the Palaeozoic lower crust. The field of 'metamorphic rocks' and the average composition of the metamorphic rocks indicate the connection between the composition of the felsic crust and the depleted sub-arc mantle (weakly or uncontaminated Jurassic-lower Cretaceous rocks) in the hybrid Cenozoic andesite and Cenozoic ignimbrite, the latter containing large proportions of early Palaeozoic crust. The deeper Andean crust comprises early Palaeozoic high-grade metamorphic rocks depleted in Rb (and U) by early Palaeozoic metamorphism and melt extraction, represented by felsic lower crustal xenoliths and initial isotope ratios of the early Palaeozoic rocks (a). The pattern of mixing of crustal and mantle material is identical in the early Palaeozoic and Cenozoic arcs. Data sources as in Figure 4 and Jurassic to lower Cretaceous: Rogers & Hawkesworth (1989); Lucassen et al (20026); andesite: Trumbull et al (1999 and references therein); lower crust xenoliths: Lucassen et al (19996).

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Fig. 6. Uranogenic Pb isotopes of the early Palaeozoic and Cenozoic rocks, (a) Pb isotope ratios from leached feldspar of magmatic and metamorphic rocks indicate an increasing homogenization of the Pb isotope composition with ongoing metamorphism and melt extraction. Late Palaeozoic granitoids plot, with the exception of a single pluton within the field of the metamorphic rocks. Data sources see Figure 5a; Argentine Precordillera xenoliths: Kay etal (1996) metamorphic rocks 27-30° S. (b) The Pb isotope composition of Jurassic-lower Cretaceous rocks represents a mixture between the composition of the Jurassic MORE mantle and Palaeozoic crust (Lucassen etal. 2002ft). The Pb isotope ratios of the Cenozoic magmatic rocks (south of 21° S) are dominated by the compositional range of the early Palaeozoic crust. The subdivision of the early Palaeozoic crust by metamorphic grade seen in the Sr isotope composition is also present in the uranogenic Pb isotope. The average composition of feldspar of Palaeozoic metamorphic rocks resembles the composition of lower crustal xenoliths, which are U depleted. The average whole-rock composition is a proxy for the local crust not depleted in U. Data sources as in Figure 5b. Compilation of average Pacific MORE composition and time correction for in situ decay: Lucassen et al (2002ft) average composition of early Palaeozoic whole rocks: Lucassen etal (2001). NHRL, (Northern Hemisphere Reference Line) from Hart (1984).

EARLY PALAEOZOIC ANDEAN OROGENESIS

show a progressive homogenization of the crust during early Palaeozoic high-grade metamorphism and crustal melting (Fig. 6a; Lucassen et al 20020). The variations (e.g. of the Pb isotope ratios) between rock units of different ages can be explained by element mobility (e.g. of U depletion by metamorphism and melt extraction) followed by long-term separation of the depleted crust during the long evolution of the protoliths of the early Palaeozoic crust. From the point of view of this paper, the orogen formed at a continental margin without continent-continent collision. In most palaeogeographical reconstructions (e.g. Piper 2000) and in geodynamic interpretations of the section of the margin discussed here (e.g. Astini et al. 1995; Acenolaza et al 2002), the western edge of South America was a continental margin in the late Proterozoic-early Palaeozoic. The Palaeozoic Orogen north of the Argentine Precordillera has been interpreted as a coherent mobile belt at the western leading edge of South America (Acenolaza & Miller 1982; Damm et al 1994; Bock et al 2000; Lucassen et al 2000) in contrast to interpretations of the Palaeozoic basement as a mosaic of accreted terranes (Fig. 2; e.g. Ramos et al 1986; Ramos 1995; Ramos & Keppie 1999). There are key arguments against Neoproterozoic and Palaeozoic terrane accretion. (1) The metamorphic and magmatic history in the Cambrian and Ordovician of the whole area is similar. (2) The isotopic composition of Neoproterozoic and early Palaeozoic rocks does not show any exotic crustal fragments but, in the case of the uranogenic Pb isotopes, an increasing homogenization of older heterogeneities with ongoing metamorphism and magmatism. (3) Suture zones marked by ophiolites, high pressure-low temperature metamorphism or accretionary wedges - all considered as indicators of continental accretion and collision (e.g. O'Brien 2000) - have not been found in the Palaeozoic Orogen north of the Argentine Precordillera and suspect mafic rocks are related to magmatic arc activity instead of oceanic rifting and obduction of oceanic crust. (4) Large-scale stacking of crustal units is absent. (5) The age structure of the protoliths as revealed by the Nd isotope composition and inherited U-Pb ages fits perfectly the age pattern of the western parts of the South American craton. The early Palaeozoic metamorphic and magmatic rocks east of the Argentine Precordillera show identical features as the basement further north, except for MOR-like basalts of ocean floor. However, this potential ocean floor is from the NeoproterozoicCambrian basin (Baldo et al 1996: Ranela et al

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1998; Pankhurst et al 1998) and, therefore, forms part of the early Palaeozoic Mobile Belt. The possible early Palaeozoic accretion of the Argentine Precordillera, which is believed widely to be a Laurentia-derived exotic terrane based on its stratigraphic and faunal record (e.g. Astini et al 1995; Ramos & Keppie 1999), cannot be discerned in the P-T-t record of the adjacent high-grade metamorphic basement of the mobile belt. U-Pb ages of high-T metamorphism are c. 530 Ma, 470 Ma, 440 Ma, 420 Ma and an upper intercept age is 1.2 Ga (28-29° S; 68°10'-68°40' W; Lucassen & Becchio 2003). If the Argentine Precordillera accreted in the early Ordovician (e.g. Astini & Davila 2004), it did not influence the existing high-T conditions in the crust and the timing of the accretion cannot be pinpointed in the age pattern of the high-T rocks east of the Precordillera. Petrological studies on the metamorphic evolution, age structure and isotopic composition of the Argentine Precordillera itself are still scarce and existing age data (c. 1.2 Ga and 460 Ma, Vujovich & Kay 1998 and references therein; Casquet et al 2001) and Pb isotope composition (Kay et al 1996) are not distinctive. The same ages and isotope compositions (Fig. 6a) are present in the early Palaeozoic Mobile Belt (Lucassen et al 20020; Lucassen & Becchio 2003). Considering a non-collisional history of the early Palaeozoic Central Andes, a Cordillerantype orogeny is assumed here, at least for the area north of the Argentine Precordillera. The Cenozoic Central Andes are a prototype of this orogeny on continental crust.

The Cenozoic high plateau of the Central Andes The Cenozoic plateau (Fig. 1; e.g. Allmendinger et al 1997; Lamb & Hoke 1997) formed after a prolonged eastward migration of magmatic arcs during the Mesozoic (e.g. Rogers & Hawkesworth 1989; Scheuber et al 1994). The plateau building is preceded in the late Cretaceous-early Tertiary by north-south striking continental extension or a failed rift, which was located in the present eastern parts of the plateau and the eastern foreland of the Andes (Fig. 1; e.g. Galliski & Viramonte 1989; Viramonte et al 1999; Sempere et al 2002). The non-collisional setting, which is assumed for the Palaeozoic Orogen is of course obvious for the Andean Orogen. Deep sections of the Mesozoic arcs or the plateau are not exposed and all information about the properties of the crust is inferred from the eeoloeical historv of

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the region, geophysical imaging and magmatic rocks of the active arc. Erosion is low and tectonic transport of deep and deeply buried crust towards the surface, which is common in continent-continent collision zones (e.g. O'Brien 2000; Liou et al 2004, and references therein) has not been observed. Blueschist-eclogite or high pressure granulite rocks are absent, similar to the Palaeozoic Orogen. Wherever the cover by Cenozoic sedimentary rocks and volcanic rocks allows, it shows a largely preserved preCenozoic surface. This indicates a tectonically rather stable condition except for the eastern fold-and-thrust belt in southern Bolivia. The crustal make-up in the plateau area is still dominated by crust of the early Palaeozoic Orogen, considering that substantial crustal reorganization did not occur until the Cenozoic. A passive margin evolution in the Devonian to early Carboniferous (Bahlburg & Herve 1997) was followed by the renewed onset of subduction in the late Carboniferous (e.g. Brown 1991) and several episodes of subduction since then (Berg & Baumann 1985; Rogers & Hawkesworth 1989; Scheuber et al 1994). Average sedimentation in late Palaeozoic and Mesozoic times was low and higher sediment volumes are confined mainly to extensional structures (e.g. Bahlburg & Breitkreuz 1991; Suarez & Bell 1992; Breitkreuz & Zeil 1994; Prinz et al. 1994; Charrier & Reutter 1994). Substantial juvenile magmatic additions to the crust are restricted to the Jurassic-lower Cretaceous magmatic arc of the coastal area (e.g. Rogers & Hawkesworth 1989; Lucassen et al 1996) and are considered to be small in the Cenozoic arc (e.g. Francis & Hawkesworth 1994). The metamorphic conditions of the Andean metamorphism cannot be determined directly. The interpretation of surface heat flow data reveals high temperature in the crust (Springer 1999), but is ambiguous in zones of active magmatism with magma chambers near the surface. However, there are independent lines of evidence for the thermal structure of the plateau from geophysical data (seismic velocities and electrical conductivity) and the young (ignimbrite) magmatism. Seismic velocities show a pronounced attenuation in a mid-crustal low velocity layer (Fig. 7b; e.g. Wigger et al 1994; Chmielowski et al 1999; Giese et al 1999; Schmitz et al 1999; Swenson et al 2000; Zandt et al 2003). High electrical conductivity was detected in several areas in the mid-crust of the plateau (Fig. 7b; Brasse et al 2002). Both features were interpreted as the effect of a partially molten mid-crust (e.g. Brasse et al 2002; Zandt et al 2003). The upper

boundary of the partially molten zone lies at a depth of c. 20 km. Large-scale rhyodacitic ignimbrite volcanism is distributed widely on the plateau. The magmas erupted from chambers in the uppermost crust forming large caldera structures after discharging (Fig. 1; de Silva 1989; Francis et al 1989; Lindsay et al 2001). The large-scale ignimbrites represent hybrid melts containing >50-70% or even more felsic early Palaeozoic crust (Fig. 5b; e.g. Francis et al 1989; Ort et al 1996; Lindsay et al 2001). The andesite rocks of the Central Andes and their genetically related small-volume ignimbrites were formed from hybrid magmas with c. 20-30% contribution from early Palaeozoic crust (Fig. 5b; e.g. Harmon et al 1984; Hildreth & Moorbath 1988; Worner et al 1994; Trumbull et al 1999; Siebel et al 2001). Magmatic rock compositions typical of deepseated, garnet-bearing mafic rocks in their residuum are rare. Formation of crustal melts and hybridization of mantle melts takes place in the middle part of the 50-70 km thick crust rather than in the lower crust, where garnet is more abundant. Geophysical data and the composition of felsic ignimbrite and hybrid andesite imply a presently ongoing high-temperature regional metamorphism above the wet granite solidus at minimum pressures of approximately 5-6 kbar. The surface expression of the Cenozoic anatexis is the large volume ignimbrite province (Fig. 1; e.g. de Silva 1989; Francis et al 1989; Ort et al 1996; Lindsay et al 2001). Migmatic rocks can be expected as a dominant rock type in the middle crust and a large amount of granitoid intrusions in the upper crust. These granitoids are either the magma chambers of the ignimbrite or diapirs related to the anatexis. P-T estimation for the Cenozoic plateau agrees quite well with the P-T data from the Palaeozoic high-T rocks. Geophysical data and Cenozoic magma compositions have important implications also for the interpretation of a possible compositional stratification of the crust. Early Palaeozoic rocks should dominate at least in the upper crust, as inferred from the geological evolution. Velocity distribution from active and passive seismic survey in the Altiplano and Puna (18-24° S) indicates a dominantly felsic mineralogy in the upper crust and in large parts of the lower crust (e.g. Wigger et al 1994; Giese et al 1999; Schmitz et al 1999; Swenson et al 2000; Beck & Zandt 2002). There are local high velocity zones at a depth of >40 km south of 21° S (Wigger et al 1994; Schmitz et al 1999), but they are not seen in the north (Swenson

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a.

Okm Upper crust with abundant intrusions (c. 560 - 460 Ma) from crustal melts or hybrid arc magmas and late Proterozoic-Ordovician sediments

15

25

Partially molten zone: deepest exposed sections of the orogen showing uniform felsic migmatite (age of high-T metamorphism c. 530 - 420 Ma) Upper amphibolite to granulite fades mainly felsic metamorphic rocks rarely exposed. The composition of this crust is inferred from the composition of magmatic rocks and lower crustal xenoliths. Substantial depletion of Rb and U occurred during early Palaeozoic melting and metamorphism and arrested the U-Pb and Rb-Sr isotope systems. This crust comprises the bulk of the present crust of the plateau.

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Fig. 7. (a) The Palaeozoic Orogen: summary of key features related to depth. These petrological key features also apply to (b), the Cenozoic Orogen. (b) The interpreted geophysical image of the plateau region indicates an upper limit of partially molten crust at a depth of c. 20-25 km from the attenuation of seismic velocities (Altiplano low velocity zone). The zone of high electrical conductivity is also interpreted as caused by partial melting in the mid-crust. Ignimbrite chemistry and crustal contribution to the andesite indicate widely distributed generation of crustal melts in the Palaeozoic mid-crust of the plateau (Figs 4, 5, 6). Modified from Brasse et al. (2002, fig. 8).

et al. 2000) and the nature of these zones remains speculative (Schmitz et al. 1999). The interpretation of the gravity field also indicates predominantly felsic crust in the plateau region (Gotze & Kirchner 1997). The isotopic composition (Fig. 5b, 6b) of the Andean magmas and the Nd model ages (Fig. 4) of the Cenozoic large-scale ignimbrite are in perfect agreement with such a model of a felsic crust mainly made

up by migmatites, gneisses and granitoids of the early Palaeozoic Orogen. This is in contrast with the commonly assumed mafic composition of the lower crust, which might have been created by mafic underplating. Indeed, magmatic thickening by juvenile mafic material from the arc has been invoked for the Altiplano (Lamb & Hoke 1997), but voluminous mafic rocks in the lower crust can neither be proven by the

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volume of surface magmatism (Francis & Hawkesworth 1994) nor by the composition of the Cenozoic magmas. Most magmatic rock compositions do not indicate the involvement of mafic rocks in the Cenozoic and early Palaeozoic. Post-Palaeozoic formation of new crust by substantial addition of mafic to intermediate mantle-derived igneous rocks was restricted to the Jurassic-lower Cretaceous magmatic arcs of the Coastal Cordillera (Figs 1, 5b; e.g. Rogers & Hawkesworth 1989; Lucassen etal 1996). These rocks now form part of the Cenozoic forearc and are not involved in the plateau formation.

What makes both orogens special? Characteristics of the early Palaeozoic Orogen and the Cenozoic plateau and their current interpretations indicate striking similarities for both orogens. 1. High temperature (T >650 °C) and abundant partial melting occurs at a depth of c. 15-25 km (0.4-0.7 GPa) according to the P-T record and abundant crustal melts of the Palaeozoic Orogen. Similar metamorphic conditions are indicated in the crust of the Cenozoic plateau according to geophysical properties and extended voluminous ignimbrite volcanism. The uniform metamorphic conditions of exhumed early Palaeozoic rocks indicate an original crustal thickness of >50 km over a large area. Such a body of prevailing felsic crust in nearisostatic equilibrium would form a high plateau, similar to the Recent Andean plateau. The Palaeozoic Orogen is therefore viewed as an exhumed and eroded analogue of the Andean Orogen. Provided no continent-continent collision occurs, erosion of the presently thickened crust of c. 50-70 km of the Cenozoic plateau to a stable surface elevation near sealevel and a crustal thickness of c. 35-40 km would exhume felsic migmatite from the zone of partial melting in the deepest exposures. Magma chambers of ignimbrite, andesite and diapir intrusions between the deepest future exposures and the present-day surface of the high plateau would form extended areas of intermediate to granitic plutonic rocks. 2. Arc-related magmatism occurs throughout the evolution of the early Palaeozoic and Cenozoic orogens. Crustal growth and thickening by juvenile additions from depleted sub-arc mantle seems to be small in both orogens. Recycling of Proterozoic crust dominates the magmatic processes in the generally hot crust. The expected basal accretion or intrusion into the lower crust of juvenile basaltic material during the long activity of magmatic arcs did

either not result in a continuous thick mafic lower crust or the mafic crust and underlying mantle lithosphere were recycled into the mantle by density-driven delamination (e.g. Meissner & Mooney 1998). 3. A longstanding heat source is necessary to maintain high temperatures in the crust. The Cenozoic deformation leading to plateau formation started in the early Tertiary at c. 40-50 Ma (Scheuber et al 1994; Kennan et al 1995) and it was preceded by c. 200 Ma of subduction. The plateau (<25 Ma) hosts the extensive ignimbrite volcanism (peak <10 Ma, e.g. de Silva 1989). Major sources of heat in the Cenozoic Orogen are the shallow position of the thermal asthenosphere-lithosphere boundary in a mantle wedge of the arc after thermal erosion or delamination of the mantle lithosphere, and accretion of magmas at the base of the overthickened crust (e.g. Meissner & Mooney 1998; Babeyko et al 2002). For the Palaeozoic Orogen the long period of high-temperature metamorphism and/or mainly crustal-derived magmatism from Cambrian (530 Ma) to the Ordovician (460 Ma) or even the Silurian (420 Ma) is best explained by the long-term stability of a similar arc setting. 4. The mainly felsic crust has undergone several metamorphic-magmatic-sedimentary cycles in the framework of the western South American craton. Quartz-feldspar rheology is dominant in this crust, i.e. ductile deformation is the response of this crust already at moderate temperatures (>400-500 °C depending on the mineral modes; e.g. Passchier & Trouw 1996). There is no evidence for large-scale complex stacking of different crustal units in the Palaeozoic or later, which is important in the formation of thick crust in collisional orogens such as the Himalayas-Tibet Orogen (e.g. Yin & Harrison 2000). The Cenozoic underthrusting of the Brazilian Shield inferred from geophysical data does not yet reach the plateau area (e.g. Beck & Zandt 2002) and does not contribute to the thick crust of the plateau. The dominant process of crustal thickening in the plateau is ductile (possibly homogeneous) deformation of weak quartz-feldspar-dominated crust (e.g. Chmielowski et al 1999, Beck & Zandt 2002) and not stacking of crustal units. 5. High pressure-low temperature metamorphic rocks (eclogite, blueschist) indicating substantial tectonic exhumation during compression are absent, e.g. obducted units from the ocean floor (ophiolites), deeply subducted rocks or accretionary wedges, which are typical of 'cold' collisional orogens. Exhumation of high-pressure, mainly blueschist

EARLY PALAEOZOIC ANDEAN OROGENESIS

facies, rocks in non-collisional settings seems to be restricted to specific conditions in the forearc favouring the formation of an accretionary wedge. The preservation of possible accretionary wedges in the Central Andes (e.g. a late Palaeozoic accretionary wedge is partly preserved south of c. 32° S; e.g. Herve 1988) or destruction was subject to differences in the forearc dynamics, e.g. subduction erosion of the continental plate (e.g. Clift & Vannuchi 2004). 6. The initiation of subduction orogeny is still speculative but both orogens initiated in a global tectonic situation of continental breakup and crustal thickening was preceded by continental extension in an area, which later formed part of the arc and back-arc. This extension did not result in the formation of ocean floor but could have caused a thermally weakened lithosphere. Thermal weakening of felsic (quartz-feldspar-dominated) crust and mantle lithosphere is considered a prerequisite of homogeneous deformation and formation of thickened continental crust (Thompson et al 2001). Thermally softened continental extensional zones (rifts and arcs) as precursors to thickened orogenic belts might be the prerequisite for arc-related plateau formation. Xenoliths studies from the late Cretaceous rift indicate high temperature in the lower crust (c. 900 °C) of a crust of normal thickness (c. 35 km) and in the upper mantle (c. 1000-1100 °C) at 26° S (Lucassen et al 19996; 2005). The Palaeozoic and Cenozoic orogens both formed in a non-collisional setting at a continental margin, which was continuously or almost continuously active during compression and crustal thickening. Single characteristics of the orogens summarized are also found in collisional orogens, e.g. crustal melting in overthickened felsic crust is also common in parts of the Himalayas-Tibet Orogen where heating was attributed to various processes including delamination of mantle lithosphere (see e.g. review of the Himalayas-Tibet Orogen by Yin & Harrison 2000). In parts of the Tibetan plateau a similar geophysical image of the crust (Alsdorf et al 1998, Li et al 2003) as seen for the Central Andes is observed. Important other features of collisional orogens (Alps, Himalayas-Tibet; e.g. O'Brien 2000, Liou et al 2004) or collision of continental fragments are, however, missing in the Central Andes as outlined above. Furthermore, the contemporaneous magmatism in the collision regime is different in composition and extent compared with the Cenozoic and early Palaeozoic arc magmatism in the Central Andes. The early Palaeozoic Orogen comprises magmatic rocks

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from an arc setting, which formed contemporaneously with the thick, hot crust. This excludes at least the collision of a continent or large continental fragment, which should result in a ceasing of arc activity and - in the case of renewed onset of subduction - to shift of the location of the arc. This is not observed in the evolution of the early Palaeozoic Orogen. This contribution is a result of the project SFB 267 'Deformation Processes in the Andes - Gl Thermal and Compositional Structure of the Lithosphere' and the authors thank all their colleagues for stimulating discussions, as well as the Deutsche Forschungsgemeinschaft for financial support. U. Riller is thanked for his comments on an earlier version of the manuscript, and U. Cordani, P. Leat and an anonymous reviewer for their critical and helpful reviews.

References ACENOLAZA, FG & MILLER, H. 1982. Early Paleozoic orogeny in southern South America. Precambrian Research, 17, 133-146. ACENOLAZA, FG, MILLER, H. & TOSELLI, A.J. 2002. Proterozoic-Early Paleozoic evolution in western South America - a discussion. Tectonophysics, 352,121-137. ALLMENDINGER, R., JORDAN, T.E., KAY, S.M. & ISACKS, B.L. 1997. The evolution of the Altiplano-Puna plateau of the Central Andes. Annual Review Earth Planetary Sciences, 25, 139-174. ALSDORF, D., BROWN, L., NELSON, K.D., MAKOVSKY, Y., KLEMPERER, S. & ZHAO, W. 1998. Crustal deformation of the Lhasa terrane, Tibet plateau from project INDEPTH deep seismic reflection profile. Tectonics, 17, 501-519. ASTINI, R.A. & DAVILA, FM. 2004. Ordovician back arc foreland and Ocloyic thrust belt development on the western Gondwana margin as a response to Precordillera terrane accretion. Tectonics, 23. Doi: 10.1029/2003TC001620. ASTINI, R.A., BENEDETTO, J.L. & VACCARI, N.-E. 1995. The Early Paleozoic evolution of the Argentine Precordillera as a Laurentian rifted, drifted, and collided terrane: A geodynamic model. Geological Society of America Bulletin, 107, 253-273. AUGUSTSSON, C. & BAHLBURG, H. 2003. Cathodoluminesence spectra of detrital quartz as provenance indicators of Paleozoic metasediments in Southern Andean Patagonia. Journal of South American Earth Sciences, 16, 15-26. BABEYKO, A.Y., SOBOLEV, S.V., TRUMBULL, R.B., ONCKEN, O. & LAVIER, L.L. 2002. Numerical model of crustal scale convection and partial melting beneath the Altiplano-Puna plateau. Earth and Planetary Science Letters, 199,373-388. BAHLBURG, H. & BREITKREUZ, C. 1991. Paleozoic evolution of active margin basins in the southern Central Andes (northwestern Argentina and northern Chile). Journal of South American Earth Sciences, 4, 171-188.

270

F. LUCASSEN & G. FRANZ

BAHLBURG, H. & HERVE, F. 1997. Geodynamic evolution and tectonostratigraphic terranes of northwestern Argentina and northern Chile. Geological Society of America Bulletin, 109, 869-884. BALDO, E.G., DEMANGE, M. & MARTINO, R.D. 1996. Evolution of the Sierras de Cordoba, Argentina. Tectonophysics, 267,121-146. BECCHIO, R., LUCASSEN, F, FRANZ, G, VIRAMONTE, J. & WEMMER, K. 1999. El basamento paleozoico inferior del noroeste de Argentina (23°-27°S) metamorfismo y geocronologia. In: BONORINO, G.G., OMARINI, R. & VIRAMONTE, J. (eds) Geologia del noroeste Argentino. XIV Congreso Geologico Argentino, Salta, Argentina, Relatorio, Artes Graficas S.A., Salta, Argentina, 58-72. BECK, S.L. & ZANDT, G. 2002. The nature of erogenic crust in the central Andes. Journal of Geophysical Research, 107. Doi: 10.1029/2001JB000124. BERG, K. & BAUMANN, A. 1985. Plutonic and metasedimentary rocks from the Coastal range of northern Chile: Rb-Sr and U-Pb isotopic systematics. Earth and Planetary Science Letters, 75, 101-115. BOCK, B., BAHLBURG, H., WORNER, G. & ZIMMERMANN, U. 2000. Tracing crustal evolution in the southern Central Andes from Late Precambrian to Permian using Nd and Pb isotopes. Journal of Geology, 108, 515-535. BRASSE, H., LAZAETA, P., RATH, V. SCHWALENBERG, K., SOYER & W. HAAK, V. 2002. The Bolivian Altiplano conductivity anomaly. Journal of Geophysical Research, 107. Doi: 10.1029/2001JB000391. BREITKREUZ, C. & ZEIL, W. 1994. The Late Carboniferous to Triassic volcanic belt in Northern Chile. In: REUTTER, K.J., SCHEUBER, E. & WIGGER, P.J. (eds) Tectonics of the Southern Central Andes. Springer Verlag, Heidelberg, 277-292. BROWN, M. 1991. Comparative geochemical interpretation of Permian-Triassic plutonic complexes of the Coastal Range and Altiplano (23°30' to 26°30' S), northern Chile. In: HARMON, R.S. & RAPELA, C.W. (eds) Andean magmatism and its tectonic setting. Geological Society of America, Special Paper, 265,157-171. BUTTNER, S.H., GLODNY, J., LUCASSEN, F., WEMMER, K., ERDMANN, S., HANDLER, R. & FRANZ, G. 2005. Ordovician metamorphism and plutonism in the Sierra de Quilmes metamorphic complex: Implications for the tectonic setting of the northern Sierras Pampeanas (NW Argentina). Lithos, in press. CASQUET, C., BALDO, E., PANKHURST, R.J., RAPELA, C.W., GALINDO, C., FANNING, C.M. & SAAVREDA, J. 2001. Involvement of the Argentine Precordillera terrane in the Famatinian mobile belt: U-Pb SHRIMP and metamorphic evidence from the Sierra de Pie de Palo. Geology, 29, 703-706. CHARRIER, R. & REUTTER, K.J. 1994. The Purilactis group of northern Chile: Boundary between arc and back-arc from late Cretaceous to Eocene. In: REUTTER, K.J., SCHEUBER, E. & WIGGER, P.J. (eds) Tectonics of the Southern Central Andes. Springer Verlag, Heidelberg, 189-202.

CHMIELOWSKI, J., ZANDT, G & HABERLAND, C. 1999. The Central Andean Altiplano-Puna magma body. Geophysical Research Letter, 26, 783-786. CLIFT, P. & VANNUCCHI, P. 2004. Controls on tectonic accretion versus erosion in subduction zones: implications for the origin and recycling of continental crust. Review of Geophysics, 42. Doi: 10.1029/2003RG000127. COIRA, B.L., MAHLBURG-KAY, S., PEREZ, B., WOLL, B., HANNING, M. & FLORES, P. 1999. Magmatic sources and tectonic setting of Gondwana margin Ordovician magmas, northern Puna of Argentina and Chile. In: RAMOS, V.A. & KEPPIE, ID. (eds) Laurentia-Gondwana Connection before Pangea. Geological Society of America, Special Paper, 336, 145-170. CORDANI, U.G., SATA, K., TEIXEIRA, W, TASSARINI, C.C.G. & BASEI, M.A.S. 2000. Crustal evolution of the South American platform. In: CORDANI, U.G., MILANI, E.I, THOMAZ FILHO, A. & CAMPOS, D. A. (eds) Tectonic evolution of South America. 31st International Geological Congress, Rio de Janeiro, Brazil, 19-40. CORDANI, U.G., D'AGRELLA FILHO, M.S., BRITONEVES, B.B. & TRINDADE, R.I.F. 2003. Tearing up Rodinia: the Neoproterozoic palaeogeography of South American cratonic fragments. Terra Nova, 15, 350-359. DAMM, K.W., PICHOWIAK, S., HARMON, R.S., TODT, W, KELLEY, S., OMARINI, R. & NIEMEYER, H. 1990. Pre-Mesozoic evolution of the central Andes; The basement revisited. In: KAY, S.M. & RAPELA, C.W. (eds) Plutonism from Antarctica to Alaska. Geological Society of America, Special Paper, 241, 101-126. DAMM, K.W, HARMON, R.S. & KELLEY, S. 1994. Some isotope and geochemical constraints on the origin and evolution of the Central Andean basement (19°-24°S). In: REUTTER, K.I, SCHEUBER, E. & WIGGER, P.I (eds) Tectonics of the Southern Central Andes. Springer Verlag, Heidelberg, 263-275. DE SILVA, S.L. 1989. Altiplano-Puna volcanic complex of the Central Andes. Geology, 17,1102-1106. DURAND, F.R. & ACENOLAZA, F.G. 1990. Caracteres biofaunisticos, paleoecologicos y paleogeograficos de la Formacion Puncoviscana (Precambrico superior-Cambrico inferior) del noroeste Argentino. In: ACENOLAZA, F.G, MILLER, H. & TOSELLI A.I (eds) El Ciclo Pampeano en el Noroeste Argentino. Universidad Nacional de Tucuman, Argentina, Serie Correlation Geologica, N4, 71-112. EGENHOFF, S.O. & LUCASSEN, F. 2003. Chemical and Isotopic Composition of Lower to Upper Ordovician Sedimentary Rocks (Central Andes/South Bolivia): Implications for their Source. Journal of Geology, 111, 487-497. FRANCIS, PW. & HAWKESWORTH, C.I 1994. Late Cenozoic rates of magmatic activity in the Central Andes and their relationship to continental crust formation and thickening. Journal of the Geological Society, London, 151, 845-854. FRANCIS, PW, SPARKS, R.S.I, HAWKESWORTH, C.J.

EARLY PALAEOZOIC ANDEAN OROGENESIS ET AL. 1989. Petrology and geochemistry of volcanic rocks of the Cerro Galan caldera, northwest Argentina. Geological Magazine, 126, 515-547. GALLISKI, M.A. & VIRAMONTE, J.G. 1988. The Cretaceous paleorift in northwestern Argentina; a petrological approach. Journal of South American Earth Sciences, 1, 329-343. GlESE, P., SCHEUBER, E., SCHILLING, E, SCHMITZ, M. &

WIGGER, P. 1999. Crustal thickening processes in the Central Andes and the different nature of the Moho-discontinuity. Journal of South American Earth Sciences, 12, 201-220. GOLDSTEIN, S.L., O'NiONS, R.K. & HAMILTON, P.J. 1984. A Sm-Nd study of atmospheric dust and particulates from major river systems. Earth and Planetary Science Letters, 70, 221-236. GOLDSTEIN, S.L., ARNDT, N.T. & STALLARD, R.F. 1997. The history of continent from U-Pb ages of zircons from Orinoco river sand and Sm-Nd isotopes in Orinoco basin river sediments. Chemical Geology, 139, 271-286. GOTZE, H.-J. & KIRCHNER, A. 1997. Interpretation of gravitiy and geoid in the Central Andes between 20° and 29°S. Journal of South American Earth Sciences, 10, 179-188. HARMON, R.S., BARREIRO, B.A., MOORBATH, S. ET AL. 1984. Regional O-, Sr-, and Pb-isotope relationships in late Cenozoic calc-alkaline lavas of the Andean Cordillera. Journal of the Geological Society, London, 141, 803-822. HERVE, E 1988. Late Paleozoic subduction and accretion in southern Chile. Episodes, 11, 183-188. HILDRETH, W. & MOORBATH, S. 1988. Crustal contributions to arc magmatism in the Andes of central Chile. Contributions to Mineralogy and Petrology, 98, 455-489. HOECKENREINER, M., SOLLNER, F. & MlLLER, H. 2003.

Dating the TIPA shear zone: an early Devonian terrane boundary between the Famatinian and Pampean systems (NW Argentina). Journal of South American Earth Sciences, 16, 45-66. JEZEK, P., WILLNER, A.P., ACENOLAZA, EG. & MILLER, H. 1985. The Puncoviscana trough - a large basin of Late Precambrian to Early Cambrian age on the Pacific edge of the Brazilian shield. Geologische Rundschau, 74, 573-584. KAY, S.M., ORRELL, S. & ABBRUZZI, J.M. 1996. Zircon and whole rock Nd-Pb evidence for a Grenville age and a Laurentian origin for the basement of the Precordillera in Argentina. Journal of Geology, 104, 637-648. KENNAN, L., LAMB, S. & RUNDLE, C. 1995. K-Ar date from the Altiplano and Cordillera Oriental of Bolivia: Implications for Cenozoic stratigraphy and tectonics. Journal of South American Earth Sciences, 8, 163-186. KLEINE,TH., MEZGER, K., ZIMMERMANN, U., MUNKER, C. & BAHLBURG, H. 2004. Crustal evolution of the early Ordovician Proto-Andean margin of Gondwana: Trace element and isotope evidence from Complejo Igneo Pocitos (Northwest Argentina). Journal of Geology, 112, 503-520. LAMB, S. & HOKE, L. 1997. Origin of the high plateau

271

in the Central Andes, Bolivia, South America. Tectonics, 16, 623-649. Li, S., UNSWORTH, M.J., BOOKER, J.R., WEI, W.,TAN, H. & JONES, A.G. 2003. Partial melt or aqueous fluid in the mid-crust of Southern Tibet? Constraints from INDEPTH magnetotelluric data. Geophysical Journal International, 153, 289-304. LINDSAY, J.M., SCHMITT, A.K., TRUMBULL, R.B., DE SILVA, S.L., SIEBEL, W. & EMMERMANN, R. 2001. Magmatic evolution of the La Pacana Caldera system, central Andes, Chile: compositional variation of two cogenetic, large-volume felsic ignimbrites and implications for contrasting eruption mechanisms. Journal of Petrology, 42, 459^86. Liou, J.G, TSUJIMORI, T, ZHANG, R.Y., KATAYAMA, I. & MAYURAMA, S. 2004. Global UHP metamorphism and continental subduction/collision: The Himalayan model. International Geology Review, 46, 1-27. LOEWY, S.L., CONNELLY, J.N. & DALZIEL, I.W.D. 2004. An orphaned basement block: The ArequipaAntofalla basement of the central Andean margin of South America. Geological Society of America Bulletin, 116, 171-187. LUCASSEN, F. & BECCHIO, R. 2003. Timing of highgrade metamorphism: Early Paleozoic U-Pb formation ages of titanite indicate long-standing high-T conditions at the western margin of Gondwana (Argentina, 26-29°S). Journal of Metamorphic Geology, 21, 649-662. LUCASSEN, E, FOWLER, C.M.R. & FRANZ, G. 1996. Extension and magmatic accretion during the Early- and Mid-Mesozoic in the North Chilean Coast Range: A geological and thermal model. Tectonophysics, 262, 263-279. LUCASSEN, E, FRANZ, G. & LABER, A. 19990. Permian high pressure rocks - the basement of Sierra de Limon Verde in N-Chile. South American Journal of Earth Sciences, 12, 183-199. LUCASSEN, F, LEWERENZ, S., FRANZ, G, VIRAMONTE, J. & MEZGER, K. 19995. Metamorphism, isotopic ages and composition of lower crustal granulite xenoliths from the Cretaceous Salta Rift, Argentina. Contributions to Mineralogy and Petrology, 134, 325-341. LUCASSEN, E, BECCHIO, R., WILKE, H.G., THIRLWALL, M.F., VIRAMONTE, I, FRANZ, G. & WEMMER, K. 2000. Proterozoic-Paleozoic development of the basement of the Central Andes (18°-26°) - a mobile belt of the South American craton. Journal of South American Earth Sciences, 13, 697-715. LUCASSEN, F, BECCHIO, R., HARMON, R. ET AL. 2001. Composition and density model of the continental crust at an active continental margin - the Central Andes between 21° and 27°S. Tectonophysics, 341,195-223. LUCASSEN, F, HARMON, R., FRANZ, G, ROMER, R.L., BECCHIO, R. & SIEBEL, W. 20020. Lead evolution of the Pre-mesozoic crust in the Central Andes (18°-27°): Progressive homogenisation of Pb. Chemical Geology, 186, 183-197. LUCASSEN, E, ESCAYOLA, M., FRANZ, G, ROMER, R.L.

272

F. LUCASSEN & G. FRANZ

& KOCH, K. 2002ft. Isotopic composition of Late Mesozoic basic and ultrabasic rocks from the Andes (23-32°S) - implications for the Andean mantle. Contributions to Mineralogy and Petrology, 143, 336-349. LUCASSEN, F., FRANZ, G., VIRAMONTE, J., ROMER, R.L., DULSKI, P. & LANG, A. 2005. The late Cretaceous lithospheric mantle beneath the Central Andes: evidence from phase equilibrium and composition of mantle xenoliths. Lithos, in press. MEISSNER, R. & MOONEY, W. 1998. Weakness of the lower continental crust: a condition for delamination, uplift and escape. Tectonophysics, 296, 47-60. MOYA, M.C. 1988. Lower Ordovician in the southern part of the Argentine Eastern Cordillera. In: BAHLBURG, H., BREITKREUZ, CH. & GIESE, P. (eds) The southern Central Andes. Lecture Notes in Earth Sciences, 17. Springer Verlag, Heidelberg, 55-69. O'BRIEN, PJ. 2000. Subduction followed by collision: Alpine and Himalayan examples. Physics of the Earth and Planetary Interiors, 127, 277-297. OMARINI, R.H., SUREDA, R.J., GOTZE, H.J., SEILACHER, A. & PFLUGER,F. 1999. Puncoviscana folded belt in northwestern Argentina: testimony of Late Proterozoic Rodinia fragmentation and pre-Gondwana collisional episodes. International Journal of Earth Sciences (Geologische Rundschau), 88, 76-97. ORT, M.H., COIRA, B.L. & MAZONI, M.M. 1996. Generation of a crust-mantle magma mixture: magma sources and contamination at Cerro Panizoz, central Andes. Contributions to Mineralogy and Petrology, 123, 308-322. PANKHURST, RJ. & RAPELA, CW. (eds) 1998. The Proto-Andean margin of Gondwana. Geological Society, London, Special Publications, 142. PANKHURST, R.J., RAPELA, C.W., SAAVEDRA, I, BALDO, E., DAHLQUIST, J., PASCUA, I. & FANNING, C.M. 1998. The Famatinian magmatic arc in the central Sierras Pampeans: an Early-to-Middle Ordovician continental arc on the Gondwana margin. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean margin of Gondwana. The Geological Society, London, Special Publications, 142, 343-367. PANKHURST, R.J., RAPELA, C.W. & FANNING, C.M. 2000. Age and origin of coeval TTG, I- and Stype granites in the Famatinian belt of NW Argentina. Transactions of the Royal Society of Edinburgh Earth Sciences, 91, 151-168. PASSCHIER, CW. & TROUW, R.AJ. 1996. Microtectonics. Springer Verlag, Berlin. PIPER, J.D.A. 2000. The Neoproterozoic Supercontinent: Rodinia or Palaeopangaea? Earth and Planetary Science Letters, 176,131-146. PRINZ, P., WILKE, H.G. & HILLEBRANDT, A.v. 1994. Sediment accumulation and subsidence history in the Mesozoic marginal basin of northern Chile. In: REUTTER, K.J., SCHEUBER, E. & WIGGER, PJ. (eds) Tectonics of the Southern Central Andes. Springer Verlag, Heidelberg, 219-232.

RAMOS, V.A. 1995. Sudamerica: un mosaico de continentes y oceanos. Ciencia hoy, 6, 24-29. RAMOS, V.A. & KEPPIE, J.D. (eds) 1999. LaurentiaGondwana Connection before Pangea. Geological Society of America, Special Paper, 336. RAMOS, V.A., JORDAN, T.E., ALLMENDINGER, R.W., MPODOZIS, C, KAY, S.M., CORTES, J.M. & PALMA, M.A. 1986. Palaeozoic terranes of the central Argentine Chilean Andes. Tectonics, 5, 855-880. RAPELA, C.W, PANKHURST, R.J., CASQUET, C., BALDO, E., SAAVEDRA, J., GALINDO, C. & FANNING, C.M. 1998. The Pampean Orogeny of the southern proto-Andes: evidence for Cambrian continental collision in the Sierras de Cordoba. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean margin of Gondwana. Geological Society, London, Special Publications, 142, 181-217. RING, S., KOMIY, T, WINDLEY, B.F., KATAYAMA, I., MOTOKI, A. & HIRATA, T. 2004. Major episodic increases of continental crustal growth determined from zircon age of river sands; implications for mantle overturns in the Early Precambrian. Physics of the Earth and Planetary Interiors, 146, 369-394. ROGERS, G. & HAWKESWORTH, CJ. 1989. A geochemical traverse across the North Chilean Andes: evidence for crust generation from the mantle wedge. Earth and Planetary Science Letters, 91, 271-285. SATO, K. & SIGA JR. O. 2002. Rapid growth of continental crust between 2.2 and 1.1 Ga in the South American Platform: Integrated Australian, European, North American and SW USA crustal evolution study. Gondwana Research, 5,165-173. SCHEUBER, E., BOGDANIC, T, JENSEN, A. & REUTTER, KJ. 1994. Tectonic development of the north Chilean Andes in Relation to Plate Convergence and Magmatism since the Jurassic. In: REUTTER, K J., SCHEUBER, E. & WIGGER, P. J. (eds) Tectonics of the Southern Central Andes Springer Verlag, Heidelberg, 7-22. SCHMITZ, M., LESSEL, K., GIESE, P. ET AL. 1999. The crustal structure beneath the Central Andean forearc and magmatic arc as derived from seismic studies - the PISCO 94 experiment in northern Chile (21°-23°S). Journal of South American Earth Sciences, 12, 237-260. SCHWARTZ, J. J. & GROMET, L.P. 2004. Provenance of a late Proterozoic-early Cambrian basin, Sierras de Cordoba, Argentina. Precambrian Research, 129,1-21. SEGEMAR (SERVICIO GEOLOGICO Y MINERO ARGENTINO) 1997. Map a geologico de la Republica Argentina, E: 2.500.00. Secretaria de Industria Comercio y Mineria de Argentina, Buenos Aires. SEMPERE,T, CARLIER, G, SOLER, P. ET AL. 2002. Late Permian-Middle Jurassic lithospheric thinning in Peru and Bolivia, and its bearing on Andean-age tectonics. Tectonophysics, 345,153-181. SlEBEL, W, SCHNURR, W.B.W, HAHNE, K., KRAEMER,

B., TRUMBULL, R.B., VAN DEN BOGAARD, P. &

EARLY PALAEOZOIC ANDEAN OROGENESIS EMMERMANN, R. 2001. Geochemistry and isotope systematics of small- to medium-volume Neogene-Quarternary ignimbrites in the southern central Andes: evidence for derivation from andesitic magma sources. Chemical Geology, 171, 213-237. SPRINGER, M. 1999. Interpretation of heat-flow density in the Central Andes. Tectonophysics, 306, 377-395. SUAREZ, M. & BELL, C.M. 1992. Triassic rift-related sedimentary basins in northern Chile (24°-29°S). Journal of South American Earth Sciences, 6, 109-121. STUART-SMITH, P.G., MIRO, R., SIMS, J.P. ET AL. 1999. Uranium-lead dating of felsic magmatic cycles in the southern Sierras Pampeanas, Argentina: Implications for the tectonic development of the proto-Andean Gondwana margin. In: RAMOS, V. & KEPPIE, D. (eds) Laurentia-Gondwana Connection before Pangea. Geological Society of America, Special Paper, 336, 87-11. SWENSON, J.L., BECK, S.L. & ZANDT, G. 2000. Crustal structure of the Altiplano from broadband regional waveform modeling: Implications for the composition of thick continental crust. Journal of Geophysical Research, 105, 607-621. THOMPSON, A.B., SCHULMANN, K., JEZEK, J. & TOLAR, V. 2001. Thermally softened continental extensional zones (arcs and rifts) as precursors to thickened erogenic belts. Tectonophysics, 332, 115-141. TRUMBULL, R.B., WITTENBRINK, R., HAHNE, K., EMMERMANN, R., BUSCH, W, GERSTENBERGER, H. & SIEBEL, W. 1999. Evidence for Late Miocene to Recent contamination of arc andesites by crustal melts in the Chilean Andes (25°-26°S) and its geodynamic implications. Journal of South American Earth Sciences, 12, 135-155. VALERIANO, C.M., MACHADO, N., SIMONETTI, A., VALLADARES, C.S., SEER, H.S. & SIMOES, L.S.A. 2004. U-Pb geochronology of the southern Brasilia belt (SE-Brazil): sedimentary provenance, Neoproterozoic orogeny and assembly of West Gondwana. Precambrian Research, 130, 27-55. VARELA, R., SATO, A.M., BASEI, M.A.S. & SIGA JR, O. 2003. Proterozoico medio y Paleozoico inferior de la sierra de Umango, antepais andino (29°S). Argentina: edades U-Pb y caracterizaciones isotopicas. Revista geologica de Chile, 30,265-284. VIRAMONTE, J.G., KAY, S.M., BECCHIO, R., ESCAYOLA,

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M. & NOVITSKI, I. 1999. Cretaceous rift related magmatism in central-western South America. Journal of South American Earth Sciences, 12, 109-121. VUJOVICH, G.I. & KAY, S.M.A. 1998. A Laurentian? Grenville.age oceanic arc/back-arc terrane in the Sierra Pie de Palo, western Sierras Pampeanas, Argentina. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean margin of Gondwana. Geological Society, London, Special Publications, 142,159-180. WlGGER, P.I, SCHMITZ, M., ARANEDA, M. ET AL. 1994.

Variation in the crustal structure of the Southern Central Andes deduced from seismic refraction investigations. In: REUTTER, K.J., SCHEUBER, E. & WIGGER, P.J. (eds) Tectonics of the Southern Central Andes Springer Verlag, Heidelberg, 23-48. WILLNER, A.P., MILLER, H. & JEZEK, P. 1985. Geochemical features of an Upper PrecambrianLower Cambrian greywacke/pelite sequence (Puncoviscana trough) from the basement of the NW-Argentine Andes. Neues Jahrbuch fur Geologie und Palaontologie, Monatshefte, 8, 498-512. WORNER, G, MOORBATH, S., HORN, S., ENTENMANN, J.,

HARMON, R.S., DAVIDSON, J.P. & LOPEZ-ESCOBAR, L. 1994. Large- and fine-scale variations along the Andean arc of Northern Chile (17.5°-22°S). In: REUTTER, K.J., SCHEUBER, E. & WIGGER, P.J. (eds) Tectonics of the Southern Central Andes. Springer Verlag, Heidelberg, 77-92. WORNER, G, LEZAUN, I, BECK, A. ET AL. 2000. Geochronology, metamorphic petrology and geochemistry of basement rocks from Belen (N. Chile) and C. Uyarani (W. Bolivian Altiplano): Implications for the evolution of Andean basement. Journal of South American Earth Sciences, 13, 717-737. YIN, A. & HARRISON, T.M. 2000. Geologic evolution of the Himalayan-Tibetan orogen. Annual Review Earth Planetary Sciences, 28, 211-280. ZANDT, G, LEIDIG, M., CHMIELOWSKI, J., BAUMONT, D. &YUAN, X. 2003. Seismic detection and characterization of the Altiplano-Puna magma body. Pure and Applied Geophysics, 160, 789-807. ZIMMERMANN, U. & BAHLBURG, H. 2003. Provenance analysis and tectocnic setting of the Ordovician clastic deposits in the southern Puna Basin; NW Argentina. Sedimentology, 50, 1079-1104.

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Northern Victoria Land terranes, Antarctica: far-travelled or local products? FRANZ TESSENSOHN & FRIEDHELM HENJES-KUNST Bundesanstalt fur Geowissenschaften und Rohstoffe, Stilleweg2,30655 Hannover, Germany (e-mail: franz. tessensohn@tiscali. de) Abstract: The early Palaeozoic Ross Orogen in northern Victoria Land, Antarctica, consists of three major fault-bounded tectonostratigraphic terranes. Their true nature, fartravelled or local part of the accretionary collage, is under discussion. The inboard Wilson terrane (WT) consists mainly of high-grade metamorphic rocks intruded by the calcalkaline magmatic arc of the Ross Orogen. The terrane nature of the WT is doubtful, as it appears more like the leading edge of the East Antarctic craton. The Bowers terrane (BT) comprises a mixed sedimentary-volcanic succession, beginning with volcanic rocks of island-arc character, followed by turbidites, mudstones, conglomerates and fossiliferous Middle Cambrian shallow-water sediments. The whole sequence is capped by a fluvial to deltaic quartzitic series several kilometres thick, with strong continental affinity. The combination of primitive forearc to back-arc volcanics at the bottom and mature continental sediments at the top poses a problem. The outer Robertson Bay terrane (RBT) is made up of a thick turbidite succession which, in one area, contains allochthonous blocks of fossiliferous Tremadocian limestones. All terrane boundaries appear to be distinct fault zones. The WT/BT boundary forms a deep-reaching continent-ocean suture associated with strongly sheared rock units, ultramafic lenses and high-pressure rocks. Coesite in eclogites of the Lanterman Range indicates a depth of burial of around 90 km. A greenschist-facies schist belt marks the BT/RBT boundary. The terranes contain evidence for subduction at an active margin setting as well as for accretion processes along major faults. The present changes of the Cambrian time-scale, such as younging of the base of the Upper Cambrian by about 30 Ma since the 1980s, allow separation of arc formation and later terrane accretion events.

Terranes are usually features of an active convergent margin. They are recognized by their tectonostratigraphic difference to the neighbouring rock units and are, therefore, in principle easy to define. A key problem, however, is usually the distinction between fartravelled (exotic) and locally derived units, particularly if no reliable palaeomagnetic data are available. At a convergent margin, the contrasting continental rocks of the overriding plate and the oceanic units of the subducting plate, including accretionary wedge and forearc basins, usually fulfil the criteria for the terrane definition, even though they may not be allochthonous. The early Palaeozoic Ross orogenic terranes of northern Victoria Land, Antarctica, provide a good example for the controversial interpretation of terrane provenance, local or exotic. Combined international research efforts have provided a wealth of information and data, with often contradictory implications. Apart from the cited single publications, relevant summaries are contained in the following

articles and volumes: GANOVEX Team (1987), Ricci (1991; 1997). The early Palaeozoic Ross Orogen is a welldefined, at least 3000 km long belt, which today traverses most of the Antarctic continent (Fig. 1) (Stump 1995), but which, at the time of formation, formed part of the palaeo-Pacific active margin of the Gondwana supercontinent (Stump 1987; 1995; Kleinschmidt & Tessensohn 1987; Borg & De Paolo 1991; Palmeri et al 1994; Ricci et al. 1997; Talarico et al. 1998; Goodge 2002). It had its equivalents in the Delamerian Orogen of Australia and the Saldanian belt of South Africa. In the inboard part, the Ross Orogen is characterized by a diffuse low-P, lowto high-T regional metamorphism spanning the whole range from low-grade to granulite facies and migmatization (Grew et al 1984; Schubert et al 1984; GANOVEX Team 1987; Talarico & Castelli 1995; Talarico et al 1995; Schussler et al 1999). The rocks show a pronounced metamorphic peak deformation with axes generally parallel to the strike of the orogen from NW to

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 275-291. 0305-8719/$15.00 © The Geological Society of London 2005.

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Fig. 1. The magmatic arc of the early Palaeozoic Ross Orogen in Antarctica. The inset shows the tectonostratigraphic terranes of northern Victoria Land at the northern end of the Transantarctic Mountains.

SE. The metamorphic rocks are intruded mainly by calc-alkaline igneous rocks. Granites, granodiorites and tonalites predominate. Orogenic plutonism started at about 550 Ma in the central Transantarctic Mountains, while maximum igneous activity probably occurred close to 500 Ma (summarized in Goodge 2002). The

Palaeozoic Ross Orogen shows typical features of subduction accretion. In this respect, it is similar to the later Andes. The latter have a generally shallower level of erosion, with volcanic instead of plutonic rocks exposed. The fact that the Ross Orogen is an accretionary (rather than a collisional) orogen is important, because this is

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one of the reasons why the original active margin setting is so well preserved. In northern Victoria Land, at the northern end of the Transantarctic Mountains, two very low-grade units occur outboard of the main Ross orogenic belt, the Bowers and Robertson Bay units (Fig. 1). These units were recognized already as specific geologic entities during the first scientific exploration of the area (Sturm & Carryer 1970; Dow & Neall 1972). However, since they were first designated as terranes (Bradshaw et al 1985«), there has been an ongoing discussion as to whether they are fartravelled 'exotic' terranes (Bradshaw et al. 1985&) or whether they are accreted units related directly to the subduction process that formed the magmatic arc (Kleinschmidt & Tessensohn 1987). This discussion is of principal importance (Fig. 2) and relevant also for other active margin assemblages.

The terranes of northern Victoria Land According to the tectonostratigraphic definition, a terrane is a fault-bounded package of rock with a distinct stratigraphy that characterizes a

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peculiar geological setting, i.e. it forms a unique piece in the orogenic puzzle (Howell 1995). In this purely descriptive sense the term can be applied well to the northern Victoria Land units (Fig. 3).

Wilson 'terrane' The inboard Wilson 'terrane' (WT) is built up mostly by metamorphic rocks of sedimentary origin. Rare amphibolites can be interpreted as evidence for subordinate synsedimentary basin volcanism. The basement rocks are characterized by variable metamorphic degrees of low-P, and low- to high-T character. Lowgrade metasedimentary units are represented by rather monotonous turbiditic series as well as more variegated shallow-water sequences (Ricker 1964; Tessensohn et al. 1981; Skinner et al. 1996; Henjes-Kunst & Schussler 2003). Neoproterozoic to Cambrian passive Gondwana margin sequences, as known from the Adelaide geosyncline in SE Australia (e.g. Preiss 1987), or Cambrian platform carbonates, as found towards the central Transantarctic Mountains (e.g. Stump 1995), are unknown.

Fig. 2. Schematic cross-sections of the two alternative possibilities for the nature of the northern Victoria Land terranes: (a) autochthonous units at an accretionary margin (textbook example); (b) collage of exotic terranes. Example from the North American Pacific margin after Monger et al. (1985).

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Fig. 3. Terrane map of northern Victoria Land, Antarctica. Location as in Figure 1.

High- to very high-grade metamorphic rocks, interpreted in parts as meta-turbidites (Schiissler et al. 1999) comprise high-amphibolite to granulite-facies gneisses and migmatites (Talarico et al 1998; Schiissler et al 1999, 2004). It has been

assumed widely that a substantial amount of the parent rocks are Precambrian in age (e.g. Gair et al. 1969; Gibson & Wright 1985; Borg et al 1987; Talarico et al 1995). However, no definite proof of this assumption has been found.

NORTH VICTORIA LAND TERRANES Instead, SHRIMP U-Pb dating of zircons (Ireland et al 1999; Henjes-Kunst et al 2004) and Ar-Ar laser probe dating of detrital mica (Henjes-Kunst 2003) support a Lower to Middle Cambrian sedimentation age for the protoliths of the basement rocks. The low-grade rocks of the Berg Group and Rennick Schist contain detrital muscovites, the youngest of which are 511 Ma and 505 Ma in age (Henjes-Kunst 2003). Nd model ages range from 1.6-2.2 Ga for the western WT to 1.4-1.9 Ga for the eastern WT (Henjes-Kunst & Schiissler 2003). Age data of detrital minerals and geochemical and isotopic data indicate a dominant sedimentary input from an evolved source with an, on average, Early Proterozoic crustal formation age but late Neoproterozoic to Cambrian crystallization ages. Towards the east, an increasing influence of a juvenile source component is indicated. The WT basement forms the host unit for the Granite Harbour Intrusives (Gunn & Warren 1962) of the Ross orogenic magmatic arc (Borg et al 1987; Borsi et al 1987; Ghezzo et al 1987; Kreuzer et al 1987; Tonarini & Rocchi 1994; Rocchi et al 1998). Along the north coast between the Rennick and Mertz glaciers, the terrane is about 1000 km wide, if the outcrops of Ross age granitoids in George V Land (Fanning et al 2002) are considered. This enormous width, particularly of the hosted magmatic arc poses a problem which will be discussed later. Towards the south, the terrane continues without major break to at least the area of the Dry Valleys in South Victoria Land. However, since there are no outer terranes present, the term Wilson terrane is not used in that area.

Bowers terrane The Bowers terrane (Fig. 3) occurs between the inboard Wilson 'terrane' and the outboard Robertson Bay terrane. The terrane is complex and consists of three rather different sequences (Fig. 4). These are discussed in stratigraphical order. (1) There is a partly spilitic volcanic sequence of lavas, pillow lavas (Fig. 5a) and agglomerates (Glasgow Formation: Laird & Bradshaw 1983) of more than 2000m thickness. The lavas vary in composition between olivine-basalt and andesite. There are also some small gabbroic bodies. Geochemically, the volcanics are of variable, primitive island-arc, forearc or back-arc character (Weaver et al 1984; Wodzicky & Robert 1986; Rocchi et al

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2003; Estrada & Jordan 2003). They indicate an oceanic rather than continental crust underneath. The marine character of the succession is indicated by the presence of pillow lavas and interbedded greywackes and unfossiliferous limestones. Because of their strong alteration the rocks, so far, have not yielded any meaningful isotopic ages. The volcanic succession interfingers with several 100 m of mudstones, sandstones and conglomerates (Molar Formation), which have yielded Middle Cambrian trilobites (Cooper et al 1990; Wolfart 1994). The sequence is highly irregular, with numerous intercalations of matrix-supported conglomerates in a background sedimentation of mudstone and turbidites, the latter characterized lithologically by a pronounced volcanic component (Casnedi & Di Giulio 2003). Geochemical and isotopic investigations indicate that the siliciclastic Molar Formation represents a strongly heterogeneous mixture of a primitive basaltic source and a continental component similar to that inferred for the WT and RBT (Henjes-Kunst & Schiissler 2003). Nd model ages range from 1.7 Ga to 2.1 Ga. Ar-Ar laser probe dating of detrital muscovites shows a dominant age population in the range 720-520 Ma, with the youngest grains dated to 500 Ma (HenjesKunst 2003). SHRIMP dating reveals a crystallization age of 511 Ma for zircon of a granitic clast in a conglomerate horizon (Bassett et al 2002). The sedimentary environment appears to be that of a fairly unstable submarine slope. (2) In some areas, the Molar Formation grades into the fossiliferous Mariner Group (Fig. 5b). This is a marine regressive sequence from mudstones through limestones to sandstones (Andrews & Laird 1976). The age of the sequence spans the Middle to Upper Cambrian boundary. I-type granitic clasts of a conglomerate interpreted to form a member of the uppermost Mariner Group yielded SHRIMP zircon ages of 500-510 Ma (Bassett et al 2002). (3) The top of the succession is formed by quartz-conglomerates and cross-bedded quartzites (Fig. 5c) of the Leap Year Group. These rocks are rather pure continent-derived fluvial to deltaic sediments with a thickness of between 2300 m and 4000 m. Trace fossils are common and Scolithus, Daedalus, Arthrophycus, Monocaterion and Rusophycus have been mentioned

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Fig. 4. Compilation diagram of accretion processes in northern Victoria Land during Ross age active margin tectonics. Data are taken from a number of sources as cited in the text. Hatched columns indicate age range of detrital micas found in turbidites from all terranes. Numbers on top of columns give the age of the youngest micas found. Note: BT data are from mid-Cambrian Sledgers Group; no data available from upper Cambrian Leap Year Group. Dates in boxes in BT give ages of granitic clasts (Bassett et al 2002) from conglomerates. E, eclogite; T, tiger-layered gabbro; S, Surgeon Island granite.

in particular (compiled in GANOVEX Team 1987). The mature sediments are similar in character to the North African Palaeozoic quartzitic series. Comparable sequences on Gondwana are the Cambrian quartzites of the Ellsworth Mountains, the Pensacola Mountains, Shackleton Range, the Falkland Islands and the Sierras Australes (Buggisch 1987). The stratigraphic age of the Leap Year Group is inferred to be Late Cambrian. No isotopic source information is available at present for the unit.

The contrast between magmatic arc-related primitive volcanic rocks and continent-derived quartz-rich sandstones within the same terrane is striking. The importance of this change in sedimentary environment has been stressed recently by Casnedi & Di Giulio (2003) and Weaver et al. (2003).

Robertson Bay terrane The Robertson Bay terrane (Fig. 3) is the outermost terrane of the collage in northern Victoria Land. It consists of turbidites, derived mainly

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Fig. 5. Selected rock types from the Bowers terrane from bottom to top: (a) pillow lavas of the primitive island-arc Glasgow Volcanics, Bowers Mountains; (b) folded Middle Cambrian limestones from Reilly Ridge, near Lanterman Range; limestone succession is about 100 m thick; (c) cross-bedded terrestrial quartzites of the Late Cambrian Leap Year Group.

from continental source rocks of, on average, Palaeoproterozoic formation age (e.g. HenjesKunst & Schiissler 2003). Lithologically, there is a rather complete lack of volcanic components in the sandstones. The estimated total thickness

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of the succession is in the order of several kilometres. The metamorphic grade is fairly uniform, from very low (trace fossils preserved) to low with a slight increase towards the interior western boundary (Buggisch & Kleinschmidt 1991). The rocks are folded in a rather regular chevron style manner (Crowder 1968; Kleinschmidt & Skinner 1981; Findlay & Field 1983; Findlay 1986; Kleinschmidt 1992). The base of the sequence is not exposed, but there are some indications that the underlying crust might be continental in character. In granitoids of a later generation - the post-Ross orogenic Admiralty Granites - Borg et al (1987) have found an eastward increasing continental Nd-isotope signature. Well within the RBT, there is an enigmatic sheared granite on Surgeon Island off the north coast of northern Victoria Land (Tessensohn et al. 1996 with earlier references). Its crystallization age is 511 Ma (Fioretti et al 20010). The age pattern of inherited zircon components in the metagranite is different from those of the RBT metasediments (Fioretti et al 2002). A likely interpretation is that the Surgeon Island metagranite presents a fragment of the Ross-age basement of the RBT (Fioretti et al 20035). Alternatively, it may be interpreted as a remnant of an exotic continental sliver (Borg & De Paolo 1991). At Handler Ridge near the boundary to the BT, there occur exotic limestone blocks as part of the succession. These blocks contain a fauna of conodonts, crinoids, trilobites and brachiopods of earliest Ordovician (Tremadoc) age (Wright et al 1984; Burrett & Findlay 1984; Buggisch & Repetski 1987). So far this is the only known outcrop of marine Ordovician rocks in the Antarctic. According to the results of SHRIMP dating of detrital zircon (Fioretti et al 20030) and Ar-Ar laser probe dating of detrital muscovite (Henjes-Kunst 2003), which yielded minimum ages of c. 500-480 Ma, a maximum Early Ordovician sedimentation age is documented also for other parts of the RBT. Similarly to the WT, the age data of detrital minerals indicate a dominant sedimentary infill from crustal rocks with late Neoproterozoic to Cambrian crystallization ages. Nd model ages vary only between 1.7 Ga and 1.9 Ga (HenjesKunst & Schiissler 2003).

Internal deformation The deformation in the WT is complex and not uniform. However, a regional schistosity, which forms the axial plane foliation of isoclinal folds, strikes NW-SE. Polyphase deformation of two

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to four phases is common in the high-grade realms (Skinner & Ricker 1968; Kleinschmidt & Skinner 1981; Bradshaw et al 1982, Sandiford 1985; Kleinschmidt et al 1984). The main deformation in the WT is coupled closely with metamorphism (metamorphic peak deformation). The WT is also affected by a number of thrusts, the Exiles and Wilson thrusts (Flottmann & Kleinschmidt 1991) in the north and the Boomerang thrust (Skinner 1991) in the south. The thrusts strike NW-SE and involve tectonic transport to the west, opposite to the thrust system along the terrane boundaries. Hightemperature deformation of pegmatites which cross-cut the Exiles and Wilson thrusts, indicates that thrusting occurred during high-grade Ross-orogenic metamorphism (Laufer & Rossetti 2003). This is also indicated by Ar-Ar dating of igneous and metamorphic mica (Schiissler et al 2004). Away from the terrane boundaries, the RBT and BT show a comparable style of deformation, i.e. a single phase of upright folding with associated penetrative cleavage. In the RBT, the sub-horizontal undulating axes trend NW-SE, with a slight angle to the terrane boundary. The regular style of upright chevrontype folding in the RBT requires a detachment horizon underneath (Kleinschmidt 1992). In the BT, folding and cleavage are obvious particularly in the well-bedded turbidites. The thick quartzites at the top of the BT sequence form a box-shaped syncline with vertical limbs, more than 100 km long. In the Mariner Glacier area, the syncline is accompanied by an anticline on the NE side (Tessensohn 1984; GANOVEX Team 1987). The terrane boundaries The contact between the Wilson terrane and rocks of the East Antarctic craton (i.e. the structural internal boundary of the WT) is not exposed, but appears to be rather abrupt. The nearest (Precambrian) cratonic rocks occur at the Mertz Glacier in George V Land (Oliver & Fanning 2002). A mylonitic shear zone in the same area (Talarico & Kleinschmidt 2003) is part of the cratonic basement. The following outcrops to the east are all in WT basement and Granite Harbour Intrusives arc granitoids. No such granitoids occur in the Precambrian cratonic realm. A major structural break in the WT may occur in the Matusevich Glacier area. According to aeromagnetic data (Ferraccioli et al 2003), there may be a mafic complex at depth comparable in signature to the Bowers Zone. The Matusevich Glacier area is also the location

of the Exiles thrust (Flottmann & Kleinschmidt 1991). The rocks on either side of this zone are similar and both the western and the eastern part of the Wilson terrane contain granitoids of the early Palaeozoic magmatic arc. All other terrane boundaries are distinct structural features. Both the WT/BT and the BT/RBT borders are characterized by the occurrence of schist belts, which are several kilometre wide (Fig. 3). The metamorphic grade is of greenschist facies (distinct from the very low to low grade of the interior BT and RBT) and the schist zones are marked by a second phase of deformation (Fig. 6), distinct from the monophase deformation of the interior of both BT and RBT terranes. In most cases, there appears to be outward-directed, west over east tectonic transport (Tessensohn 1984; GANOVEX Team 1987; Palmeri et al 1994), sometimes with a dextral component (Capponi et al 2003).

Fig. 6. Example from the schist belt between Bowers terrane and Robertson Bay terrane. Note two phases of deformation. Millen Range.

NORTH VICTORIA LAND TERRANES

WT/BTboundary The WT/BT boundary is a first-order tectonic contact, which is well exposed between Lady Newnes Bay of the Ross Sea and the Lanterman Range. North of this range, it is hidden beneath the Rennick Glacier. The tectonic nature of the structure is that of a major shear zone. Conglomerates show strongly flattened pebbles, tonalitic rocks contain sheared inclusions and there are cases (Mariner Plateau, Capponi et al 2003) where slivers of WT tonalite are incorporated tectonically in BT schists. Tectonic transport along the fault zone in this area is mainly orthogonal with a dextral component. In addition, the WT/BT boundary is marked by a number of special features, which provide information on the nature of the fault zone (Capponi et al. 2003). First, there are two strongly sheared conglomerate bodies in the Lanterman Range along the fault zone. The interior one consists of basement clasts, which are not only flattened but, in some cases, also folded. The adjacent mafic conglomerate has a clast population of about 95% boninite, which is a lava composition typical for primitive forearc volcanism. Other conglomerates occur close to the Ross Sea along the same fault (Gibson et al 1984). Slivers of ultramafic rocks are also found at the Ross Sea coast and close to the fault in the Lanterman Range. These include a layered gabbro intrusion near the mouth of the Mariner Glacier (Tiger gabbro, GANOVEX Team 1987) and eclogites in the Lanterman Range (Ricci et al 1996). Along the fault zone, but still within the WT, there are indications of high-pressure metamorphism (Grew et al 1984), but the most important discovery is that of the high-pressure quartz modification coesite which was recently found in eclogite of the Lanterman Range (Ghiribelli et al 2001; Palmeri et al 20030, b). The host rocks of the WT must have been at a depth of at least 90 km during the time of formation of this mineral. Petrological investigations (Talarico et al 1998) further reveal that the Lanterman Range represents a stacked sequence of basement units, which were metamorphosed under increasing pressure conditions towards the WT/BT boundary. Geochronological investigations constrain an age close to 500 Ma for peak-metamorphic conditions in the basement units of the Lanterman Range including the ultra-high pressure metamorphic rocks (Goodge et al 1995; Goodge & Dallmeyer 1996; Di Vincenzo et al 1997; Di Vincenzo & Palmeri 2001). These findings strongly support the interpretation of

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the fault zone as major suture and trace of a subduction zone (Ricci & Tessensohn 2003). The fact that rocks of the Granite Harbour Intrusives magmatic arc occur very close to the subduction zone indicates that this cannot be the primary setting and that the outboard part of the original WT crust was removed by later tectonic processes. In the fault segment close to the Ross Sea, there are two additional slivers of distinct rock types between WT and BT that are too small to be counted as separate terranes. These are the Dessent Schists, an amphibolite-facies intermediate-pressure metamorphic unit (Kleinschmidt et al 1984; Talarico et al 1998) and, adjacent to it, a small tectonic wedge of Rossorogenic tonalitic rocks with strong internal deformation (sc-c structures) in places (Musumeci et al 1994). In this area, the BT has a schist margin showing a second deformation (Gibson et al 1984). In the same area, there is also the strongest evidence for east-directed tectonic transport on thrust faults and reversed faults (Tessensohn 1984).

BT/RBTboundary The schists at the BT/RBT boundary are generally quartz-rich, indicating RBT turbidites as possible protoliths. In a few places, however, there occur also volcanogenic rocks, i.e. in the Millen Range as mafic volcanics at Crosscut Peak (Findlay 1992) and pillow lavas on the second ridge NW of Turrett Peak (authors' observation) and in the Barber Glacier area near the north coast of northern Victoria Land as serpentinite and other mafic rocks (Kleinschmidt 1992; Capponi et al 19990, b, c). This indicates that the schists contain also protoliths of BT rocks. A good example of the complicated field relations of this boundary is found in the area of Crosscut Peak in the Millen Range (Wright & Dallmeyer 1991; Findlay 1992; Capponi et al 1994), where northeast-directed thrusting is obvious, but where lithological units, phases of deformation and metamorphic grades do not fit a simple explanation scheme.

Implications The WT/BT boundary appears to consist of two major elements, the main suture and an undisputable trace of a subduction zone within the WT and the greenschist belt on the western boundary of the BT. The subduction zone must also have been responsible for the formation of the Granite Harbour magmatic arc, which is restricted to the Wilson terrane.

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The greenschist-metamorphic shear zone at the western BT boundary, as well as the schist belt between BT and RBT, mark the accretion of the two outer terranes to the orogenic collage by a process of underplating by which the leading edge of the WT was thrustfaulted over the BT and RBT. This event occurred later than the formation of the magmatic arc because some of the magmatic-arc granitoids (tonalite belt) were sheared during the process (Musumeci et al 1994; Musumeci 1999). The BT/RBT boundary forms a reversed fault or thrust between the two units. As an overburden of 10-15 km is required to explain the metamorphic grade in the RBT, Buggisch & Kleinschmidt (1991) suggested that the higher thrust units have formed this overburden tectonically. This relates the metamorphism in the RBT directly to the coupling process.

Age constraints The age spectra of detrital mica (hatched columns on Fig. 4) from turbidites of all three terranes (Henjes-Kunst 2003) show a dominant age population in the range of 720-520 Ma. The dates on top of the columns give the age of the youngest reworked mica in each sampled unit. The gradual younging of the reworked material from inboard to outboard is in agreement with a gradual accretion process. The presence of late Neoproterozoic to Early Cambrian material in all age spectra indicates a close spatial relationship of the various turbidite basins. It is obvious that a significant amount of the detrital mica comes from a source up to 150 Ma older than the main Ross event. This is supported by SHRIMP age data of detrital zircon from metasediments of the three terranes (Ireland et al. 1999; Fioretti et al. 20030; Henjes-Kunst et al. 2004). Nd model ages (Henjes-Kunst & Schtissler 2003) also indicate a common crustal source with model ages of 1.6-1.9 Ga common to all terranes (Fig. 4, bottom). The upper limit of accretion is well defined geologically by overlap sequences and stitching plutons. However, these rocks are so much younger that they do not provide tight age constraints for the accretion events. Overlap sequences are the subaerial Gallipoli Volcanics of the Devonian-Carboniferous boundary, which occur in isolated patches on all three terranes (Fioretti et al. 2001 b). The DevonianCarboniferous Admiralty granites, which are coeval with the Gallipoli Volcanics (HenjesKunst & Kreuzer 2003), form stitching plutons occurring on all three terranes and, in a few cases, straddling the terrane boundaries

(GANOVEX Team 1987). The Permo-Triassic Gondwana sequence forms another, more widespread overlap sequence on the Wilson and Bowers terranes (GANOVEX Team 1987).

Discussion It is easy to decide whether active margin assemblages are exotic terranes or part of the local orogenic collage if the rocks of the terrane units differ substantially in age from each other and from the inboard continental edge. If this is not the case, then the analysis becomes very difficult.

Autochthonous terranes? First consider the case of an active margin setting. The magmatic arc in the Wilson terrane is the strongest argument for this scenario. The intrusions of the granitoids took place over some time (Fig. 4), with a peak activity around 500 Ma. Regional metamorphism with peak ages also around 500 Ma (Henjes-Kunst et al. 2004 and references therein) indicates the close temporal relationship of subduction-related mantle melting and metamorphic heating processes of the crust. Products of this welldefined magmatic arc-forming episode are found all along the Transantarctic Mountains (e.g. Stump 1995) and the event is one of the main characterizing features of the Ross Orogen. Additional arguments for the subduction scenario are found associated with the suture zone at the outboard margin of the WT. Highpressure metamorphism, ultramafic slivers and the extreme pressure and burial values of the eclogites and the coesite relics in them are particularly significant. This boundary is not just a strike-slip accretion fault; it must be regarded as ultimate proof for a deep-reaching subduction process. The total width of about 1000 km of Ross-age plutonism in northern Victoria Land, Oates Land and George V Land, however, poses a severe problem to the concept of the Wilson terrane forming the active continental margin of the Ross Orogen. This would indicate a rather unusual subduction process. The situation led Bradshaw (1987) to the assumption that part of the unit is a terrane coming from somewhere else along the same active margin. However, Bradshaw's suggested interior terrane boundary, the Aviator-Rennick line, reduces the total width of the WT only slightly. To the west, there remain still about 900 km of WT and Granite Harbour magmatic-arc granitoids.

NORTH VICTORIA LAND TERRANES Recent investigations of the geology in the area of the Mertz Glacier, George V Land indicate a major crustal boundary separating Ross-age continental crust to the east from Palaeoproterozoic cratonic crust to the west (Fanning et al 2002; Henjes-Kunst et al pers. comm.). Mantle fragments in metasedimentary basement rocks of the Matusevich Glacier area, Gates Land may indicate another suture zone separating Ross-age crustal blocks (Roland et al 2001). Other possible terrane boundaries remain to be shown. But what is the role of the two outer terranes in this scenario? The Nd model ages of the crustal source (Henjes-Kunst & Schiissler 2003) for all three terranes are very similar (Fig. 4). Leaving aside the terrestrial quartzites in the upper BT for the moment, both terranes can be taken as part of a marine oceanic realm. In the active margin scenario, the BT might represent an accretionary wedge of a subduction complex, the RBT a turbidite basin of the trench. In the accretionary wedge, volcanic rocks of the BT would have been scraped off the down-going ocean floor. This would account for the mixed character of the volcanic rocks, the irregular sedimentation, the amount of exotic material and conglomeratic bodies and the common reworked turbidite components in these conglomerates as well as for some tectonic repeating as found on Reilly Ridge. However, the subduction zone responsible for the accretionary wedge is too close and times of sedimentation and accretion are too young to be related to the magmatic-arc forming process of the Granite Harbour Intrusives. In the RBT, the continent-derived turbidites lack volcanic components. This seems to indicate that the transport from the source did not pass through the accretionary wedge, but in some way went around it and then became arcparallel. However, the continental signature deduced for the underlying crust seems to exclude the trench position. If one assumes a position in a forearc basin instead, this would be on the wrong side of the accretionary wedge. It appears that the accretionary wedge activity would have come to a stop during the regressive Mariner Group sedimentation. The overlying terrestrial top sequence of the BT (Leap Year Group) must come from a large continent, as indicated by maturity and fluvial depositional features. It could indicate the transgression of a river delta over the accretionary wedge, with the river(s) coming from the interior of Antarctica or Gondwana. The discussion shows that there are a number of problems associated with the active margin

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scenario, particularly in the outer terranes, although there is direct evidence for a subduction zone. The possible positions in a subduction complex are not really convincing and the time of sedimentation does not coincide with the arc formation.

Exotic terranes In an exotic terrane setting, the first event would have been the formation of an oceanic forearc to back-arc environment in the Middle Cambrian, now represented by the Glasgow Volcanics of the BT. One can imagine easily an oceanic subduction setting, e.g. of the South Sandwich Islands type, producing such a feature. Boninites, slope sediments, conglomerates, and limestone fringes could all be associated with such an oceanic island chain. However, continental components in the BT sediments and granitic clasts in the conglomerates argue strongly for nearby continental crust experiencing strong exhumation and denudation. The terrestrial Leap Year sequence of the BT (Figs 4, 5c) can be derived only from a large continent. The similarity with other coeval terrestrial series in Ellsworth Mountains, Pensacola Mountains, Shackleton Range, Falkland Islands, Cape Town area and Argentina supports the assumption of interior Gondwana as the source area. In the exotic terrane scenario, this means that BT must have been coupled with WT before the beginning of the Leap Year sedimentation. Otherwise, the deposition of the terrestrial quartzites onto the BT primitive magmatic-arc terrane in an exotic area elsewhere is difficult to explain. A WT/BT coupling after Mariner Group sedimentation in the Middle Cambrian would coincide with the magmatic-arc activity on the WT. However, a BT/RBT coupling during this interval is impossible. As shown by the fossils in the exotic limestone blocks and by the ages of detrital minerals, sedimentation in the RBT lasted until the Early Ordovician. Folding of the rocks in the terrane can have occurred only later. Accretion tectonics in the schist belt of the RBT/BT boundary was even subsequent to that.

Time frame For the interpretation of the sedimentary, plutonic and tectonic processes a combination of stratigraphic and isotopic dates has to be used. Stratigraphic markers are found mainly in the two outboard terranes, whereas isotopic data are available mainly for the inner WT. To link the isotopic data with the stratigraphic

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results, an up-to-date time-scale is the most important tool. The recent changes in the Cambrian time-scale (GTS 2004) have been of considerable help, as a number of former contradictions can now be resolved. The dates are used in the synoptic table of events (Fig. 4). The changes must also be considered when dealing with earlier interpretations, including our own (Kleinschmidt & Tessensohn 1987). For example, an age younger than 530 Ma for the base of the Late Cambrian seemed improbable 15 years ago (Bradshaw 1989), whereas the same horizon is now considered to lie close to 500 Ma. Magmatic-arc plutonism as well as metamorphism in the WT have their peaks around 500 Ma. According to the new timescale, this is at the MC-LC boundary. Age data of the eclogite in the Lanterman Range suggest a similar age (500 Ma) for the maximum burial in the subduction zone. The ages of detrital mica in the low-grade units of the WT terminate before the 500 Ma peak. The stratigraphic ages in sediments of the BT and RBT are based on conodonts (RBT) and trilobites (BT) and are, therefore, rather precise (Fig. 4). The Early Ordovician age of sedimentary exotic blocks in the RBT puts a limit to the deformation event in RBT and also BT. As the deformation is monophase in both terranes and of comparable style, it must be later than the limestone age, i.e. later than the Tremadoc, probably in a range between 480 Ma and 485 Ma. Also, the deformation at the BT/RBT boundary must be younger than Tremadoc as it is subsequent to the folding event. Possibly, the boundary event is manifest in the second phase of schist belt deformation (s2) (Wright & Dallmeyer 1991). All isotopic deformation ages older than Tremadoc - and there is quite a number of those in the RBT are meaningless geologically. The cause for this seems to lie more in the unsuitable material (mixture of detrital and newly grown micas) rather than in the methods, as indicated by the fact that the Ar/Ar ages (e.g. Wright & Dallmeyer 1991) are not very different from earlier K-Ar ages (e.g. Adams & Kreuzer 1984). The inferred Late Cambrian stratigraphic age for the terrestrial sediments of the BT gives a less precise upper limit for the following deformation, but the RBT docking event at closely before 480 Ma could also have been responsible for docking of the BT. The most important conclusion is the existence of two main events in northern Victoria Land (Fig. 4) - subduction and magmatic-arc formation at around 500 Ma and terrane accretion at around 485 Ma. This sequence of

events allows combination of both the subduction process (Kleinschmidt & Tessensohn 1987) and the terrane accretion (Bradshaw et al. 1985Z?) scenarios. The enormous width of the Ross-age crust from the Rennick Glacier area in northern Victoria Land to the Mertz Glacier in George V Land makes it unlikely that subduction and magmatic-arc forming process took place in the present relative position of the orogen. If, however, the WT is a terrane, as the lack of a transition at the interior boundary suggests, then it might as well consist of two terranes with a boundary at the Matusevich Glacier. Part of the magmatic arc, formed at around 500 Ma elsewhere along the Gondwana active margin, would then have been sheared off and moved as a terrane to its present position. A duplication of this segment at the northern Victoria Land margin would explain the enormous width of terrane and magmatic arc; it would be simply two segments of the same active margin. Some proof of this scenario could be given if the age for the younger accretion event could be found in the Matusevich Glacier shear zone. Both outer terranes could have a different origin and could have been added to the collage at around 485 Ma. In support of this assumption, Ar-Ar dates on phengites indicate a 486 Ma age for the greenschist metamorphism in the area of the Lanterman Range (Goodge & Dallmeyer 1996; Di Vicenzo & Palmeri 2001). This is the preferred option at present. Alternatively, the accretion of the BT may have occurred earlier, which would provide an explanation for the formation of the lid of terrestrial sediments on the BT. The accretion could then have taken place in the following, more complicated, but not impossible process: (a) oceanic magmatic-arc formation; (b) BT terrane accretion in the early Late Cambrian without any obvious deformation phase; (c) fluviatile terrestrial sedimentation; and (4) coupling of the RBT in the Early Ordovician and reactivation of the WT/BT boundary, including folding of the Leap Year sediments. Data on detrital minerals in the Leap Year sediments could help in the future to decide on the options.

Conclusions In the early Palaeozoic Ross Orogen of northern Victoria Land, Antarctica, there is evidence for both subduction and terrane accretion processes. Main evidence for subduction is the magmatic arc in the WT and the relict subduction zone at the eastern WT margin. Both were contemporaneous at about 500 Ma.

NORTH VICTORIA LAND TERRANES Evidence for terrane accretion is present mainly in the outboard terranes which have younger sedimentation ages up to 485 Ma, followed by a deformation event between 485 Ma and

480 Ma. The inner WT may be autochthonous, but then the Ross-age magmatic arc would have an enormous width. Alternatively, the WT may have formed by subduction in another segment of the active margin of Gondwana. Sheared off and transported to the present position, there may have been a split and tectonic duplication of the segment so that the WT today consists of at least two terranes, an eastern terrane from the Lanterman Range to the Wilson Hills east of the Matusevich Glacier and a western terrane from Matusevich Glacier to the Mertz Glacier. The outer terranes seem to originate from truly exotic areas, from a primitive forearc to back-arc magmatic arc in the case of the BT and from a continental margin in the case of the RBT. Final terrane accretion occurred about 15 Ma later than the formation of the granitic arc. There is evidence that this coupling event affected both outer terranes simultaneously (schist belts at the terrane boundaries) between

485 Ma and 480 Ma. It is possible, however, that the BT was accreted earlier at the MC-LC boundary (close to 500 Ma), because a coupling at that time would provide a better explanation for the continent-derived mature sandstones that form the uppermost unit of the BT. Proof of c. 485 Ma old shear zones in the Matusevich Glacier area and dates on detrital micas and zircons in the Leap Year Group could help to decide between the options in the future.

References ADAMS, CJ. & KREUZER, H. 1984. Potassium/argon age studies of slates and phyllites from the Bowers and Robertson Bay terranes, North Victoria Land, Antarctica. Geologisches Jahrbuch, B60, 265-288. ANDREWS, PB. & LAIRD, M.G. 1976. Sedimentology of a Late Cambrian regressive sequence (Bowers Group), northern Victoria Land, Antarctica. Sedimentary Geology, 16, 21-44. BASSETT, K.N., WEAVER, S., BRADSHAW, J.D. & IRELAND, T. 2002. Dating of the accretion of the Cambrian intra-oceanic arc, Bowers Terrane, northern Victoria Land, Antarctica. Abstract, Gondwana 11, Correlations and Connections. University of Canterbury, Christchurch. BORG, S.G. & DE PAOLO, DJ. 1991. Crustal structure and tectonics of the Antarctic margin of Gondwana and implications for the tectonic

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development of southeastern Australia. Tectonophysics, 196, 674-685. BORG, S.G., STUMP, E., CHAPPELL, B.W., MCCULLOCH, M.T., WYBORN, D., ARMSTRONG, R.L. & HALLOWAY, J.R. 1987. Granitoids of northern Victoria Land, Antarctica: implications of chemical and isotopic variations to regional crustal structure and tectonics. American Journal of Science, 287,127-169. BORSI, L., FERRARA, G. & TONARINI, S. 1987. Rb-Sr and K-Ar data on Granite Harbour Intrusives from Terra Nova Bay and Priestley Glacier (Victoria Land, Antarctica). Memorie delta Societa Geologica Italiana, 33,161-170. BRADSHAW, J.D. 1987. Terrane boundaries and terrane displacement in northern Victoria Land, Antarctica: some problems and constraints. In: LEITCH, E. & SCHEIBNER, E. (eds) Terrane Accretion and Orogenic Belts. American Geophysical Union, Geodynamic Series, 19, 207-219. BRADSHAW, J.D. 1989. Terrane boundaries in North Victoria Land. Memorie delta Societa Geologica Italiana, 33, 9-16. BRADSHAW, J.D., LAIRD, M.G. & WODZICKY, A. 1982. Structural style and tectonic history in northern Victoria Land, Antarctica. In: CRADDOCK, C. (ed.) Antarctic Geoscience. University of Wisconsin Press, Madison, 809-816. BRADSHAW, J.D., BEGG, J.C., BUGGISCH, W., BRODIE, C., TESSENSOHN, F. & WRIGHT, TO. 19850. New data on Paleozoic stratigraphy and structure in North Victoria Land. New Zealand Antarctic Record, 6 (3), 1-6. BRADSHAW, J.D., WEAVER, S. & LAIRD, M.G. 19856. Suspect terranes and Cambrian tectonics in northern Victoria Land, Antarctica. In: HOWELL, D.G. (ed.) Tectono-Stratigraphic Terranes of the Circum-Pacific Region. Circum-Pacific Council for Energy and Mineral Resources, Earth Sciences Series 1, 467-479. BUGGISCH, W. 1987. Stratigraphy and very low-grade metamorphism of the Sierras Australes de la Provincia de Buenos Aires (Argentina) and implications for Gondwana correlations. Zentralblattfiir Geologic und Palaontologie, 1, 819-837. BUGGISCH, W. & KLEINSCHMIDT, G. 1991. Recovery and recrystallization of quartz and 'crystallinity' of illite in the Bowers and Robertson Bay terranes, northern Victoria Land, Antarctica. In: THOMSON, M.R.A., CRAME, J.A. & THOMSON, J.W (eds) Geological Evolution of Antarctica. Cambridge University Press, Cambridge, 155-160. BUGGISCH, W. & REPETSKI, J.E. 1987. Uppermost Cambrian (?) and Tremadocian conodonts from Handler Ridge, Robertson Bay Terrane, Antarctica. Geologisches Jahrbuch, B66,145-185. BURRETT, C.F. & FINDLAY, R.H. 1984. Cambrian and Ordovician conodonts from the Robertson Bay Group, Antarctica and their tectonic significance. Nature, 307, 723-725. CAPPONI, G, CAROSI, R., MECCHERI, M & OGGIANO, G. 1994. Strain determination and tectonic evolution of the Millen Range area (northern Victoria Land, Antarctica). Terra Antartica, 1, 63-66.

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CAPPONI, G., CRISPINI, L. & MECCHERI, M. 19990. The metaconglomerates of the eastern Lanterman Range (northern Victoria Land, Antarctica): new constraints for their interpretation. Antarctic Science, 11, 217-227. CAPPONI, G., CRISPINI, L. & MECCHERI, M. 19996. Structural history and tectonic evolution of the boundary between the Wilson and Bowers terranes, Lanterman Range, northern Victoria Land, Antarctica. Tectonophysics, 312, 249-266. CAPPONI, G, MECCHERI, M., PETRINI, R. & VILLA, I.M. 1999c. The meta-igneous rocks from the Barber Glacier area (northern Victoria Land, Antarctica): a clue for the enigmatic Borchgrevink orogeny? Proceedings of the 8th International Symposium on Antarctic Earth Sciences, Abstracts, 63, Royal Society of New Zealand, Wellington. CAPPONI, G, KLEINSCHMIDT, G, PERTUSATI, PC., RICCI, C.A. & TESSENSOHN, F. 2003. Terrane relationships in the Mariner Glacier area of northern Victoria Land. In: TESSENSOHN, F. & RICCI, C.A. (eds) The Mariner Glacier Area, Antarctica - Aspects of a Suture Zone. Geologisches Jahrbuch, B85, 49-78. CASNEDI, R. & Di GIULIO, A. 2003. Sandstone petrography of the Bowers Supergroup and the Robertson Bay Group (Mariner Glacier area, Antarctica): preliminary results and remarks on the Paleotectonics. In: TESSENSOHN, F. & RICCI, C.A. (eds) The Mariner Glacier Area, Antarctica - Aspects of a Suture Zone. Geologisches Jahrbuch, B85,195-221. COOPER, R.A., BEGG, J.G. & BRADSHAW, J.D. 1990. Cambrian trilobites from Reilly Ridge, northern Victoria Land, Antarctica, and their stratigraphic implications. New Zealand Journal of Geology and Geophysics, 33, 55-66. CROWDER, D.F. 1968. Geology of a part of north Victoria Land, Antarctica. US Geological Survey Professional Paper, 600D, 95-107. Di VINCENZO, G. & PALMERI R. 2001. An 40Ar-39Ar investigation of high-pressure metamorphism and the retrogressive history of mafic eclogites from the Lanterman Range (Antarctica): evidence against a simple temperature control on argon transport in amphibole. Contributions to Mineralogy and Petrology, 141, 15-35. Di VINCENZO, G, PALMERI R., TALARICO F, ANDRIESSEN P.A.M. & RICCI C.A. 1997. Petrology and geochronology of eclogites from the Lanterman Range, Antarctica. Journal of Petrology, 38,1391-1417. Dow, J.A.S. & NEALL, V.E. 1972. Summary of the geology of the lower Rennick Glacier. In: ADIE, R. (ed.) Antarctic Geology and Geophysics. Universitaetsforlaget, Oslo, 339-344. ESTRADA, S. & JORDAN, H. 2003. Early Palaeozoic island arc volcanism in the Bowers Terrane of northern Victoria Land, Antarctica. Geologisches Jahrbuch, B95, 183-207. FANNING, C.M., MENOT, R.P., PEUCAT, JJ. & PELLETIER, A. 2002. New insights into the development of the early Palaeozoic Pacific margin of Gondwana on the Mawson Continent. Abstract, Gondwana

11, Correlations and Connections. University of Canterbury, Christchurch. FERRACCIOLI, F, DAMASKE, D, Bozzo, E. & TALARICO, F. 2003. The Matusevich aeromagnetic anomaly over Gates Land, East Antarctica. Terra Antartica, 10, 221-228. FINDLAY, R.H. 1986. Structural geology of the Robertson Bay and Millen terranes, northern Victoria Land, Antarctica. In: STUMP, E. (ed.) Geological Investigations in Northern Victoria Land, Antarctica. American Geophysical Union, Antarctic Research Series, 46, 91-114. FINDLAY, R.H. 1992. The age of cleavage development in the Ross orogen, northern Victoria Land, Antarctica: evidence from 40Ar/39Ar whole-rock slate ages: discussion. Journal of Structural Geology, 14, 887-890. FINDLAY, R.H. & FIELD, B. 1983. Tectonic significance of deformations affecting the Robertson Bay Group and associated rocks, northern Victoria Land, Antarctica. In: OLIVER, R.L., JAMES, PR. & JAGO, J.B. (eds) Antarctic Earth Science. Australian Academy of Science, Canberra, 107-112. FIORETTI, A.M., CAPPONI, G, BLACK, L.P., VARNE, R. & VISONA, D. 20010. Relationships between Surgeon Island and Robertson Bay Terrane: new U-Pb zircon and structural data. EUG XI, Strasbourg, European Union of Geosciences. Journal of Conference Abstracts, 6, 380. FIORETTI, A.M., BLACK, P. & VARNE, R. 20016. U-Pb zircon SHRIMP dating of the Gallipoli Volcanics, northern Victoria Land (Antarctica). EUG XI, Strasbourg, European Union of Geosciences. Journal of Conference Abstracts, 6, 380. FIORETTI, A.M., CAPPONI, G, BLACK, L.P., VARNE, R. & VISONA, D. 2002. Inherited zircon pattern of Surgeon Island granite: evidence for Proterozoic crust at the eastern end of northern Victoria Land. Abstract, Gondwana 11, Correlations and Connections. University of Canterbury, Christchurch. FIORETTI, A.M., BLACK, P., HENJES-KUNST, F. & VISONA, D. 20030. Detrital zircon age patterns from a large gneissic xenolith from the Cape Phillips granite and from the Robertson Bay Group metasediments, northern Victoria Land, Antarctica. 9th International Symposium on Antarctic Earth Sciences, Potsdam, Programme and Abstracts. Terra Nostra, 2003/4, 94-95. FIORETTI , A.M., CAPPONI, G, BLACK, L.P. & VISONA, D. 20036. Surgeon Island Granite: implications for regional geological evolution of northern Victoria Land from structural observations and SHRIMP zircon dating. Terra Antartica Reports, 9, 171-175. FLOTTMANN, T. & KLEINSCHMIDT, G. 1991. Opposite thrust systems in northern Victoria Land, Antarctica: imprints of Gondwana's Paleozoic accretion. Geology, 19, 45-47. GAIR, H.S., STURM, A., CARRYER, SJ. & GRINDLEY, GW. 1969. The geology of northern Victoria Land. Antarctic Map Folio Series, Folio 12, Plate XII. American Geographical Society, New York.

NORTH VICTORIA LAND TERRANES GANOVEX Team 1987. Geological map of North Victoria Land, Antarctica, 1:500 000, explanatory notes. Geologisches Jahrbuch, B66, 7-79. GHEZZO, C, BALDELLI, C, BIAGINI, R. ET AL. 1987. Granitoids from the David Glacier-Aviator Glacier segment of the Transantarctic Mountains, Victoria Land, Antarctica. Memorie delta Societd Geologica Italiana, 33,143-160. GHIRIBELLI, B., FREZOTTI, M.L. & PALMERI, R. 2001. Coesite in eclogites of the Lanterman Range (Antarctica): evidence from textural and raman spectroscopy studies. European Journal of Mineralogy, 14, 355-360. GIBSON, G.M. & WRIGHT, TO. 1985. Importance of thrust faulting in the tectonic development of northern Victoria Land, Antarctica. Nature, 315, 480-483. GIBSON, G.M., TESSENSOHN, F. & CRAWFORD, A. 1984. Bowers Supergroup rocks west of the Mariner Glacier and possible greenschist-facies equivalents. Geologisches Jahrbuch, B60, 289-318. GOODGE, J.W. 2002. From Rodinia to Gondwana: supercontinent evolution in the Transantarctic Mountains. In: GAMBLE, J.A., SKINNER, D.N.B. & HENRYS, S. (eds) Antarctica at the Close of a Millennium. Royal Society of New Zealand, Bulletin, 35, 61-74. GOODGE, J.W. & DALLMEYER, R.D. 1996. Contrasting thermal evolution within the Ross orogen, Antarctica: evidence from mineral 40Ar/39Ar ages. Journal of Geology, 104, 435-458. GOODGE, J., WALKER, N.W & DALLMEYER. R.D. 1995. Thermal and kinetic history of the Ross orogen, Antarctica. Geological Society of America, Abstracts with Programs, 27, 126. GREW, E.S., KLEINSCHMIDT, G. & SCHUBERT, W 1984. Contrasting metamorphic belts in North Victoria Land, Antarctica. Geologisches Jahrbuch, B60, 253-263. GTS (GEOLOGICAL TIME SCALE) 2004. International Commission on Stratigraphy, Paleozoic Comparison Chart, www.stratigraphy.org/paleu.pdf. GUNN, B.M. & WARREN, G. 1962. Geology of Victoria Land between Mawson and Mulock Glaciers, Antarctica. New Zealand Geological Survey, Bulletins, 71, 1-157. HENJES-KUNST, F. 2003. Single-crystal Ar-Ar laser dating of detrital micas from metasedimentary rocks of the Ross erogenic belt at the Pacific margin of the Transantarctic Mountains, Antarctica. 9th International Symposium on Antarctic Earth Sciences, Potsdam, Programme and Abstracts. Terra Nostra, 2003/4, 150-151. HENJES-KUNST, F. & KREUZER, H. 2003. Mid-Paleozoic igneous activity in northern Victoria Land, Antarctica: implications of new geochronological data. Geologisches Jahrbuch, B85, 271-302. HENJES-KUNST, F. & SCHUSSLER, U. 2003. Metasedimentary units of the Cambro-Ordovician Ross Orogen in northern Victoria Land and Gates Land: geochemical and Nd/Sr isotope comparison, implications for their provenance and geotectonic consequences. Terra Antartica, 10,105-128. HENJES-KUNST, F, ROLAND, N.W, DUNPHY, J. &

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FLETCHER, I. 2004. SHRIMP U-Pb dating of high-grade migmatites and related magmatites from northwestern Gates Land (East Antarctica): evidence for a single high-grade event of Ross-orogenic age. Terra Antartica, 11, 67-84. HOWELL, D.G. 1995. Principles of Terrane Analysis. New Applications for Global Tectonics (2nd edn). Chapman & Hall, London. IRELAND, T.R., WEAVER, S.D., BRADSHAW, J.D., ADAMS, C., GIBSON, G.M. 1999. Detrital age spectra from the Neoproterozoic-early Paleozoic Gondwana margin in Antarctica. Proceedings of the 8th International Symposium on Antarctic Earth Sciences, Abstracts, 150, Royal Society of New Zealand, Wellington. KLEINSCHMIDT, G. 1992. Structural observations in the Robertson Bay Terrane and their implications. Polarforschung, 60, 128-132. KLEINSCHMIDT, G. & SKINNER, D.N.B. 1981. Deformation styles in the basement rocks of North Victoria Land, Antarctica. Geologisches Jahrbuch, B41, 155-199. KLEINSCHMIDT, G. & TESSENSOHN, F. 1987. Early Palaeozoic westward-directed subduction at the Pacific margin of Antarctica. In: McKENZiE, G.D. (ed.) Gondwana Six: Structure, Tectonics and Geophysics. American Geophysical Union, Geophysical Monographs, 40, 89-105. KLEINSCHMIDT, G, ROLAND, N.W. & SCHUBERT, W. 1984. The metamorphic basement complex in the Mountaineer Range, North Victoria Land, Antarctica. Geologisches Jahrbuch, B60, 213-251. KREUZER, H., HOHNDORF, A., LENZ, H., MULLER, P. & VETTER, U. 1987. Radiometric ages of preMesozoic Rocks from Northern Victoria Land, Antarctica. In: McKENZiE, G.D. (ed.) Gondwana Six: Structure, Tectonics and Geophysics. American Geophysical Union, Geophysical Monographs, 40, 1-47. LAIRD, M.G. & BRADSHAW, ID. 1983. New data on the lower Palaeozoic Bowers Supergroup, northern Victoria Land, Antarctica. In: OLIVER, R.L., JAMES, PR. & JAGO, J.B. (eds) Antarctic Earth Science. Australian Academy of Science, Canberra, 123-126. LAUFER, A. & ROSSETTI, F. 2003. Late-Ross ductile deformation features in the Wilson terrane of Northern Victoria Land (Antarctica) and their implications for the western front of the Ross Orogen. Terra Antartica, 10, 179-196. MONGER, J.W.H., CLOWES, R.M., PRICE, R.A., SIMONY, S., RIDDIHOUGH, R.P. & WOODSWORTH, G.J. 1985. Juan de Fuca plate to Alberta plains. Geological Society of America, Centennial Continent/Ocean Transect 1, B-2. MUSUMECI, G. 1999. Magmatic belts in accretionary margins, a key for tectonic evolution; the tonalite belt of North Victoria Land (East Antarctica). Journal of the Geological Society, London, 156, 177-189. MUSUMECI, G, CAPPONI, G, MECCHERI, M. & OGGIANO, G. 1994. Deformed intrusives along the Wilson-Bowers terrane boundary (northern

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Victoria Land): structural features and tectonic evolution. Terra Antartica, 1, 54-58. OLIVER, R.L. & FANNING, C.M. 2002. Proterozoic geology east and southeast of Commonwealth Bay, George V Land, Antarctica, and its relationship to that of adjacent Gondwana terranes. In: GAMBLE, J.A., SKINNER, D.N.B. & HENRYS, S. (eds) Antarctica at the Close of a Millennium. Royal Society of New Zealand, Bulletin, 35, 51-58. PALMERI, R., PERTUSATI, PC., Rica, C.A. & TALARICO, F. 1994. Late Proterozoic(?)-Early Paleozoic evolution of the Active Pacific margin of Gondwana: evidence from the southern Wilson Terrane (northern Victoria Land, Antarctica). Terra Antartica, 1, 5-9. PALMERI, R., GHIRIBELLI, B., TALARICO, F. & Rica, C.A. 20030. Ultra-high-pressure metamorphism in felsic rocks: the garnet-phengite gneisses and quartzites from the Lanterman Range, Antarctica. European Journal of Mineralogy, 15, 513-525. PALMERI, R., TALARICO, F. & RICCI, C.A. 2003ft. Ultrahigh-pressure metamorphism at the palaeo-Pacific margin of Gondwana: the Lanterman Range in Antarctica. 9th International Symposium on Antarctic Earth Sciences, Potsdam, Programme and Abstracts. Terra Nostra, 2003/4, 249-250. PREISS, W.V 1987. The Adelaide geosyncline - Late Proterozoic stratigraphy, sedimentation, palaeontology and tectonics. Geological Survey of South Australia, Bulletin, 53. RICCI, C.A. (ed.) 1991. Proceedings of the Meeting 'Earth Science Investigations in Antarctica', Siena. Memorie della Societa Geologica Italiana, 46. RICCI, C.A. (ed.) 1997. The Antarctic Region: Geological Evolution and Processes. Terra Antartica Publications, Siena. RICCI, C.A. & TESSENSOHN, F. 2003. The LantermanMariner suture: Antarctic evidence for active margin tectonics in Palaeozoic Gondwana. In: TESSENSOHN, F. & RICCI, C.A. (eds) The Mariner Glacier Area, Antarctica - Aspects of a Suture Zone. Geologisches Jahrbuch, B85, 305-332. RICCI, C.A., TALARICO, F, PALMERI, R., Di VICENZO, G & PERTUSATI, PC 1996. Eclogite at the Antarctic Palaeo-Pacific margin of Gondwana (Lanterman Range, northern Victoria Land, Antarctica). Antarctic Science, 8, 277-280. RICCI, C.A., TALARICO, F. & PALMERI, R. 1997. Tectonothermal evolution of the Antarctic Palaeo-Pacific margin of Gondwana: a northern Victoria Land perspective. In: RICCI, C.A. (ed.) The Antarctic Region: Geological Evolution and Processes. Terra Antartica Publications Siena, 213-218. RICKER, J. 1964. Outline of the geology between the Mawson and Priestley glaciers, Victoria Land. In: ADIE, RJ. (ed.) Antarctic Geology. NorthHolland Publishing Company, Amsterdam, 265-275. ROCCHI, S., TONARINI, S., ARMIENTI, P., INNOCENTI, F.

& MANETTI, P. 1998. Geochemical and isotopic structure of the early Palaeozoic active margin of

Gondwana in northern Victoria Land, Antarctica. Tectonophysics, 284, 261-281. ROCCHI, S., CAPPONI, G, CRISPINI, L., Di VINCENZO, G, GHEZZO, C, MECCHERI, M. & PALMERI, R. 2003. Mafic rocks at the Wilson-Bowers terrane transition and within the Bowers Terrane: implications for a geodynamic model of the Ross orogen. Terra Antartica Reports, 9, 145-148. ROLAND, N., HENJES-KUNST, F, KLEINSCHMIDT, G. & TALARICO, F. 2001. Petrographical, geochemical, and radiometric investigations in northern Victoria Land, Oates Land and George V Land: towards a better understanding of plate boundary processes in Antarctica. Terra Antartica Reports, 5, 57-66. SANDIFORD, M. 1985. Structural evolution of the Lanterman metamorphic complex, northern Victoria Land, Antarctica. New Zealand Journal of Geology and Geophysics, 28, 443-458. SCHUBERT, W., OLESCH, M. & SCHMIDT, K. 1984. Paragneiss-orthogneiss relationships in the Kavrayskiy Hills, North Victoria Land, Antarctica. Geologisches Jahrbuch, B60, 187-211. SCHUSSLER, U, B ROCKER, M, HENJES-KUNST, F. & WILL, T. 1999. P-T-t evolution of the Wilson Terrane metamorphic basement at Oates Coast, Antarctica. Precambrian Research, 93, 235-258. SCHUSSLER, U, HENJES-KUNST, F, TALARICO, F. & FLOTTMANN, T. 2004. High-grade crystalline basement of the northwestern Wilson Terrane at Oates Coast: new petrological and geochronological data and implications for its tectonometamorphic evolution. Terra Antartica, 11, 15-24. SKINNER, D.N.B. 1991. Metamorphic basement contact relations in the southern Wilson Terrane, Terra Nova Bay, Antarctica - the Boomerang thrust. Memorie della Societa Geologica Italiana, 46, 163-178. SKINNER, D.N.B. & RICKER, J. 1968. The geology of the region between the Mawson and Priestly glaciers, north Victoria Land, Antarctica. Part I Basement meta-sedimentary and igneous rocks. New Zealand Journal of Geology and Geophysics, 11,1009-1040. SKINNER, D.N.B., JORDAN, H. & SCHMIDT-THOME, M. 1996. The Berg Group of Oates Land, East Antarctica. In: ROLAND, N.W. (ed.) German Antarctic North Victoria Land expedition GANOVEX. From Oates Land to Marie Byrd Land. Geologisches Jahrbuch, B89, 271-293. STUMP, E. 1987. Construction of the Pacific margin of Gondwana during the pannotios cycle. In: McKENZiE, G.D. (ed.) Gondwana Six: Structure, Tectonics and Geophysics. American Geophysical Union, Geophysical Monographs, 40, 77-87. STUMP, E. 1995. The Ross Orogen of the Transantarctic Mountains. Cambridge University Press, Cambridge. STURM, A. & CARRYER, SJ. 1970. Geology of the region between the Matusevich and Tucker Glaciers, northern Victoria Land, Antarctica. New Zealand Journal of Geology and Geophysics, 13, 408^35.

NORTH VICTORIA LAND TERRANES TALARICO, F. & CASTELLI, D. 1995. Relict granulites in the Ross orogen of northern Victoria Land (Antarctica), I. Field occurrence, petrography and metamorphic evolution. Precambrian Research, 75, 141-156. TALARICO, F. & KLEINSCHMIDT, G. 2003. Structural and metamorphic evolution of the Mertz Shear Zone (East Antarctic Craton, George V Land): implications for Australia/Antarctica correlations and East Antarctic craton/Ross orogen relationships. Terra Antartica, 10, 229-248. TALARICO F, BORSI L. & LOMBARDO B. 1995. Relict granulites in the Ross Orogen of northern Victoria Land (Antarctica), II. Geochemistry and palaeotectonic implications. Precambrian Research, 75, 157-174. TALARICO, F, GHIRIBELLI, B., SMITH-SIDDOWAY, C, PALMERI, R. & RICCI, C.A. 1998. The northern Victoria Land segment of the Antarctic paleoPacific margin of eastern Gondwana: new constraints from the Lanterman and Mountaineer Ranges. Terra Antartica, 5, 245-252. TESSENSOHN, F 1984. Geological and tectonic history of the Bowers structural zone in North Victoria Land, Antarctica. In: ROLAND, N.W. (ed.) German Antarctic North Victoria Land Expedition 1982/83, GANOVEX III, Vol. I. Geologisches Jahrbuch, B60, 371-396. TESSENSOHN, F, DUPHORN, K., JORDAN, H. ETAL. 1981. Geological comparison of basement units in North Victoria Land, Antarctica. Geologisches Jahrbuch, B41, 31-88. TESSENSOHN, F, KREUZER, H., HENJES-KUNST, F, KLEINSCHMIDT, G. & VETTER, U. 1996. Geologic Map of the Yule Bay Quadrangle, Victoria Land, Antarctica with Geological Description. Bundesanstalt fiir Geowissenschaften und Rohstoffe, Hannover.

291

TONARINI, S. & ROCCHI, S. 1994. Geochronology of Cambro-Ordovician intrusives in northern Victoria Land: a review. Terra Antartica, 1,46-50. WEAVER, S.D., BRADSHAW, J.D. & LAIRD, M.G. 198 Geochemistry of Cambrian volcanics of the Bowers Supergroup and implications for the Early Palaeozoic tectonic evolution of northern Victoria Land, Antarctica. Earth and Planetary Sciences Letters, 68, 128-140. WEAVER, S., BASSETT, K., BRADSHAW, J. & IRELAND, T. 2003. Accretion of the Cambrian Bowers Terrane arc to the Gondwana margin, northern Victoria Land: evidence from conglomerate provenance. 9th International Symposium on Antarctic Earth Sciences, Potsdam, Programme and Abstracts. Terra Nostra, 2003/4, 335-336. WODZICKY, A. & ROBERT, R. JR. 1986. Geology of the Bowers Supergroup, central Bowers Mountains, northern Victoria Land. In: STUMP, E. (ed.) Geological Investigations in Northern Victoria Land. American Geophysical Union, Antarctic Research Series, 46, 39-68. WOLFART, R. 1994. Middle Cambrian faunas (brachiopoda, mollusca, trilobita) from exotic limestone blocks, Reilly Ridge, North Victoria Land, Antarctica; and their biostratigraphic and paleobiogeographic significance. Geologisches Jahrbuch, B84, 1-161. WRIGHT, TO. & DALLMEYER, R.D. 1991. The age of cleavage development in the Ross orogen, northern Victoria Land, Antarctica: evidence from 40Ar/39Ar whole-rock slate ages. Journal of Structural Geology, 13, 677-690. WRIGHT,TO., Ross, R.J. & REPETSKI, IE. 1984. New discovered youngest Cambrian or oldest Ordovician fossils from the Robertson Bay Terrane (formerly Precambrian), North Victoria Land, Antarctica. Geology, 12, 301-305.

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Investigating the deep structure of terranes and terrane boundaries: insights from earthquake seismic data ANYA M. READING Research School of Earth Sciences, Australian National University, Canberra, ACT, 0200, Australia (e-mail: [email protected]) Abstract: The crust and upper mantle structure beneath a single, three-component broadband seismic station may be determined using energy from distant earthquakes. Features such as the depth of the seismic Moho, the abruptness of the change in seismic velocity across the Moho, the velocity profile through the crust and the seismic velocity of the upper mantle may be found using receiver function techniques. Waveforms are analysed which contain energy arising from the interaction of incoming signals with the Earth structure beneath the receiving station. Seismic characteristics of the deep crust are often continuous or slowly varying across a single terrane and, moreover, show sudden contrasts across terrane and other major tectonic boundaries. Such techniques are, therefore, appropriate tools in the exploration of deep structure and, hence, in the understanding of large-scale continental assembly. They complement geological and geochronological investigations of surface rocks by providing the third spatial dimension. In regions of no surface basement exposure, seismic and other geophysical methods provide the only means of mapping the lateral extent of a given terrane and/or tracing province boundaries. Receiver function methods are presented here in the context of terrane tectonics with illustrative examples from former Gondwanan provinces and locations on the Pacific margin.

Extensive regions of continental lithosphere were formed through the assembly of allochthonous geological provinces or terranes (e.g. Coney et al. 1980). Defined as 'fault-bounded blocks of the Earth's crust characterized by a geological history distinct from that of adjacent blocks' (Friend et al. 1988), terranes and related concepts are central to the contemporary view of the evolution of the continents. Advances in understanding both detailed regional geology and, more generally, interactions at plate boundaries, depend on terrane models. Terranes, however, are not merely surface veneers, nor even (in spite of the above definition) confined to the Earth's crust, but may also form part of the upper mantle. The allochthonous nature of large terranes could even extend to full lithospheric depth. This contribution explores the ways in which seismic structure, derived from records of distant earthquakes, complements geological and geochronological investigations with examples from former Gondwanan provinces, including those on the Pacific margin of Australia. Insights from deep structure strengthen observations made on the surface and, importantly, enable extrapolations to be made into regions with no surface exposure.

Terranes at depth and the continental lithosphere The relationship between the crustal terranes observed at the surface and the sub-crustal lithospheric mantle (SCLM) was illustrated for a range of Archaean to Proterozoic terranes by O'Reilly et al. (2001). They used xenoliths and xenocrysts to construct empirical palaeogeotherms and, hence, constrain the geochemical structure at depth and over time (O'Reilly & Griffin 1996). In general, mantle domains that stabilized during the Archaean and Proterozoic form highly buoyant roots of SCLM beneath the associated crust to approximately 200 km deep, whereas Palaeozoic SCLM is denser, prone to delamination, leaving much thinner lithosphere. Keppie et al. (2003) examined shallower structures and discussed the diverse methods by which terranes may be transferred, adding to the larger, existing continental craton/landmass. With few exceptions, larger terranes take their lower crust and upper mantle foundations with them when they amalgamate to form new continental lithosphere. Less extensive regions of continental crust, e.g. the Antarctic Peninsula, also consist of allochthonous terranes (Vaughan & Storey 2000) and, again, the deep structure places important constraints on the processes by which the continental lithosphere has assembled and

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 293-303. 0305-8719/$15.00 © The Geological Society of London 2005.

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evolved. From many perspectives, it is indeed the case that 'understanding the architecture of the lower crust and mantle of continental regions is critical to understanding the largescale processes responsible for the development of the Earth's continents' (O'Reilly et al. 2001). Seismic methods have a significant role to play in advancing the understanding of crust and upper mantle architecture at a variety of scales. In some locations, marine, ship-to-shore or land-based active seismic experiments may be appropriate (e.g. Stern et al. 2002) but these are major, necessarily costly, undertakings. From the diverse range of seismic methods, receiver function analysis (e.g. Vinnik et al. 1996) of earthquake records has the advantage of obtaining crustal structure using naturally occurring sources. The resolution of structure obtained is lower than active source reflection/refraction methods, but is more than adequate for comparing lithospheric characteristics associated with contrasting terranes. Images obtained from seismic tomography (e.g. Kennett 2002) are invaluable in establishing the deep structural framework of continental lithosphere. The scale of structure that may be examined using teleseismic earthquake data to produce tomographic images extends from regional studies using (e.g.) short-period body-waves through to (e.g.) long-period surface-wave investigations of the whole Earth. In general, seismic tomography tends to smooth lateral contrasts in seismic structure and is not well suited to the delineation of terranes or their boundaries. Seismic tomography generally requires multiple recording stations whereas receiver function methods can produce new determinations of structure using data from a single location.

Terrane boundaries and deep structural contrasts Seismic characteristics of the deep crust are often continuous or slowly varying across a single terrane and, moreover, show sudden contrasts across terrane and other major tectonic boundaries. Features such as the depth of the seismic Moho, the abruptness of the change in seismic velocity across the Moho, the velocity profile through the crust and the seismic velocity of the upper mantle may be found routinely using receiver function techniques. The coarse-scale mapping of deep structural characteristics enables the lateral extent of a given terrane and the location of terrane boundaries to be determined. Mantle velocities and the nature of the Moho and crustal

discontinuities also show evidence for structural processes, such as underplating and rifting. It should be noted that discussion of features such as the 'Moho' is often techniquedependent. The geochemical Moho, for example, is a compositional boundary, whereas the seismic Moho is a discontinuity in physical characteristics that may or may not coincide with the chemical boundary. While such disparities can often give rise to illuminating observations regarding deep structure, care should be taken when comparing quantities such as the depth of the Moho determined in different ways.

Lithospheric evolution and changes in deep structure When using seismic methods that image contemporary structure to constrain the plate tectonic processes responsible for the observed architecture, it is important to consider both continental assembly and lithospheric evolution taking place after continental assembly. Condie & Chomiak (1996) concluded that crust formed by accretion during the Early Proterozoic evolved directly into mature continental crust, whereas Mesozoic accretion (at least in North America) involves the transformation of mafic oceanic terranes into continental crust. The Mesozoic crust was also more susceptible to fragmentation and displacement along transcurrent faults. The picture in the sub-continental lithospheric mantle is the same, with ancient crust being protected by its buoyant stable roots (Poudjom Djomani et al. 2001; also, see the Applications and Discussion section). Crustal erosion and uplift can also play a part in changing some aspects of deep structure, e.g. Moho depth. Significant surface erosion, causing the isostatic uplift of the crust, will reduce the depth of the Moho. Typical uplift values for post-Gondwanan margins (e.g. Lisker 2002) are of the order of 2 km to 4 km. Receiver function methods constrain the Moho depth to within 2 km, suggesting that it is important to consider crustal erosion and uplift but there are few cases where the change in Moho depth, due to erosion and uplift, would cause real problems in using contemporary seismic structure in the understanding of past processes. Constraints imposed on the use of seismic methods by terrane fragmentation and influences on the crust, such as extensive igneous emplacement, are discussed in detail in the next section. In general, mapping the deep structure of Archaean and Proterozoic lithosphere using receiver function methods is a straightforward

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Seismic receiver functions as analytical tools

Fig. 1. Sketch showing how a ID seismic structure may be derived from a teleseismic earthquake: (a) simplified Earth structure and incoming earthquake energy; (b) simplified radial receiver function calculated from three-component data; (c) the corresponding derived seismic velocity profile.

undertaking. Mapping Mesozoic terranes is also possible with sufficient numbers of stations but is limited to terrane fragments greater than 100 km wide.

Seismic receiver function methods are valuable tools in the determination of first-order crustal and upper mantle structure. The upcoming Pwave energy from a distant earthquake (between 30° and 80° epicentral distance) refracts as it passes through the Earth and reaches the layered structure beneath a seismic recording station at a steep angle (Fig. 1). A proportion of the energy changes from P-wave to S-wave energy at the Moho and continues at a slower speed, reaching the seismic station about 5 s later than the large initial pulse from the direct P-wave. Some of the P-wave energy passing through the Moho unchanged is converted to S-wave energy at discontinuities within the crust. This pulse (a P to S conversion from a shallow discontinuity) reaches the recording station after the initial P-pulse but before the P to S converted pulse from the Moho. The pulses due to the structure beneath the receiving station are isolated from the other energy arriving in the P-wave coda by deconvolution (Ammon 1991; Stein & Wysession 2003). This process requires three-component seismic data rotated into vertical, horizontal/radial (i.e. the direction from source to receiver) and horizontal/transverse directions. Figure Ib shows an idealized radial receiver waveform for a layered crust with two discontinuities in seismic velocity, in the upper crust and at the Moho. Figure Ic shows an idealized seismic velocity profile associated with the rays and receiver function in Figure la and Ib. The change in seismic velocity across the Moho is greater than across the shallow discontinuity (resulting in a higher amplitude pulse). In practice, a best-fit seismic velocity profile often shows a high-velocity gradient zone (HVGZ) at the Moho rather than a very sharp step and the Moho is taken to be the base of the HVGZ. The observed receiver functions in this work are calculated using the method of Shibutani et al. (1996) and are then compared with synthetic receiver functions taking account of the highly non-linear relationship between receiver functions and seismic structure. An appropriate inverse approach, the Neighbourhood Algorithm (NA, Sambridge 1999), is used to investigate many possible velocity models to find the best-fitting solution. The NA searches the parameter space using an efficient, adaptive technique and is well suited to non-linear applications. The ray paths illustrated in Figure la are somewhat misleading. Seismic energy passing from source to a receiver at teleseismic distance

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Fig. 2. Station locations (large open triangles) and summary tectonic structure for the examples in this contribution, (a) SW, Southwest terrane; SC, Southern Cross terrane; EG, Eastern Goldfields terrane; Ida, Ida Fault; WT02-10 and WV01-07 are stations of the WACRATON seismic deployments, terrane boundaries after Myers & Hocking (1998), redrawn from Reading et al. (2003). (b) The Lambert Glacier region, East Antarctica. MAW, Mawson Antarctic Base and seismic station; BVLK, Beaver Lake summer camp and seismic station; DAY, Davis Antarctic Base; DAY, REIN, NMES, KOMS, CRES, BVLK and JACK are stations of the SSCUA seismic deployment 1 (2002/03). Terrane structures and boundaries (none shown) are the subject of continued debate. The Rayner Province (e.g. Fitzsimons 2003) extends southward from Mawson. (c) Tasmania and southern Victoria, Australia and summary tectonic boundaries. Ben, Bendigo zone; Sel, Selwyn Block; Tab, Tabberabbera zone (e.g. Cayley et al. 2002); R Cape, Rocky Cape element; Dun, Dundas element; NE Tas, North East Tasmania element (e.g. Munker & Crawford 2000; Meffre et al. 2000). Stations TA01-13 and TB01-04 formed the broadband component of the TIGGER seismic deployment.

has a maximum frequency of around 1 Hz. It travels in a banana-shaped ray-tube, which intersects the Moho discontinuity in an elliptical region approximately one wavelength across (8 km for a I Hz wave, the width of the first Fresnel zone). Imbricate slivers of crust, thinner than the crustal depth, would violate the assumption approximating the crust to layers with ID symmetry beneath an individual station. In summary, as a rule of thumb, terrane structure may be investigated using receiver functions providing the width of the terrane is greater than the [Moho depth] + [twice the Fresnel zone], assuming that fairly steep-angled faults form the terrane boundaries. For a 60 km thick crust, the minimum terrane width for seismic study would be about 80 km. The Fresnel zone limits, in a fundamental way, the extent to which the receiver function method can resolve details in Moho topography. This situation is, however, of great practical advantage because it means that structures such as igneous intrusions, of the order of < 4 km across at the Moho, are transparent to seismic receiver functions (although multiple intrusions may have an effect on the best-fit seismic velocity). Dipping structures affect the determination of seismic velocity if they are steeper than approximately 20° (Cassidy 1992). Where terrane boundaries are shallow-dipping, boundary faults may lead to complex observed receiver

functions and care must be taken in interpretation. Illustrative examples In this section, examples of crust and upper mantle velocity structures determined using the receiver function method are presented. Figure 2 shows the locations of the seismic stations. Cases illustrated are (1) that of consistent structure within a given terrane, (2) contrasting structure across a terrane or other major tectonic boundary and (3) complex structures arising at a terrane boundary A single terrane Consistency of structure within a single terrane, in this case, the Southwest terrane within the Archaean Yilgarn craton, Western Australia, is shown in Figure 3a, b for stations WT05 and WT06, part of the broader WACRATON experiment (Reading etal 2003). Although not located in a Pacific-margin environment, this Gondwanan example nevertheless provides a clear end-member case of consistent structure within a stable craton (Betts et al. 2002). The geological exposure, and geochronological record is sufficient to identify the major terrane boundaries of the Yilgarn craton positively (Myers & Hocking 1998), while the location in the

DEEP SEISMIC STRUCTURE AND TERRANES continental interior ensures a relatively good signal to noise ratio for the observed receiver functions. Both WT05 and WT06 show the following characteristics: a high upper mantle velocity, a Moho at similar depths (41 km ± 2 km and 40 km, respectively), a broad, positive velocity gradient above the Moho (10 km ± 2 km and 8 km thick, respectively), a positive velocity gradient in the shallow crust (between 2-12 km ±2 km and 2-14 km, respectively) and little change in the P-velocity to ^-velocity ratio throughout the profile. Within the confidence limits of the inversion for structure, the velocity profiles determined from the receiver functions are very similar. A comparable result is found in southern Africa (Nguuri et al. 2001), where the central terrane of the Kaapvaal craton has a consistent structure where it is undisturbed by later tectonic events. Figures 3c, d shows first results from a deployment in Antarctica. In this case, the geological data are much more sparse, although those that do exist are consistent with the two stations being located in the same terrane. The data, especially from Mawson station, MAW, are noisy and affected by reverberations. Some shortcomings in the fit of the waveform from Beaver Lake station, BVLK, are apparent, but will be improved as more data from this station become available. The features identified are, nonetheless, sufficiently robust for comparison to be made between the stations. Both MAW and BVLK show a low-velocity upper mantle, a Moho at 44 km (±2 km), a negative velocity gradient above the Moho, a mid-crustal discontinuity (at 26 km ± 2 km and 28 km, respectively), a shallow discontinuity at 10 km and a big drop in P-velocity:S-velocity ratio between the shallow and mid-crust discontinuities. They certainly fall within the same terrane from the perspective of deep crust and upper mantle seismic structure. Assuming that adding more data would improve the fit of the BVLK structure, the Moho would become more step-like, more similar to the step-like Moho observed beneath MAW. In some instances, it is possible to correlate crustal structures derived from receiver functions across passive continental margins. Reading (2004) describes crustal velocity structures from Antarctica and southern Western Australia that are very similar. This could indicate that the region close to Casey Station, East Antarctica and the Albany-Fraser Orogen, Australia are former neighbours in a reconstructed Gondwana continent, owing their deep structure to an origin within a single or related terrane.

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Across a terrane boundary Contrast in structure across a terrane boundary is illustrated using newly determined structure beneath seismic stations (part of the TIGGER experiment) in southern Victoria and Tasmania, Australia (Fig. 2c). Figure 4a shows determinations of structure beneath station TB03, located in the Selwyn Block in southern central Victoria (Cayley et al. 2002). This profile is characterized by a 10 km (±2 km) thick high-velocity upper crust, divided from a similarly high-velocity mid-crust by a low-velocity layer. The waveform is not well fitted, and it is impossible to draw anything other than these basic conclusions. A peak in the observed waveform at 8.5 s might indicate a deep discontinuity (>50 km) or it might be due to multiple paths in the upper crust. In contrast, the modelled profile beneath TB01, located in the Tabberabbera zone in eastern Victoria (Cayley et al. 2002) shows a medium velocity upper mantle, a sharp Moho at 39 km (±2 km) deep, with a discontinuity at 29 km, and a sharp discontinuity at 15 km deep. The two stations, located in different terranes show markedly different deep structures. Stations TA02 and TA04 are located either side of the Arthur Lineament, which divides the Rocky Cape and Dundas elements in northwest Tasmania (Holm & Berry 2002). The structure beneath TA02 is characterized by a shallow discontinuity at 6 km. Between 6 km and 21 km deep there are high velocities and a negative velocity gradient. Deeper than 21 km, the velocity increases again with depth, to a fairly shallow Moho at 31 km. There is no strong discontinuity across the Moho beneath the Rocky Cape element. Station TA04 is characterized by a deeper Moho at 35 km (±2 km) with a more pronounced HVGZ between 32 km and 35 km. A positive velocity gradient zone exists between 2 km and 10 km deep, but there are no marked high or low velocity zones. Again, there is a strong contrast between seismic velocity profiles on either side of the major tectonic boundary.

At a terrane boundary Structural features associated with terrane boundaries, where it becomes difficult to use receiver function methods in the analysis of deep structure, are given in Figure 5. Both WT08 and WV02 are located in the Southern Cross terrane in the centre of the Yilgarn craton, Western Australia, well away from any terrane boundary. They both show Moho discontinuities with a strong velocity contrast,

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Fig. 3. Seismic structure beneath a single terrane: (a) WT05 and (b) WT06 are located beneath the Southwest terrane, Western Australia (see also Fig. 2a); (c) MAW and (d) BVLK are located in the Rayner Province, East Antarctica (see also Fig. 2b). Upper plots show the observed (solid) and synthetic (dashed) receiver functions. Observed receiver functions are stacked records derived from between four and ten earthquakes. Lower plots show the best-fit seismic velocity profile (S-velocity against depth) corresponding to the synthetic receiver function (heavy white line) plotted above the models (grey region) searched by the Neighbourhood Algorithm (Sambridge 1999). To the left of each plot (fine dark line) is the P-velocity:S-velocity ratio which is part of the inversion parameterization and is, therefore, shown for completeness. Features of the velocity profiles that are discussed in the text are indicated by arrows for this figure only. HVG2, high-velocity gradient zone.

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Fig. 4. Seismic structure beneath contrasting terranes: (a) TB03, over the Selwyn Block and (b) TB01, in the Tabberabbera zone, southern Victoria (Fig. 2c); (c) TAOl in the Rocky Cape element and (d) TA04 in the Dundas element, Tasmania (Fig. 2c). (See caption to Fig. 3 for explanation of plots.)

38 km (±2 km) deep beneath WT08 and 40 km deep beneath WV02. In contrast, WT09 and WV03 are located above the deep expression of the eastward-dipping Ida Fault (Goleby et aL

2000). They both show complicated upper crustal discontinuities and low-velocity zones that mask the structure beneath, which is unconstrained. To examine the structure

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Fig. 5. Seismic structure at a terrane boundary: (a) WT08, Southern Cross terrane and (b) WT09 over the Ida Fault; (c) WV02, Southern Cross terrane and (d) WV03, over the Ida Fault (Fig. 2a). (See caption to Fig. 3 for explanation of plots.)

beneath WT09 and WV03 further, it would be necessary to use other methods, for example active source seismic reflection and refraction,

to determine the deeper structure. Nevertheless, the difference in the characteristic structure beneath WT08 (i.e. mid-terrane) and

DEEP SEISMIC STRUCTURE AND TERRANES WT09 (over the terrane boundary) is striking. In this location, there is additional higher-resolution geophysical information to inform the interpretations, but, in principle, the terrane boundary could be mapped even if there was no surface exposure. Coupled with models of appropriate or analogous tectonic setting (e.g. Miiller et al. 1996), receiver functions can provide constraints on the basic form and position of such boundary structures.

Applications and discussion Investigating consistent or contrasting seismic structure with reference to known terranes and terrane boundaries, as illustrated in this contribution, now allows use of seismic receiver functions as a tool that complements geological and geochronological information for a given region. There are two main classes of application: (1) providing the third, spatial dimension in structural mapping of large crustal blocks; and (2) mapping the extent of terranes and the location of terrane boundaries beneath regions of no surface exposure. In mapping large crustal blocks, it is often the case that patchy surface exposure is extrapolated using lineations and block structure from aerogeophysical surveys. The success of this approach is dependent on the depth of the structures giving rise to the magnetic anomalies and/or the height at which the survey is flown. Over regions of economic importance, surveys are flown low, at high resolution and often the deep structure is not recovered. Regional-scale surveys (e.g. Ferraccioli & Bozzo 1999) flown in Antarctica can be much more successful, but it is often the case that geological interpretations are supplemented only by shallow geophysics. Adding portable seismic stations to a full-season geological expedition, given that the energy sources (distant earthquakes) come free of charge, is a relatively straightforward undertaking. A structural profile in the third spatial dimension, which is resolved using seismic data, is a fundamental constraint on the morphology and, hence, the structural history of the tectonic Earth in that location. This complements the insights to be gained from surface structural mapping and geochronology. Where there is no surface exposure, the basement is covered by younger cover rocks/drift or ice sheets, seismic receiver functions may be used to extrapolate known terrane blocks and block boundaries. Nowhere is this more important than the Pacific margin of Antarctica, with rock exposed only around the periphery of East Antarctica and key regions of

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Central West Antarctica covered with ice sheets. Further north, along the Pacific margin of Australia, there are vast regions covered with sand and recent deposits. Initial receiver function work (Clitheroe et al. 2000) has provided a contour map of the Moho at a coarse scale but further, denser deployments are required to delineate terrane boundaries further. The use of seismic receiver functions in the investigation of the deep structure and extent of terranes is limited by the resolution imposed by the frequency of earthquake waves to tectonic units greater than 100 km across. An exception would be regions of especially shallow crust, where it would be possible to examine blocks down to about 40 km across using receiver function methods. From a practical point of view, it is important to use data with a good signal to noise ratio. Coastal stations would therefore require a longer recording time to build a sufficient number of suitable records than their (low-noise) counterparts deep in the continental interior. In both cases, a geographical position within 30-40° of an active seismic zone also reduces the station deployment time required to obtain a suitable dataset. Although events may be used out to 80° or 90°, smaller magnitude and, hence, more frequent, events are recorded suitably at closer distances. Ideally, earthquakes should be recorded at a range of azimuths from the receiving station to provide insights into structural heterogeneity and seismic anisotropy beneath the station, although a detailed discussion of these important technical issues is beyond the scope of this contribution. As a working procedure, if the majority of the converted seismic energy is not on the radial component of the rotated threecomponent seismic record, and significant converted energy is seen on the transverse receiver function, then the structure beneath the station is insufficiently close to ID to be investigated using this receiver function method (Shibutani et al. 1996). Finally, it is important to consider the changes in the crust and SCLM, which may have taken place after terrane accretion, between the major phases of continental growth and the present day. Any process that affects the physical properties of crust and SCLM should be accounted for when using the contemporary structures 'visible' using seismic techniques to infer the structures and modes of continental accumulation. As mentioned previously, fairly large intrusions may not affect receiver function determinations of seismic structure. Also, faulting within a given terrane will not give rise

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to large contrasts in seismic structure and, hence, the method will remain viable. An area of concern might be the instability of Proterozoic SCLM and the consequent exposure of the upper lithosphere to the influence of the asthenosphere. If the lithosphere suffered a partial delamination and a large terrane was subject to heterogeneous influence of the asthenosphere, then the deep structure of a buried terrane might resemble two contrasting provinces, causing a terrane boundary to be predicted falsely. Provided the entire terrane is affected similarly, the methods outlined in the paper remain valid. Summary The consistency of deep structure within a single terrane and the contrast of deep structure across terrane boundaries have been illustrated. Seismic receiver function methods provide a means of (1) resolving crust and upper mantle structure at depth, complementing surface observations and (2) of extrapolating surface observations over regions where there is no surface cover. Technical and research staff members of the Seismology Group, RSES, ANU played an essential role in experimental support and field data acquisition. Mawson seismic station is operated by Geoscience Australia. Australian Antarctic Division supported the field deployment at Beaver Lake. The manuscript was improved by comments from referees Ed King and Richard England.

References AMMON, CJ. 1991. The isolation of receiver effects from teleseismic P waveforms. Bulletin of the Seismological Society of America, 81, 2504-2510. BETTS, P.G., GILES, D., LISTER, G.S. & FRICK, L.R. 2002. Evolution of the Australian lithosphere. Australian Journal of Earth Sciences, 49, 661-695. CASSIDY, J.F. 1992. Numerical experiments in broadband receiver function analysis. Bulletin of the Seismological Society of America, 82,1453-1474. CAYLEY, R.A.,TAYLOR, D.H., VANDENBERG, A.H.M. & MOORE, D.H. 2002. Proterozoic-Early Palaeozoic rocks and the Tyennan orogeny in central Victoria: the Selwyn block and its tectonic implications. Australian Journal of Earth Sciences, 49, 225-254. CLITHEROE, G., GUDMUNDSSON, O. & KENNETT, B.L.N. 2000. The crustal thickness of Australia. Journal of Geophysical Research, 105, 13 697-13 713. CONDIE, K. & CHOMIAK, B. 1996. Continental accretion: contrasting Mesozoic and Early Proterozoic tectonic regimes in North America. Tectonophysics, 265,101-126.

CONEY, P.J., JONES, D.L. & MONGER, J.W.H. 1980. Cordilleran suspect terranes. Nature, 288, 329-333. FERRACCIOLI, F. & Bozzo, E. 1999. Inherited crustal features and tectonic blocks of the Transantarctic Mountains: an aeromagnetic perspective (Victoria Land, Antarctica). Journal of Geophysical Research, 104, 25 297-25 319. FITZSIMONS, I.C.W. 2003. Proterozoic basement provinces of southern and southwestern Australia, and their correlation with Antarctica. In: YOSHIDA, M., WINDLEY, B.F. & DASGUPTA, S. (eds) Proterozoic East Gondwana: Supercontinent Assembly and Breakup. Geological Society, London, Special Publications, 206, 93-130. FRIEND, C.R.L., NUTMAN, A.P & MCGREGOR, V.R. 1988. Late Archaean terrane accretion in the Godthab region, southern West Greenland. Nature, 335, 535-538. GOLEBY, B.R., BELL, B., KORSCH, RJ. ET AL. 2000. Crustal structure and fluid flow in the Eastern Goldfields, Western Australia. Results from the AGCRC's Yilgarn deep seismic reflection survey and fluid flow modelling projects. Australian Geological Survey Organisation Record, 2000/34. HOLM, O.K. & BERRY, R.F 2002. Structural history of the Arthur Lineament, northwest Tasmania: an analysis of critical outcrops. Australian Journal of Earth Sciences, 49,167-185. KENNETT, B.L.N. 2002. The Seismic Wavefield Volume II: Interpretation of Seismograms on Regional and Global Scales. Cambridge University Press, Cambridge, 461^87. KEPPIE, J.D., NANCE, R.D., MURPHY, J.B. & DOSTAL, J. 2003. Tethyan, Mediterranean, and Pacific analogues for the Neoproterozoic-Paleozoic birth and development of peri-Gondwanan terranes and their transfer to Laurentia and Laurussia. Tectonophysics, 365, 195-219. LISKER, F. 2002. Review of fission trace studies in northern Victoria Land, Antarctica - passive margin evolution versus uplift of the Transantarctic Mountains. Tectonophysics, 349, 57-73. MEFFRE, S., BERRY, R.F. & HALL, M. 2000. Cambrian metamorphic complexes in Tasmania: tectonic implications. Australian Journal of Earth Sciences, 47, 971-985. MULLER, W.U., DAIGNEAULT, R., MORTENSEN, IK. & CHOWN, E.H. 1996. Archaean terrane docking: upper crust collision tectonics, Abitibi greenston belt, Quebec, Canada. Tectonophysics, 265, 127-150. MUNKER, C. & CRAWFORD, AJ. 2000. Cambrian arc evolution along the SE Gondwana active margin: a synthesis from Tasmania-New ZealandAustralia-Antarctica correlations. Tectonics, 19, 415-432. MYERS, J.S. & HOCKING, R.M. 1998. Geological Map of Western Australia, 1:2 500 000 (13th edition). Geological Survey of Western Australia, Perth, Australia. NGUURI, T.K., GORE, I, JAMES, D.E. ET AL. 2001. Crustal structure beneath southern Africa and its implications for the formation and evolution of

DEEP SEISMIC STRUCTURE AND TERRANES the Kaapvaal and Zimbabwe cratons. Geophysical Research Letters, 28, 2501-2504. O'REILLY, S.Y. & GRIFFIN, W.L. 1996. 4-D Lithosphere mapping: methodology and examples. Tectonophysics, 262, 3-18. O'REILLY, S.Y., GRIFFIN, W.L., POUDJOM DJOMANI, Y.H. & MORGAN, P. 2001. Are lithospheres forever? Tracking changes in subcontinental lithospheric mantle through time. GSA Today, 11 (April), 4-10. POUDJOM DJOMANI, Y.H., O'REILLY, S.Y, GRIFFIN, W.L. & MORGAN, P. 2001. The density structure of subcontinental lithosphere through time. Earth and Planetary Science Letters, 184,605-621. READING, A.M. 2004. The seismic structure of Wilkes Land/Terre Adelie, East Antarctica and comparison with Australia: first steps in reconstructing the deep lithosphere of Gondwana. Gondwana Research, 7, 21-30. READING, A.M., KENNETT, B.L.N. & DENTITH, M.C. 2003. Seismic structure of the Yilgarn Craton, Western Australia. Australian Journal of Earth Sciences, 50, 427^38. SAMBRIDGE, M.S. 1999. Geophysical inversion with a neighbourhood algorithm. I. Searching a para-

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meter space. Geophysical Journal International, 138, 479-494. SHIBUTANI, T, SAMBRIDGE, M. & KENNETT, B. 1996. Genetic algorithm inversion for receiver functions with application to crust and uppermost mantle structure beneath Eastern Antarctica. Geophysical Research Letters, 23, 1829-1832. STEIN, S. & WYSESSION, M. 2003. An Introduction to Earthquakes, Seismology and Earth Structure. Blackwell Publishing, Oxford. STERN, T, OKAYA, D. & SCHERWATH, M. 2002. Structure and strength of a continental transform from onshore-offshore seismic profiling of South Island, New Zealand. Earth Planets and Space, 54,1011-1019. VAUGHAN, A.P.M. & STOREY, B.C. 2000. The eastern Palmer Land shear zone: a new terrane accretion model for the Mesozoic development of the Antarctic Peninsula. Journal of the Geological Society, London, 157, 1243-1256. VINNIK, L.P., GREEN, R.W.E., NICOLAYSEN, L.O., KOSAREV, G.L. & PETERSEN, N.V. 1996. Deep seismic structure of the Kaapvaal craton. Tectonophysics, 262, 67-75.

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The accretionary history of southern South America from the latest Proterozoic to the Late Palaeozoic: some palaeomagnetic constraints AUGUSTO E. RAPALINI INGEODAV, Departamento de Cs. Geologicas, Fac. Cs. Exactas y Naturales, Universidad de Buenos Aires, Pabellon 2, Ciudad Universitaria, C1428EHA, Buenos Aires, Argentina (e-mail: [email protected]) Abstract: It is now accepted that southern South America was formed from several terranes of diverse origin and evolution. However, a detailed history of the accretionary processes has not been unravelled yet. Palaeomagnetism can play an important role in such an endeavour. Palaeomagnetic constraints on the tectonic evolution of this region in the Proterozoic and Palaeozoic are reviewed and discussed. Data from the Rio de la Plata craton suggest that this block was already attached to most major Gondwana blocks by the end of the Proterozoic and may have formed a single continental mass with Congo-Sao Francisco, West Nile and Arabia throughout most of the Vendian. A large ocean separating these cratons from Amazonia and West Africa, prior to Gondwana assembly, is supported by available palaeomagnetic data. To the west of the Rio de la Plata craton is the Pampia terrane. Despite lack of palaeomagnetic data, geological evidence supports a model of Early Cambrian collision between these blocks. An Early Ordovician magmatic arc, the Famatina-Eastern Puna belt, which had developed on the western margin of the already accreted Pampia terrane, shows a systematic pattern of large clockwise rotation that has been interpreted as representative of the whole terrane. The favoured tectonic model portrays a continental magmatic arc with a back-arc basin to the east that was closed when the terrane rotated. There is little doubt of a Laurentian origin for the Cuyania (Precordillera) terrane, given the amount and diversity of evidence, including palaeomagnetism. The tectonic mechanism for accretion and its timing are still controversial. New palaeomagnetic data from Late Ordovician rocks of Cuyania support the 'Laurentian plateau' hypothesis, which suggests that Cuyania was still linked to Laurentia well into the Ordovician. Nevertheless, these new data do not rule out the more generally favoured 'microcontinent model'. To the west of Cuyania is the Chilenia terrane, separated by a belt of ophiolites of Late Ordovician age. Very little is known about this terrane, although some U-Pb ages and Nd model ages point to a Laurentian origin for its basement. Lack of palaeomagnetic data precludes determining its kinematic evolution. The ArequipaAntofalla block may actually be a composite terrane. Palaeomagnetic data obtained so far come exclusively from the southern Antofalla block. Recently acquired data in the western Puna of Argentina confirm the originally proposed distribution of Early Palaeozoic palaeomagnetic poles, which, despite several uncertainties, delineate a pattern of significant counterclockwise rotations with a possible anomaly in palaeolatitude for the late Cambrian. The data suggest a major tectonic discontinuity between the Eastern and Western Puna of Argentina in the Early Palaeozoic. Four palaeomagnetic poles of Devonian to Permian age from the North Patagonian Massif are consistent in position and age with the Gondwana apparent polar wander path, suggesting that both continental masses have not experienced major relative displacement since the Devonian. The data do not, however, rule out a restricted separation of Patagonia orthogonal to its northern boundary in the Early or Middle Palaeozoic and subsequent collision in the Late Palaeozoic.

Introduction From a global palaeogeographical perspective, the Late Proterozoic can be bracketed between the dispersal of the hypothetical supercontinent of Rodinia, at approximately 750 Ma (Hoffman 1991; Powell et al 1993; Meert & Torsvik 2003; Cordani et al 2003) and the assembly of the Gondwana supercontinent at around 550 Ma

(e.g. Meert 2001). Palaeogeographical reconstructions for the Late Proterozoic rely heavily on palaeomagnetism, given the lack of preserved ocean floor magnetic anomalies, no recognized hotspot tracks and the paucity of the fossil record. Palaeomagnetic records for the Proterozoic face severe problems that may affect their reliability and potential. These include a higher likelihood of remagnetizations,

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 305-328. 0305-8719/$15.00 © The Geological Society of London 2005.

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inaccurate rock ages and more reduced and incomplete exposures. This has meant that despite continuous efforts and slow improvements, the Proterozoic palaeomagnetic database (e.g. Meert & Torsvik 2003) is still scarce, with considerable gaps both in time and space for every major continental block. In spite of all these problems, palaeomagnetism is still a major resource for improving knowledge of the global and regional palaeogeographical evolution in the Late Proterozoic and Early Palaeozoic. Recent palaeomagnetic results have cast doubts on the validity of the most popular proposals of Rodinia configuration (e.g. Torsvik et al 2001; Cordani et al 2003), leaving this as a virtually open field that awaits a much larger palaeomagnetic database from different tectonic blocks. Somewhat better constraints are available for the process of Gondwana assembly at the end of the Proterozoic (Meert 2001; Powell & Pisarevsky 2002; Sanchez Bettucci & Rapalini 2002; and many others). The kinematics of each continental block, their relations with neighbouring blocks, and timing and mode of accretion are known only approximately, even in the best cases. In particular, South America seems to have undergone a complex tectonic history from the Late Proterozoic to the Late Palaeozoic (see, for example, Pankhurst & Rapela 1998; Ramos & Keppie 1999; Cordani et al. 2000; Kroner & Cordani 2003). This region witnessed the amalgamation of several Late Archaean to Early Proterozoic cratonic nucleii during Gondwana assembly and, later on, constituted an active margin, with a very complex tectonic evolution that apparently involved the accretion and displacement of several terranes. In the last decade or so, important debates have taken place on the origin, time and mode of accretion of these terranes (e.g. Ramos et al. 1986; Dalla Salda et al. 1992; Astini et al. 1995; Conti et al. 1996; Balburgh & Herve 1997; Pankhurst & Rapela 1998; Acenolaza et al. 2002; Astini & Rapalini 2003). Some of these debates have not yet been settled but, in any case, there seems to be little doubt that the Palaeozoic tectonic evolution of the southwestern Gondwana (South American) margin cannot be separated ultimately from the final stages of Gondwana assembly. Therefore, a broad picture of the formation of this supercontinent

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should take into consideration the possible accretion and displacement of allochthonous and/or parautochthonous terranes along its southwestern margin during the Early to Late Palaeozoic. This contribution is aimed at presenting the available palaeomagnetic information regarding the assembly of southern South America and discusses the implications and shortcomings of the data. The discussion will focus on the Rio de la Plata craton and the terranes of Pampia, Cuyania (Precordillera), Famatina-Eastern Puna, Chilenia, Antofalla and Patagonia (Fig. 1). A previous review by Rapalini et al. (1999) was restricted exclusively to the Palaeozoic Andean terranes. A more comprehensive review is intended here, as well as the incorporation of palaeomagnetic and geological evidence obtained since then.

The Rio de la Plata craton The Rio de la Plata craton (RP) is a relatively small block that comprises parts of southern Brasil (states of Santa Catarina and Rio Grande do Sul), Uruguay and most of the province of Buenos Aires in Argentina (Fig. 1). It is bounded to the east by the supposedly allochthonous Rocha, Cuchilla Dionisio or Punta del Este terrane in Uruguay (Fragoso Cesar 1980; Basei et al 2000), a possible Kalahari remnant in South America, and the Pampia terrane (Ramos et al. 1993) to the west. A large proportion of the western part of the RP is under the thick Cenozoic cover of the Chaco-Pampean plains and is virtually unknown. In a Gondwana reconstruction it is surrounded by Pampia, Patagonia, Kalahari and Congo-Sao Francisco blocks. The timing and sequence of events that led to the assembly of these blocks during the formation of Gondwana has been a matter of great controversy and no consensus has been reached yet (e.g. Grunow et al. 1996; Prave 1996; Basei et al 2000; Passchier et al. 2002; Kroner & Cordani 2003, among many others). Reliable palaeomagnetic data from the RP can help to discriminate between different hypotheses. Late Proterozoic to Early Palaeozoic palaeomagnetic poles for the RP are presented in Table 1. Sanchez Bettucci & Rapalini (2002) recently obtained two palaeomagnetic poles from the Sierra de las Animas

Fig. 1. Simplified map of main Late Proterozoic to Palaeozoic terranes in southern South America. Question marks indicate terrane boundaries with hypothetical locations or unknown tectonic relationships. NP, North Patagonian Massif; D, Deseado Massif. Simplified from Ramos (1988; 1995), Conti et al. (1996), Bahlburg & Herve (1997) and Chernicoff & Zappettini (2003).

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Table 1. Summary of palaeomagnetic poles from the accreted blocks of southern South America Terrane Rio de la Plata craton

Formation Sierra de las Animas Complex (SA1) Sierra de las Animas Complex (SA2) Sierra de los Barrientos clay stones (LB)

Campo Alegre lavas (CA)

Cuyania

Cerro Totora F. (CT)

Pavon F. (PV)

Famatina-Eastern Puna

Western Puna

Patagonia

Palaeomagnetic pole (PP) Lat. Long. A95 (dp/dm)

5.9° 338.1° (19.6°/26.7°) -16.9° 250.9° (15.9°/21.5°)

-8.4° 242.6° (10.5°/15.9°)

-57.0° 223.0°

9.0°

37.0° 314.1°

5.8°

3.6° 346.4° (2.9°, 4.6°)

Acoyte F., Chiquero F. (PO)

-17.7° 357.9° 13.0°

Cuchiyaco Gr. (CY)

-30.0°

SuriF.(FS)

-13.0° 356.3°

6.0°

Overall mean (PF)

-21.8° 358.7°

7.5°

0.5° 10.0°

Pluton de Choschas + lavas presiluricas (CHO)

58.1° 309.2° (3.8°/7.5°)

Chachas and Taca-Taca plutons + Vega Pinato ignimbrites (CHA) Tucucaro, Tilipozo and Alto del Inca Plutons (OP) PreSilurian lavas (OL)

16.9 260.3 (20.4°/27.8°) 0.0° 271.0° (12.4°/13.7°) -25.2° 260.5° (13.7°/15.0°) -59.1° 289.0° (19.1°/22.0°)

Las Vicunas Fm + Vega Pinato rhyolites (FV) Sierra Grande F. (upper member) (SG2) Tepuel Group (TP)

-42.0° 283.5° 20.0°

Chilean Intrusives (CI)

-57.4° 323.5° 19.0°

Sierra Grande F. (SG3)

-77.3° 310.7° (7.0°/8.0°)

Choiyoi F. (CHY)

-21.0° 232.3°

-31.7° 316.1° (15.0°/16.0°)

8.0°

Palaeomag. results*

AgeofPPt

AF/Th demag, N = 7 (n = 33), IRM. Q = 3 AF/Th demag, N = 6 (n = 27), IRM, both polarities. Q = 4 AF/Th demag, N= 12 (n = 56), both polarities, reversal of EMF recorded. Q = 6 AF/Th demag, N = 6 (n = 46), both polarities. Q = 5 AF/Th demag, +FT, N = 10 (n = 65), IRM, Q =5 AF/Th demag, +FT, n = 22(N = 4), both polarities, AMS, Q = 5 AF/Th demag, +FT, IRM, N = 5 (n = 21), Q =4 AF/Th demag, IRM,

Camb. (520-500 Ma) L. Vendian (550 Ma)

Ref.5 1 1

Vendian

2

E. Vendian (595 Ma)

3

E. Camb. (525 Ma)

4

L. Ord. (455 Ma)

5

E. Ord.

6

E. Ord.

6

AF/Th demag, N = 6 (n = 18), Q - 3 +FT, Consistency test

E. Ord.

6

E. Ord.

6

AF/Th demag, both polarities, n - 24 samples, Q = 4 AF/Th demag, tilt test, N = 6 sites, Q = 4

Latest Camb. (500 Ma)

7

E.-M. Ord. (480-465 Ma)

8

AF/Th demag, tilt test, n = 20 samples, Q = 3 AF/Th demag, both polarities, n = 20 samples, Q = 3 AF/Th demag, tilt test, TV = 6 sites, Q = 3 Th/Ch demag, +FT, N-2(n = 10), Q = 5 AF/Th demag, +FT, IRM, detrital rem.,

L. Ord.

7

Silurian ?

7

L. Ord.-E. Silurian E. Dev.

8,9

M. Carb.

11

L. Carb.-E. Perm.

12

L. Perm.?

13

L. Perm.

14

AF/Th demag, both polarities, N = 7 (n = 58), Q = 4 AF/Th demag, IRM, FT syntectonic, N=13 (n = 88), Q = 4 AF/Th demag, +FT,

10

* AF/Th demag means complete demagnetization by alternating field or thermal cleaning; +FT refers to positive fold test; IRM, isothermal remanence experiments; AMS, anisotropy of magnetic susceptibility control of remanence; N, number of sites; n, number of samples; Q refers to palaeomagnetic quality index based on the number of reliability criteria proposed by Van der Voo (1990) fullfilled by each palaeomagnetic pole. t£, Early; M, Middle; L, Late. § References: 1, Sanchez Bettucci & Rapalini (2002); 2, Rapalini (2003); 3, D Agrella & Pacca (1988); 4, Rapalini & Astini (1998); 5, Rapalini & Cingolani (2004); 6, Conti et al (1996); 7, Forsythe et al (1993); 8, Rapalini et al (2002); 9, Rapalini etal (1999); 10, Rapalini & Vilas (1991); 11, Rapalini etal (1994); 12, Beck etal (1991); 13, Rapalini (1998); 14, Rapalini et al. (1989).

ACCRETION HISTORY OF SOUTHERN S. AMERICA

magmatic complex, exposed in the Piriapolis area, southern Uruguay. These two poles, of ages c. 550-560 Ma and c. 520-500 Ma, respectively, are consistent with a single apparent polar wander path (APWP) for Gondwana (e.g. Meert & Van der Voo 1996; Meert 2001). In particular, SA2 is consistent, within error, with the 547 Ma Sinyai dolerite pole (Meert & Van der Voo 1996) for the Congo craton and the 560-540 Ma pole for the Mirbat sandstones (Kempf et al 2000) from the Arabian block (Fig. 2a, b). These three contemporaneous and coincident poles in a Gondwana reconstruction indicate that, by c. 550 Ma, the RP, Congo and Arabia blocks were already in a Gondwana configuration, suggesting that most of Western Gondwana was already assembled. The palaeo-

309

pole position for the RP has been confirmed recently by Rapalini (2003) with a high quality pole from the Sierra de los Barrientos, although the age of this pole, while compatible with 550-560 Ma, is less certain. Meert & Van der Voo (1996), Meert (2001) and Sanchez Bettucci & Rapalini (2002), among others, proposed that most Western Gondwana blocks were already assembled by c. 550 Ma, based on the coincidence of palaeomagnetic poles of that age or younger in a Gondwana reconstruction. The latter authors have suggested that some of the blocks may have been contiguous since at least 600 Ma. Less reliable poles of approximately that age from the RP (CA, D'Agrella & Pacca 1988), the West Nile craton (BS, NB, Saradeth et al 1989), and

Fig. 2. (a) Palaeomagnetic poles for the Rio de la Plata craton (RP, see Table 1) and other Gondwanan continents, Sinyai Dolerite (SD), (Meert & van der Voo 1996). The approximate ages along the Gondwana path are indicated, (b) Vendian to Cambrian palaeomagnetic poles from the RP, Congo, West Nile, Arabia and West Africa blocks in a Gondwana reconstruction (Lottes & Rowley 1990). Note the coincidence of poles of c. 600 Ma for the RP, West Nile and Arabian blocks and their disagreement with a pole of similar age from the West African craton; Ntonya Ring (NR, Briden et al 1993); Mirbat sandstone (MB, Kempf et al. 2000); AJ (Kellogg & Beckmann 1983). (c) Laurentia (LAU), Western Africa (WA), Amazonia (AM) and Pampia (PA) in a Rodinia reconstruction similar to those of Hoffman (1991) and Weil et al (1998), and palaeomagnetic pole positions for c. 615 Ma poles for Laurentia (Long Range (LR) dykes, Murthy et al 1992) and for Western Africa (AD, Adma Diorite, Morel 1981). Mean palaeomagnetic pole for Laurentia (580) for c. 580 Ma (Meert et al 1994) is also shown. North American coordinates. See discussion in the text. (Modified from Sanchez RptturH & Ranalrni 9009 ">

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A. E. RAPALINI

Arabia (AJ, Kellogg & Beckmann 1983) are coincident in a Gondwana reconstruction (Fig. 2b). This was interpreted by Sanchez Bettucci & Rapalini (2002) as evidence that these blocks, plus the Congo-Sao Francisco craton, that lies in between, formed a single continental mass for those times. In a simplified model of the assembly of Gondwana these authors proposed that these blocks formed a continent that they named Central Gondwana (following previous ideas by Trompette 1997 and Meert 2001). On the other hand, the Amazonian and West African cratons may have remained attached to Laurentia (as remnants of Rodinia) until very late in the Neoproterozoic. The oldest palaeomagnetic pole from any of these blocks that is consistent with the Gondwana APWP is that from the Ntonya ring structure of 522 Ma (Briden et al 1993). Furthermore, the c. 610 Ma pole from the Adma diorite (AD, Morel 1981; Fig. 2), in the West African craton, does not agree with similar-aged poles from other Gondwana blocks (Fig. 2b), suggesting that, at that time, West Africa and Amazonia were not part of the same continent with them. The AD pole is consistent, however, with the Long Ranges dykes pole (c. 615 Ma, Murthy et al 1992) for Laurentia, if West Africa and Amazonia are kept in a Rodinia-like configuration (Fig. 2c). Although both AD and LR poles are not without controversies, and remagnetizations cannot be ruled out yet, they are suggestive of a separate evolution of Amazonia and West Africa from the remaining Western Gondwana blocks until nearly the end of the Proterozoic. Sanchez Bettucci & Rapalini (2002) interpreted that a major ocean was consumed between the RP and Congo-Sao Francisco on one side and Amazonia and West Africa on the other. This has been suggested independently, mainly on geological grounds, by Kroner & Cordani (2003) and Cordani et al (2003), who proposed a large and long-lived ocean named the Pampian-Goias-Pharusian Ocean. The Rio Apa (in Paraguay) and Pampia blocks, for which no palaeomagnetic data are available (see below), are considered to be associated with Amazonia and West Africa (Brito Neves et al 1999). The palaeomagnetic data are also compatible with the proposal of Cawood et al (2001) of final separation of Amazonia (and Pampia?) from Eastern Laurentia around 570 Ma. Drifting apart of these blocks probably caused final closure of the Pampian-Goias Ocean, of which collision of Pampia against the RP in the Early Cambrian may be seen as the last event (see below). The picture is further complicated when the East

Gondwana blocks are considered. Previous ideas of a single East Gondwana block during the Neoproterozoic (Powell et al 1993) no longer seem tenable. As proposed recently by Powell & Pisarevsky (2002) and Meert (2003), Eastern Gondwana may have not been completely assembled until the very end of the Proterozoic. Furthermore, the palaeogeographical evolution of the Kalahari block in southern Africa is very much disputed and no palaeomagnetic data are available to constrain its evolution during the period under consideration. Nevertheless, most models suggest that Kalahari underwent somewhat independent kinematic evolution until its collision with the Congo and RP cratons by the end of the Proterozoic (e.g. Prave 1996; Basei et al 2000; Passchier et al 2002). Recent advances in the palaeomagnetism of late Proterozoic formations of the Rio de la Plata craton, and the existence of several potentially suitable units exposed from Argentina to southern Brasil suggest that a better and more detailed palaeogeographical evolution of this block and its role in the assembly of Gondwana may not be so far away. Pampia To the west of the Rio de la Plata craton is a terrane named Pampia (Ramos 1988; Ramos et al 1993; Kraemer et al 1995) of over 1500 km long and 250 km wide (without palinspastic restoration, Fig. 1). This terrane represents the basement of most of the Sierras Pampeanas of Argentina. Its structure and lithological content is best exposed along the Sierras de Cordoba, around 32° S (Martino et al 1995; Baldo et al 1996), but it is inferred that it extends over a thousand kilometres up to the extreme north of the Salta province of Argentina (22° S) and several hundred kilometers to the La Pampa province in the south (38° S, Chernicoff & Zappettini 2003). A belt of ultramafic and mafic rocks, poorly dated as Late Neoproterozoic, named by Ramos et al (2000) as the Western Cordoba belt, extends discontinuously for 150 km and has been interpreted as an ophiolite (Villar 1975) that may constitute its eastern boundary. Studies by Escayola et al (1996) indicated that these rocks have a N-Morb signature and were interpreted by Kraemer et al (1995) as remnants of oceanic floor obducted during collision of the Pampia terrane with the Rio de la Plata craton. A second belt of ophiolites, located some 50 km to the east, shows back-arc geochemical signatures (Escayola et al 1996) and these have been interpreted as remnants of

ACCRETION HISTORY OF SOUTHERN S. AMERICA

a back-arc basin closed on the western margin of the Rio de la Plata craton (Ramos et al 2000). To the west, Pampia is bounded by the Laurentian-derived Cuyania terrane along a suture located on the Valle Fertil megashear and characterized by a belt of mylonitic rocks with conspicuous Early Palaeozoic deformation (e.g. Ramos et al 1998), that have been related to docking of the Cuyania terrane. The basement of Pampia is represented mainly by the thousands of metres of metasediments with some intercalations of metavolcanics of the Puncoviscana Fm. (e.g. Acenolaza & Toselli 1981), with ages that encompass the Late Vendian to the Early Cambrian. These rocks are better exposed in NW Argentina, while supposedly equivalent rocks representing deeper structural levels are exposed towards the south (Rapela et al 1998a). The Puncoviscana Fm. is interpreted generally as a passive margin depositional sequence (Jezek et al 1985), although radically different interpretations have also been published (e.g. 'foreland basin deposits' of Keppie & Bahlburg 1999). These rocks were folded tightly in the Early (?) Cambrian (Tilcaric event, see Astini 2003), showing a superb angular unconformity with the Middle to Late Cambrian clastic sediments of the Meson Group in NW Argentina. Detailed geochronological studies by Rapela et al (19980) have shown that an active margin existed in the Early Cambrian (c. 530 Ma) to the east of Pampia (on the Rio de la Plata craton ?). This was followed shortly by a major collisional event (Tilcaric event) that produced spectacular deformation of the Puncoviscana sedimentary rocks in NW Argentina and crustal thickening and high-grade metamorphism to the south (c. 525 Ma). Finally, isothermal uplift and widespread low-P anatexis is recorded in the Sierras de Cordoba at around 520 Ma. This evolution has been interpreted by Rapela et al (19980) as the signature of the collision of Pampia against the Rio de la Plata craton, which is then dated as Early Cambrian. Very recently, Schwartz & Gromet (2004) presented U-Pb ages on detrital zircons from Late ProterozoicEarly Cambrian metasediments from the Sierras de Cordoba, which are generally interpreted as the southward continuation of the Puncoviscana basin. A dominant peak of Mesoproterozoic (950-1050 Ma) ages suggest that the source areas cannot be located on the RP and are probably on Amazonia, ArequipaAntofalla or the Kalahari craton. The latter case, favoured by Schwartz & Gromet (2004), would imply a post-Early Cambrian significant displacement of Pampia along the Gondwana

311

margin. This is still a highly speculative scenario but, in any case, these results support Pampia as an allochthonous block accreted at some time during the Cambrian. As mentioned above, collision of Pampia probably was among the last stages of final Gondwana assembly and produced the final closure of the Pampian-Goias-Pharusian Ocean between West African-Amazonian and Central Gondwana blocks. However, there are no available Late Proterozoic to Early Palaeozoic palaeomagnetic data from Pampia to constrain its evolution and test current tectonic models.

Famatina-Eastern Puna magmatic arc An outstanding geological feature of northwest Argentina is a north-trending belt of Lower Ordovician magmatic rocks (Fig. 1) that comprises the Famatina system and the 'Faja Eruptiva' of Eastern Puna (Puna Oriental). The age of this belt is very well defined, by means of fossil assemblages found in intercalated sediments, as Tremadoc-Early Llanvirn (Coira 1973; Koukharsky & Mirre 1974; Coira & Koukharsky 1991; Acenolaza 1992; Vaccari et al 1992, among others), as well as by radiometric ages from intrusive bodies (Lork & Bahlburg 1993; Pankhurst et al 1998; Rapela et al 1998ft). This volcanism has been assigned to a magmatic arc on the basis of its geochemical signature and geotectonic setting (Coira et al 1982; Ramos 1986; Mannheim 1993; Toselli et al 1996; Saavedra et al 1998; Rapela et al 1998ft). According to Astini et al (1995), the Famatina volcanic arc (and its probable northward extension in Eastern Puna) was produced by the consumption of oceanic crust associated with the approach of the Precordillera (Cuyania) terrane towards the Gondwana margin (see below). Since the original paper by Conti et al (1996) and the review by Rapalini et al (1999) no new palaeomagnetic data have been obtained from this magmatic belt. Therefore, from a palaeomagnetic point of view no significant changes in interpretation have occurred in the last few years. Conti et al (1996) found a consistent pretectonic magnetization in different units of the same age along this belt, and computed three palaeomagnetic poles from four localities of Early Ordovician age for the Puna Oriental (Eastern Puna), Cafayate area (Cuchiyaco Granodiorite) and the Famatina system (Table 1). The same magnetization direction was found in different lithologies carried by different magnetic minerals. The Early Ordovician palaeomagnetic poles from both the Famatina

312

A. E. RAPALINI

system and the Eastern Puna belt are coincident and indicate no palaeolatitudinal anomaly with respect to Gondwana, suggesting that it was a peri-Gondwana magmatic arc. However, they show (Fig. 3) a systematic anomaly in declination that suggests a clockwise rotation of 52.6° ± 11.1°. This was interpreted by Conti et al (1996) as evidence of structural continuity between Famatina and the Eastern Puna. These authors proposed that this was a short-lived magmatic arc that rotated as a single terrane before its accretion to the Gondwana margin, probably in Middle Ordovician times. This model is consistent with early proposals of the Famatina (and Eastern Puna?) arc as being developed on oceanic or quasi-oceanic crust (Ramos 1986; Toselli et al 1996), although recently these interpretations have been disputed greatly. Recent palaeomagnetic data on the Late Ordovician Pavon Fm. (Rapalini & Cingolani 2004, see below) allow an interpretation of the kinematics of accretion of the Laurentian Cuyania terrane that is consistent with this model. A similar palaeogeographical reconstruction to that produced from palaeomagnetic data was obtained for the Famatina system by Benedetto & Sanchez (1996) and Astini & Benedetto (1996) on the basis of biogeographical considerations. However, this model has been disputed by Pankhurst et al (1998), Saavedra et al (1998), Rapela et al (1998ft) and Astini (2003) among others, in favour of a model of a continental magmatic arc

Fig. 3. Comparison of the Early Ordovician palaeomagnetic poles from the Famatina-Eastern Puna terrane with the 475 Ma mean palaeomagnetic pole for Gondwana in a Gondwana reconstruction. Details of palaeomagnetic poles in Table 1. (Modified from Conti et al 1996.)

developed on the Gondwana margin. This latter model is based mainly on the geochemical signature of the Famatinian magmatic products which, in general, do not support an intraoceanic environment for this magmatism. If Famatina (and Eastern Puna) is an Andeantype magmatic arc, the interpretation of the Early Ordovician palaeomagnetic data by Conti et al (1996) as a rotated peri-Gondwanan terrane cannot be sustained, and independent large crustal block rotations of the same sense and similar magnitudes for each of the four palaeomagnetic localities should be considered. New and more systematic palaeomagnetic data are needed for a definite answer to this controversy. Rapela et al (1998ft) and Miller et al (2003), however, have proposed a tectonic model for the Famatina system that involves a volcanic arc on continental crust but separated from mainland Gondwana by a back-arc system with development of oceanic crust that was closed in the Middle-Late Ordovician (Rapela et al 1998ft) or Early Devonian (Miller et al 2003). This latter model, if confirmed, could be compatible with both the palaeomagnetic and geochemical data. In any case, continuity of the Famatinian arc into the Eastern Puna magmatic belt also needs further palaeomagnetic confirmation. Cuyania (Argentine Precordillera) terrane The Argentine Precordillera has been interpreted as an Early Palaeozoic Laurentianderived exotic terrane (e.g. Dalla Salda et al. 1992; Benedetto 1993; 1998; Astini et al 1995; Mahlburgh Kay et al 1996; Thomas & Astini 1996; Dalziel 1997; Keller et al 1998). Ramos (1995) proposed that the Precordillera is part of a larger composite terrane that he named 'Cuyania' (Fig. 1) and that includes the San Rafael Block and the Pie de Palo Range in the Western Sierras Pampeanas (Fig. 1). These latter two areas are characterized by Grenvillian-aged (c. 1.1 Ga) basement that is consistent with the age of basement xenoliths in Tertiary volcanics in the Precordillera itself (Mahlburgh Kay et al 1996). Most authors agree that Cuyania most likely originated from the Ouachita Embayment of North America. However, there are models which interpret the origin of the Precordillera to be autochthonous (Gonzalez Bonorino & Gonzalez Bonorino 1991) or that it represents a parautochthonous displaced terrane (Baldis etal 19890; Acenolaza et al 2002; Finney et al 2003). See also Astini & Rapalini (2003) and Acenolaza et al (2002; 2003) for opposite views on this topic.

ACCRETION HISTORY OF SOUTHERN S. AMERICA

The Argentine Precordillera is undoubtedly the most thoroughly studied terrane of southwestern South America, in part, at least, due to the exceptional outcrops of Early Palaeozoic unmetamorphosed sedimentary successions. These rocks also have a remarkable biostratigraphic control (e.g. Baldis et al. 1989b; Benedetto 1993; Benedetto et al 1999). A thorough review of the publications produced around the topic of the origin and evolution of the Cuyania terrane is outside of the scope of this paper and the reader is referred to recent volumes by Pankhurst & Rapela (1998), Ramos & Keppie (1999) and Benedetto (2003) for indepth reviews. There is significant evidence in favour of an allochthonous (Laurentian) origin for the Argentine Precordillera. 1.

2.

Biogeographical evidence (Benedetto 1993; 1998; Benedetto et al 1999; Canas 1999). The Cambrian successions of the Argentine Precordillera contain benthic faunas (trilobites, brachiopods, sponges) of exclusively Laurentian affinity (particularly Appalachian) that contrast strongly with assemblages in neighbouring geological areas in South America (e.g. Cordillera Oriental, Sierras Subandinas, on Pampean basement). These faunas include archetypal Olenellus fauna of Early Cambrian times. Some faunas in the Precordillera and Southern Appalachians have been correlated to the level of genus and species. In Early Ordovician times these faunas start to diverge from each other with an increasing number of endemic genera through the Ordovician. The highest level of endemicity is attained during the Caradoc (Benedetto 1998; Keller et al 1998; Benedetto et al 1999). By Silurian times the Precordilleran faunas are virtually indistinguishable from those typical of Gondwana. Stratigraphic evidence (Ramos et al 1986; Astini et al 1995; Astini 1998). The Cambrian to Ordovician Stratigraphic succession of the Argentine Precordillera is strikingly comparable with that of the Southern Appalachians, including Early Cambrian syn-rift deposits that are covered by transgressive successions of limestones and dolomites. Very similar subsidence curves (Bond et al 1984; Astini 1998) to those expected for conjugate margins are further compelling pieces of evidence of previous juxtaposition of Cuyania and SE Laurentia.

3.

313

Age and isotopic signature of the basement (Mahlburgh Kay et al 1996; Vujovich & Mahlburgh Kay 1998). Basement rocks of the Cuyania terrane exposed in the San Rafael block and Pie de Palo ranges show dominant Grenvillian ages between 1.0 Ga and 1.1 Ga. A more extended age span is suggested by the 1.2 Ga age for the Las Matras intrusion in the La Pampa province, considered to be the southernmost extension of Cuyania (Sato et al 2000). These Grenvillian ages have also been found in basement xenoliths in Tertiary volcanics of the Argentine Precordillera and are very much consistent with dominant ages of the Appalachian basement (e.g. Mosher 1998). Exclusively Grenvillian ages have also been reported in detrital zircons of Ordovician clastic rocks in Precordillera (Finney et al 2003), although in this case they have been interpreted differently. Pb/Pb isotopes of the Precordilleran xenoliths have also been found to resemble those from the Southern Appalachians, and the Llano uplift, and are significantly different from the lead isotope signatures of other neighbouring geological provinces of southern South America (Mahlburgh Kay et al 1996).

The fourth significant line of evidence in favour of an allochthonous (Laurentian) origin for Cuyania is palaeomagnetic (Rapalini & Astini 1998; Rapalini & Cingolani 2004). Three different tectonic scenarios have been postulated for the transference of Precordillera (Cuyania). The first was postulated by Dalla Salda et al (1992) and implied a continental collision between Laurentia and Gondwana in mid (Late?) Ordovician times, followed shortly after by a separation leaving the Laurentian Precordillera attached to Gondwana as a tectonic tracer (Dalziel 1993). This hypothesis has been refuted, mainly on the basis of its incompatibility with the biogeographical evolution that suggests a progressive separation of Cuyania from Laurentia during Cambrian and Ordovician times and the lack of mixing of Laurentian and Gondwanan faunas in midOrdovician times (e.g. Benedetto 1998). Furthermore, Thomas etal (2002) recently have also postulated that main Ordovician orogenesis in eastern Laurentia (Taconic) and western South America (Ocloyic) are not contemporaneous. A second model, proposed originally by Benedetto (1993) and Astini et al (1995) and refined subsequently by numerous contributions (Thomas & Astini 1996; 1999; 2003;

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Thomas etal 2002; Astini 1998; Benedetto 1998; Benedetto et al 1999), proposes that the Cuyania terrane rifted apart from the Ouachita Embayment in the Early Cambrian and was transferred as a microplate across the lapetus Ocean during Cambrian and Ordovician times to become accreted to the western South America (southwestern Gondwana) margin. Canas (1999) has suggested that separation of Precordillera from Laurentia did not occur before the Early Ordovician. The Astini et al (1995) model is compatible with the biogeographical evidence as well as stratigraphic and tectonic interpretations from both the southeastern Laurentian and southwestern Gondwana continental margins for the Early Palaeozoic. However, it has been disputed by Dalziel (1997), who claimed that it is not actualistic as it would imply a very large ridge jump during the rift-drift transition stage. To avoid this problem, Dalziel (1997) proposed an alternative model that can be viewed as a modified first model. It proposes that the Argentine Precordillera was part of a Laurentian plateau during the latest Proterozoic and Early Palaeozoic, in a fashion similar to the present-day Malvinas-Falkland plateau with respect to South America. According to this model, the progressive faunal diversity found between Precordillera and Laurentia would be due to extension along the plateau. Dalziel (1997) also suggested that accretion of Precordillera was produced by a 'soft' collision between Laurentia and Gondwana in the Ordovician. Keller et al (1998) modified the Laurentian plateau model in such a way that it rests half way between those of Dalziel (1997) and Astini et al (1995). In his proposal, Cuyania, as part of the so-called Texas plateau, would have acted as the unknown source of sediments (Llanoria) already postulated for the Precambrian on SE Laurentia (see Keller et al 1998). According to this model, this plateau would have undergone very significant crustal extension during the Cambrian and Ordovician, reaching final break-up in the Caradoc, with formation of oceanic crust between Cuyania and Laurentia only at that stage. To avoid a Laurentia-Gondwana collision, this model proposes a Silurian to Devonian age for the accretion of Precordillera. Rapela et al (1998Z?) have argued in favour of that age for the accretion of Cuyania. However, virtually the same authors have later preferred an Ordovician age of accretion (Casquet et al 2001). Thomas & Astini (2003) analysed recently the different proposals, reaching the conclusion that the microcontinent hypothesis, with a Middle to

Late Ordovician age for accretion, is the most compatible with the evidence. However, the controversy is still not settled.

Palaeomagnetic evidence Several palaeomagnetic studies have been attempted in the Cambrian-Ordovician carbonate platform of Cuyania (Rapalini & Tarling 1993; Truco & Rapalini 1996; Rapalini et al. 2000; Rapalini & Astini 2005). All these studies failed to recover the original remanence of these rocks that are affected by a widespread and pervasive regional remagnetization of probable Permian age. This event has been called the SanRafaelic remagnetization (Rapalini 1993) and it has been associated with a main tectonic phase that affected large areas of western Argentina in the Late Palaeozoic. Up to now only two studies have recovered the primary remanence of Palaeozoic rocks from Cuyania successfully. They correspond to the Early Cambrian Cerro Totora Formation (Rapalini & Astini 1998), exposed in the northern Precordillera, and the Early Caradoc Pavon Formation (Rapalini & Cingolani 2004) exposed in the San Rafael block. Palaeomagnetic data and pole positions for these two formations are presented in Table 1. The Cerro Totora palaeomagnetic pole (CT) is anomalous with respect to the Gondwana APWP, indicating a much lower palaeolatitude than that expected for the present position of Precordillera in South America in a Gondwana reconstruction (Fig. 4a). On the other hand, if Cuyania is placed against the Ouachita Embayment of SE Laurentia, as proposed by Dalla Salda et al. (1992) and Thomas & Astini (1996), CT is perfectly consistent with the Early Cambrian expected position in the Laurentian path (Fig. 4b). This is a strong line of evidence in favour of the allochthonous nature and Laurentian origin of Cuyania that supports the above-mentioned biogeographical, stratigraphic and isotopic evidence. Finney et al. (2003) obtained U-Pb age data from detrital zircons of almost exclusively Grenvillian ages in Ordovician clastic units of the Precordillera. These ages match those expected from the Cuyania basement as well as those from the Llano Uplift in SE Laurentia (Mosher 1998). Claims by Finney et al. (2003) of a Gondwanan origin for the Argentine Precordillera on the base of detrital zircon ages from the Cerro Totora Formation have been invalidated by the fact that the published ages did not correspond to this formation due to a laboratory error, as recently acknowledged by Finney et al. (2004).

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Fig. 4. (a) Position of palaeomagnetic poles from Cuyania (Table 1) after rotation into a Gondwana palaeoreconstruction (Lottes & Rowley 1990), placing Cuyania in its present position in South America, and latest Proterozoic-Early Palaeozoic reference poles for Gondwana. (b) Position of both poles from Cuyania after rotation of the Cuyania terrane to match the Ouachita Embayment, following rotation parameters by Rapalini & Astini (1998). Reference poles as presented by Rapalini & Cingolani (2004). Ages of poles are indicated in Ma. (Modified from Rapalini & Cingolani 2004.)

All the above-mentioned lines of evidence indicate strongly that Cuyania is a Laurentianderived exotic terrane that probably originated in the Ouachita Embayment of SE North America. Recently, Rapalini & Cingolani (2004) presented new palaeomagnetic data from the Early Caradoc Pavon Fm. Positive fold and reversal tests, as well as anisotropy of magnetic susceptibility (AMS) results indicate that the PV pole is a reliable recorder of the Late Ordovician palaeomagnetic field. PV is also anomalous with respect to its expected position in the Gondwana path (Fig. 4a). However, in this case the anomaly is only in decimation and can be solved assuming a moderate (around 30°) clockwise rotation of the study area. A secondary remanence isolated from the same rocks suggests that if such rotation ever occurred it must have happened in prePermo-Triassic times. A good resolution of the palaeolatitude anomaly of PV with respect to the Gondwana APWP is hampered by the low quality and large error associated with the Late Ordovician Gondwana reference pole (McElhinny & McFadden 2000). Figure 4b illustrates the result obtained by Rapalini & Cingolani (2004). If Cuyania is placed against

the Ouachita Embayment, PV is coincident with the Late Ordovician reference pole for Laurentia (McElhinny & McFadden 2000). This coincidence is even better if 500 km of stretching normal to the margin is considered between Cuyania and Laurentia. This seems to be a strong argument in favour of a Late Ordovician link between both blocks and the Laurentian plateau model proposed by Keller et al (1998). However, since no significant palaeolatitude anomaly can be determined yet between Cuyania and Gondwana, the argument is less compelling. Rapalini & Cingolani (2004) presented three possible scenarios to explain the position of the PV pole. The first is the aforementioned Laurentian plateau model. The second and third consider the consistency with the Laurentian path as merely coincidental. In the second one the 30° CW rotation is proposed to be representative of the whole of Cuyania and associated with the kinematics of accretion (therefore implied to be post-Early Caradoc) and similar to the rotation proposed for the Famatina-Eastern Puna belt (see above). The third alternative suggests a preTriassic local crustal block rotation of unknown origin that would post-date the accretion, therefore considered to be pre-Caradoc. From

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a palaeomagnetic point of view all three alternatives are equally permissible.

Chilenia Nearly two decades ago Ramos et al (1986) proposed that the basement of the Central Andes of western Argentina and Chile, to the west of the Argentine Precordillera, is an allochthonous terrane (Fig. 1) that was accreted to the western Gondwana margin in the Middle Palaeozoic. This hypothesis was based on the oceanic character (E-Morb signature) of Late Ordovician basaltic pillow-lavas, dykes and gabbros exposed along the western border of the Precordillera (Haller & Ramos 1984; Mahlburg Kay et al. 1984) that have since been interpreted as ophiolites (see Ramos et al 2000). The palaeogeographical interpretation was that an open ocean existed in the Late Ordovician to the west of Cuyania (and Gondwana) and, therefore, the basement of the main Cordillera had to be accreted later. A main tectonic phase in the Late Devonian, very well represented in several areas of western Argentina, has been interpreted as due to the collision of this terrane labelled as 'Chilenia'. Some authors, however, completely disagree with this interpretation, suggesting that the Late Ordovician ophiolites are the product of extensional processes that led to the development of an oceanic basin (e.g. Dalla Salda et al. 1992; Keller et al. 1998; Davis et al. 1999). Exposures of pre-Late Palaeozoic rocks of Chilenia are very scarce and there are no biostratigraphical data that permit any comparison with the Early Palaeozoic successions of Cuyania and Antofalla. However, U-Pb ages on zircons of 1069 ± 36 Ma (Ramos & Basei 1997) and Nd model ages of 1.4-1.7 Ga suggest a Grenvillian basement and point to a possible Laurentian link for Chilenia (Keppie & Ramos 1999). Rapela et al. (19986) have suggested that Chilenia may be part of Cuyania and not an independent terrane. There is no palaeomagnetic data for Chilenia, so far, to constrain its kinematic history and suitable rocks for palaeomagnetic studies are still to be found in this terrane.

Western Puna (Antofalla terrane) The existence of Early to Middle Proterozoic crystalline rocks along the southern coast of Peru (Dalmayrac et al. 1977; Shackleton et al. 1979), generally referred to as the Arequipa Massif, led to speculations on their origin (e.g. Ramos 1988). In several models, this block has

been extended into northern Chile to include outcrops of Middle to Late Proterozoic metamorphic rocks (see Ramos 2000, for a recent update on this). Ramos (1988) has labelled this enlarged Arequipa block as the ArequipaAntofalla Massif. However, outcrops in northern Chile are very small, disconnected and are generally younger than those in Peru, suggesting that claims of tectonic continuity with the Arequipa Massif are, at least, disputable. Furthermore, Bahlburg & Herve (1997) suggested that there may be two separate tectonic blocks: an Early to Middle Proterozoic Arequipa block in the north and a Late Proterozoic Antofalla block in the south (Fig. 1). An Early Palaeozoic magmatic belt extends through the Puna plateau of Argentina and Chile with a NNW trend between approximately 26° S and 23° S. This belt has been named 'Faja Eruptiva de la Puna Occidental' ('Western Puna Eruptive Belt', Palma et al. 1986) and continues into northern Chile along the Sierra de Almeida. It is located close to the eastern border of a hypothetical Antofalla terrane (Bahlburg & Herve 1997) and consists of Late Cambrian to Silurian plutons and Ordovician volcanics exposed at numerous localities, but for which systematic ages are still lacking. Basic to ultrabasic rocks of oceanic signature - interpreted by some authors as ophiolites (Allmendinger et al. 1983; Blasco et al. 1996; Ramos 2000) - are located in several areas of the Western Puna of Argentina, to the east of the Chilean Neoproterozoic outcrops and the Western Puna eruptive belt, and to the west of the Eastern Puna of Argentina (Fig. 1). Forsythe et al. (1993) were the first to produce palaeomagnetic data from some of these rocks, i.e. Late Cambrian to Silurian volcanics and intrusives from the Sierra de Almeida in Chile. Later, Rapalini et al. (1999) reported some palaeomagnetic results from the Early Ordovician volcaniclastic rocks of the Las Vicunas Formation in the Western Puna of Argentina. In both cases, data come exclusively from the Antofalla block. Forsythe et al. (1993) considered their original data as representative of the entire Arequipa-Antofalla block and interpreted them as supporting a tectonic model in which a parautochthonous Arequipa Massif rotated consecutively clockwise in the Proterozoic and counterclockwise in the Early Palaeozoic, opening and closing a small oceanic basin in the southern part between this block and southwestern Gondwana. Rapalini et al. (1999) cast doubt on the validity of this model and suggested that the data should be applied only to describe the kinematics of the Antofalla

ACCRETION HISTORY OF SOUTHERN S. AMERICA

block, and that the previous model did not account for a palaeolatitude anomaly of over 1000 km shown by the Late Cambrian-Early Ordovician palaeomagnetic results. Rapalini et al (2002) recently presented new palaeomagnetic data from Early Palaeozoic units of the Western Puna. These new data come from Tremadoc to Arenig hyaloclastites and rhyolites exposed at Vega Pinato, the Early Ordovician (?) Chachas magmatic complex and the Early Ordovician (469 ± 4 Ma) Taca-Taca batholith. All available Early Palaeozoic palaeomagnetic poles from this region are presented in Table 1 and Figure 5. The precise age of some of these poles (e.g. (OL) Tucucaro, Tilipozo and Alto de Inca plutons, (OP) preSilurian lavas, (FV) Las Vicunas Fm. and Vega Pinato rhyolites) is disputable, while palaeohorizontal control of some of the sites in others (e.g. (CHO) Choschas pluton, (CHA) Chachas

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complex) is dubious. In any case, the data are only representative for the Antofalla block. The overall distribution of palaeomagnetic poles suggests that they define the Late Cambrian to Silurian APWP for Antofalla as originally proposed by Forsythe et al. (1993). However, the FV pole position is problematic, as the rocks from which it was computed are dated accurately as Tremadoc-Arenig and the data yielded a positive fold test suggesting that magnetization pre-dates the Ocloyic (latest Ordovician) tectonic event. A possible explanation is that it represents a pre-tectonic remagnetization in the latest Ordovician (or Early Silurian), associated with the Ocloyic tectonic phase (Rapalini et al 1999). The tectonic interpretation of the available palaeomagnetic data is controversial. The original model of Forsythe et al (1993) suggests a counterclockwise rotation of a parautochthonous

Fig. 5. Early Palaeozoic palaeomagnetic poles for the Western Puna (Antofalla terrane, Table 1) and coeval reference poles for Gondwana with its hypothetical apparent polar wander path. Numbers indicate age of the reference poles in Ma. All poles have been rotated into African coordinates according to Lottes & Rowley (1990). (Modified from Rapalini et al. 2002.)

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Arequipa-Antofalla terrane. However, Rapalini et al (1999) pointed out that the CHO pole of c. 500 Ma shows a large palaeolatitude anomaly of at least 1000 km. The new Early to Middle Ordovician CHA pole does not show that anomaly in palaeolatitude, although its oval of confidence precludes determining anomalies of less than 10°. Nevertheless, a significant counterclockwise rotation with respect to the available Gondwana reference poles for the Early and Middle Ordovican is evident. The lack of well-defined reference poles for Gondwana for the Middle and, particularly, the Late Ordovician precludes the unambiguous determination of whether the OP and OL poles show any relative movement between Antofalla and Gondwana. The FV pole is consistent with the Silurian and Early Devonian poles of Gondwana. Despite the uncertainties and ambiguities, the simplest interpretation of the palaeomagnetic data obtained so far suggests some kind of relative displacement between Antofalla and Gondwana prior to the latest Ordovician, which includes large counterclockwise rotation and, perhaps, significant latitudinal displacement. No matter how speculative are the tectonic interpretations of the available data, the palaeomagnetic information retrieved so far from Early Palaeozoic units in the Puna shows a clear difference between sites in the Western and Eastern Puna magmatic belts (see above). This probably implies a major tectonic discontinuity for the Early Palaeozoic between the Western and Eastern Puna. Conti et al (1996) found systematic clockwise rotations with no palaeolatitude anomaly in Early Ordovician rocks from the Eastern Puna and Famatina. The Antofalla block, on the other hand, shows a systematic pattern of counterclockwise rotations of the Early Palaeozoic units and a possible palaeolatitude anomaly for the Late Cambrian-Early Ordovician. Rapalini et al. (1999) discussed in some detail the evidence available in favour of separate Antofalla and Arequipa blocks, as originally proposed by Bahlburg & Herve (1997). A major tectonic discontinuity between Western and Eastern Puna magmatic belts has also been suggested by many authors previously, either as consumed oceanic crust (Coira et al. 1982; Allmendinger et al. 1983; Forsythe et al. 1993; Blasco et al. 1996) or as a major strike-slip fault (Coira et al. 1999). The Western Puna magmatic belt is also approximately coincident with a conspicuous positive gravity anomaly (Gotze & Kirchner 1997; Omarini et al. 1999), which has been interpreted as an Early Palaeozoic or Late

Proterozoic terrane boundary (Gangui & Goetze 1996; Omarini et al. 1999). However, this mobilistic view of the tectonic evolution of this region has been challenged by Lucassen et al. (2000; 2001) on the basis of isotopic and geochemical data from Precambrian and Palaeozoic metamorphic and igneous rocks. These authors interpreted that all magmatic rocks in the region were produced from reworked continental crust of similar characteristics with minor addition of juvenile material. Zimmermann & Bahlburg (1999) and Astini (2003) recently also suggested a main authochthonous and ensialic Palaeozoic tectonic evolution for this region. Furthermore, in a recent paper Astini & Davila (2004) suggested that the Western Puna magmatic belt could be a northward continuation of the Famatina system, although this is difficult to reconcile with the available palaeomagnetic information. The conflicting evidence mentioned above indicates that a precise reconstruction of the palaeogeographical and tectonic evolution of the Western Puna in the Early Palaeozoic is still elusive, due in part to the scarcity of reliable isotopic ages and palaeomagnetic results.

Patagonia Since the early twentieth century, geologists have recognized the particular geological characteristics of Patagonia. This led Windhausen (1931) to consider, well before plate tectonics, that Patagonia had evolved in some way isolated from the rest of South America. The extra-Andean Patagonia comprises two nucleii of Proterozoic to Early Palaeozoic crystalline rocks: the North Patagonian or Somuncura Massif in the north and the Deseado Massif in the south. To a large extent these rocks are covered by sedimentary rocks and volcanics of Late Palaeozoic to Cenozoic age. The tectonic relationship between both massifs is virtually unknown (Ramos 2002). Martinez (1980) and Dalmayrac et al. (1980) proposed that Patagonia underwent a major displacement along the SW Gondwana margin during the Late Palaeozoic. This was followed by Ramos (1984), who demonstrated that such an idea was untenable and proposed a radically different hypothesis. His proposal (see also Ramos 1988) is that Patagonia was an independent microplate from Gondwana in the Early to Middle Palaeozoic that underwent a frontal collision with it in the Late Palaeozoic, producing folding and thrusting on the Sierra de la Ventana fold system in the SW of the Buenos

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Aires province (Argentina) and possibly the Cape Fold Belt in South Africa (Winter 1984). In this hypothesis, poorly dated Late Palaeozoic calk-alcaline magmatic rocks in the North Patagonian Massif were interpreted as the magmatic arc developed previous to the collision. Palaeomagnetic data obtained so far from Palaeozoic rocks in Patagonia have been reviewed by Rapalini (1998) and Rapalini et al (1999). No new data have been produced since then. Available palaeomagnetic poles are presented in Table 1 and Figure 6. As already concluded by the above-mentioned authors, four out of six palaeomagnetic poles for Patagonia, between the Devonian and Permian, are consistent with the Gondwana path. One of the anomalous poles has been interpreted as due to a local block rotation (Choiyoi Fm.) and the remaining one has not been considered reliable. The available palaeomagnetic data, therefore, are more consistent with models that favour an autochthonous origin for Patagonia (e.g. Forsythe 1982; Dalla Salda et al 1990). However, it must be considered that all data belong exclusively to the North Patagonian Massif and the uncertainties in some of the poles are large enough to allow displacements between Patagonia and Gondwana of several hundred kilometres. In particular, the Middle to Late Carboniferous Tepuel pole has a large A95 (15°) and falls on the Gondwana path somewhat to the north (Fig. 6) of the Late Carboniferous Gondwana pole. This permits a separation of Patagonia and Gondwana of roughly 1000 km orthogonal to the northern boundary of Patagonia in the Middle to Late Carboniferous. An even larger separation (up to c. 1500 m) could be accommodated, given the limits of the palaeomagnetic error. The controversy surrounding the tectonic origin of Patagonia has been revived recently by contributions in favour of both positions. Gonzalez et al (2002) have recently correlated latest Proterozoic to Early Palaeozoic sediments of northern Patagonia with the Puncoviscana sedimentary cover of the Pampia terrane, and Rapela & Pankhurst (2002) have interpreted an Early Palaeozoic Gondwana link for Patagonia from U-Pb dates from detrital zircons of the El Jaguelito Fm. In contrast, von Gosen (2002) inferred a Late Palaeozoic collisional event from a structural study of NE Patagonia, favouring an allochthonous origin for this terrane. Undoubtedly new palaeomagnetic data are needed for a definite answer to this controversy. In the meantime, a par autochthonous origin of the North Patagonian Massif, that rifted away

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(to c. 1000 km) from SW Gondwana in the Late Proterozoic or Early Palaeozoic to collide in the Late Palaeozoic, seems to be the model most compatible with all lines of evidence. Whether recent reports of a Cambrian episode of rifting in the SW of the Rio de la Plata craton (Rapela et al 2003) and the Ellsworth Mountains in West Antarctica (Curtis 2001) are related to Patagonia should be investigated.

Some palaeogeographical remarks In the last two decades the classical view of South America as a single plate in preGondwana times has been replaced slowly by a model in which numerous cratonic blocks and terranes interacted in a complex manner before Gondwana was assembled completely. In particular, southern South America comprises a handful of terranes, whose relationships are known only partially and with large uncertainties. The potential of palaeomagnetism to place kinematic and temporal constraints in their evolution has not been achieved fully, given that few poles, sometimes of low reliability, have been obtained from many of these terranes. In the last decade some progress has been made, which permits a broad, albeit somewhat blurred, picture to emerge of the accretionary history of southern South America when the palaeomagnetic data are analysed together with evidence from other sources. Four schematic diagrams showing a possible palaeogeographical evolution of southern South American main blocks and their relation with Laurentia and other Gondwanan elements from the latest Proterozoic to the Late Ordovician is presented in Figure 7. As already mentioned, the picture presented is still highly hypothetical. By the late Vendian (Fig. 7a), almost all Western Gondwana blocks were already accreted, except for the Kalahari craton and Pampia blocks, although no palaeomagnetic constraints exist for them. The Cuyania terrane was part of southeastern Laurentia, which was separated from the Gondwana blocks by a large lapetus Ocean. By the late Early Cambrian (Fig. 7b) Pampia and Kalahari had just accreted into Gondwana, the assembly of which was essentially completed. A rift developed between Laurentia and Cuyania, while Gondwana moved over the south pole reducing the size of the lapetus Ocean. In the Early Ordovician (Fig. 7c) a peri-Gondwana magmatic arc developed along southwestern Gondwana (Famatina-Eastern Puna), while a position of the Western Puna terrane outboard of the Gondwana margin is proposed. A larger

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Fig. 6. Palaeozoic palaeomagnetic poles from Patagonia with their 95% confidence ovals (Table 1) and the Early Devonian-Triassic APWP of Gondwana. The Gondwana path is based on individual high quality poles from Early Devonian to Late Carboniferous and mean Western Gondwana poles from Late Carboniferous to Triassic-Jurassic boundary (see Rapalini et al. 1999). See discussion in the text. (Modified from Rapalini et al. 1999.) Dl, Lower Devonian; Du, Upper Devonian; D-C, Devonian-Carboniferous boundary; C-P, latest Carboniferous to earliest Permian; PI, Lower Permian; Pu, Upper Permian; Trm, Middle Triassic; Tru, Upper Triassic.

separation between Cuyania and Laurentia is inferred, either as a rifted-apart terrane or as a developing plateau. By the Late Ordovician (Fig. 7d) two alternative positions for Cuyania can be presented, either still attached to Laurentia as a large plateau (CY1) or already accreted to Gondwana (CY2). Conclusions Palaeomagnetic data from the Rio de la Plata craton confirm previous suggestions that most cratonic blocks of Western Gondwana were

already or close to being assembled by c. 550 Ma. Although uncertain, a major ocean may have existed by Vendian times between the Rio de la Plata-Congo-Sao Francisco and AmazoniaWest African cratons. The latter cratons may have remained attached to Laurentia until very late in the Neoproterozoic (c. 570 Ma). To the west of the Rio de la Plata craton is the Pampia terrane, which may have been part of, or associated with, Amazonia. Collision of Pampia against Rio de la Plata occurred in the Early Cambrian (c. 525 Ma) producing the obduction of the western Cordoba ophiolites.

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Fig. 7. Palaeogeographical reconstructions showing the hypothetical evolution of southern South American main blocks and their relation to Laurentia and other Gondwanan elements, from the late Vendian to the Late Ordovician: (a) late Vendian (c. 550 Ma); (b) late Early Cambrian (c. 525 Ma); (c) Early Ordovician (c. 480 Ma); (d) Late Ordovician (c. 455 Ma); AM, Amazonia; AR, Arequipa; C, Congo; CY, Cuyania; CY1, Cuyania as part of a Laurentian plateau; CY2, Cuyania as a terrane accreted to Gondwana; K, Kalahari; Lau, Laurentia; PA, Pampia; PT, Patagonia; RP, Rio de la Plata; SF, Sao Francisco; WP, Western Puna; WN, Western Nile; WA, Western Africa; FP, Famatina-Puna terrane. See discussion in the text.

However, palaeomagnetic data from this terrane are lacking. In the Early Ordovician a magmatic arc developed closed to or on the SW Gondwana margin (Pampia basement). This short-lived magmatic belt is exposed along the Famatina system and the Eastern Puna of Argentina. Palaeomagnetic data from four localities along this belt show no palaeolatitude anomaly but a systematic clockwise rotation of around 50°,

which was originally interpreted as a long single parautochthonous rotated terrane. A model of a magmatic arc developed on continental crust but separated from Gondwana by a back-arc basin that was closed when the terrane rotated is, perhaps, the most compatible with most of the evidence. The Cuyania (Argentine Precordillera) terrane has been accepted generally as a Laurentian-derived continental block. Apart

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from some recent disputes, the amount and diversity of evidence, including palaeomagnetic, in favour of such an origin remain unchallenged. The timing of accretion, however, and the tectonic mechanism for transferring Cuyania from Laurentia to Gondwana is still controversial. Recent palaeomagnetic data from Late Ordovician rocks of Cuyania support the 'Laurentian plateau' hypothesis, although they can also be reconciled with the 'microcontinent model'. The Chilenia terrane is located to the west of Cuyania, which indicates that it must be allochthonous too. Knowledge of this terrane is scarce, with no available palaeomagnetic data to constrain its evolution. Hypotheses in the literature range from an origin as an independent block that collided in Devonian times to a fragment of Cuyania transported to the west due to its collision with Gondwana. The Arequipa-Antofalla block in northern Chile, NW Argentina and SW Peru may actually be a composite terrane. The Antofalla block in the south is the only one for which Early Palaeozoic palaeomagnetic data have been obtained. Recently acquired data in the Western Puna of Argentina confirm the originally proposed pattern of palaeomagnetic poles. Despite poor reference poles and uncertainties in rock age, tectonic control and age of remanence for many of the data, a pattern of significant counterclockwise rotations is shown by most palaeomagnetic poles, with a possible anomaly in palaeolatitude for the late Cambrian. This contrasts greatly with the pattern of clockwise rotations shown by Early Ordovician rocks from the Eastern Puna, suggesting a major tectonic discontinuity between the Eastern and Western Puna of Argentina in the Early Paleozoic. The relationship between Patagonia and the Gondwana blocks to the north is a matter of acute debate. Four palaeomagnetic poles of Devonian to Permian age from the North Patagonian Massif are consistent in position and age with the Gondwana APWP, suggesting that both continental masses did not experience major relative displacement since the Devonian. However, the data do not rule out separation of Patagonia of around 1000 km orthogonal to its northern boundary in the Early Palaeozoic and subsequent collision in the Late Palaeozoic. The Universidad de Buenos Aires and the Consejo Nacional de Investigaciones Cientificas y Tecnicas (CONICET, Argentina) gave institutional support to these investigations. Further support was received from the Antorchas Foundation of Argentina. Numerous scientists contributed by different means

and at different stages to a better comprehension of the complex tectonic evolution of southern South America. Among many, the author would like to explicitly give thanks to R. Astini, C. Cingolani, M. Lopez de Luchi, P. Pazos, V. Ramos, C. Rapela, L. Sanchez Bettucci, D. Tarling and C. Vasquez. Special thanks go to Bob Pankhurst for the invitation to submit this contribution. Constructive critical reviews by C. MacNiocall, J. Meert and P. Schimdt and the editor, A. Vaughan, improved the final version substantially. Colleagues at the INGEODAV are thanked for the chaotic but fruitful working environment.

References ACENOLAZA, F.G. 1992. El sistema Ordovicico en Latinoamerica. In: GUTIERREZ MARCO, J.G, SAAVEDRA, J. & RABANO, I. (eds) El Paleozoico Inferior de IberoAmerica. Universidad de Extremadura, Spain, 85-118. ACENOLAZA, F.G. & TOSELLI, A.J. 1981. Geologia del noroeste argentino. Facultad de Ciencias Naturales, Universidad Nacional de Tucuman, Argentina, Publication Especial. ACENOLAZA, F.G, MILLER, H. & TOSELLI, A.J. 2002. Proterozoic-Early Paleozoic evolution in western South America - a discussion. Tectonophysics, 354, 121-137. ACENOLAZA F.G, MILLER, H. & TOSELLI, A.J. 2003. Reply to Discussion of Proterozoic-Early Paleozoic evolution in western South America - a discussion, by Astini and Rapalini. Tectonophysics, 366, 149-151. ALLMENDINGER, R.W., RAMOS, V.A., JORDAN, T.E., PALMA M. & ISACKS, B.L. 1983. Paleogeography and Andean structural geometry, NW Argentina. Tectonics, 2, 1-16. ASTINI, R.A. 1998. Stratigraphic evidence supporting the rifting, drifting and collision of the Laurentian Precordillera terrane of western Argentina. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 11-33. ASTINI, R.A. 2003. The Ordovician Proto-Andean basins. In: BENEDETTO, J.L. (ed.) Ordovician fossils of Argentina. Secretaria de Ciencias y Tecnologfa, Universidad Nacional de Cordoba, 1-74. ASTINI, R.A. & BENEDETTO, J.L. 1996. Paleoenvironmental features and basin evolution of a complex volcanic-arc region in the pre-Andean western Gondwana: the Famatina belt. Third International Symposium on Andean Geodynamics, St. Malo, France, Extended Abstracts, ORSTOM editions, Paris, 755-758. ASTINI, R.A. & DAviLA, F.M. 2004. Ordovician back arc foreland and Ocloyic thrust belt development on the western Gondwana margin as a response to Precordillera terrane accretion. Tectonics, 23, DOI: 10.1029/2003TC001564. ASTINI, R.A. & RAPALINI, A.E. 2003. Discussion to

ACCRETION HISTORY OF SOUTHERN S. AMERICA 'Proterozoic-Early Paleozoic evolution in western South America - a discussion'. Tectonophysics, 366, 143-148. ASTINI, R.A., BENEDETTO, J.L. & VACCARI, N.E. 1995. The early Paleozoic evolution of the Argentine Precordillera as a Laurentian rifted, drifted and collided terrane: a geodynamic model. Geological Society of America Bulletin, 107, 253-273. BAHLBURG, H. & HERVE, F. 1997. Geodynamic evolution and tectonostratigraphic terranes of northwestern Argentina and northern Chile. Geological Society of America Buletin, 109, 869-881. BALDIS, B., PERALTA, S. & VILLEGAS, C. 19890. Esquematizaciones de una posible transcurrencia del Terrane Precordillera. In: Proyecto IGCP 270, Correlation Geologica, San Miguel de Tucuman, 5, 81-100. BALDIS, B., BORDONARO, O., ARMELLA, C., BERESI, M., CABALERI, N., PERALTA, S. & BASTIAS, H. 19896. La Cuenca Paleozoica Inferior de la Precordillera Argentina. In: CHEBLI, G. & SPALLETTI, L.A. (eds) Cuencas Sedimentarias Argentinas. Universidad Nacional de Tucuman, Tucuman, Buenos Aires, 101-121. BALDO, E., CASQUET, C. & GALINDO, C. 1996. El metamorfismo de la Sierra Chica de Cordoba (Sierras Pampeanas), Argentina. Geogaceta, 19, 48-51. BASEI, M.A.S., SIGA JR., O., MASQUELIN, H., HARARA, O.M., REIS NETO, J.M. & PRECIOZZI PORTA, F. 2000. The Dom Feliciano Belt and the Rio de la Plata Craton: tectonic evolution and correlation with similar provinces of southwestern Africa. In: CORDANI, U.G., MlLANI, E.J., THOMAZ FlLHO, A.

& CAMPOS, D.A. (eds) Tectonic Evolution of South America, 31st International Geological Congress, Rio de Janeiro, Brazil, 311-334. BECK, M.E., GARCIA, A., BURMESTER, R.F., MUNIZAGA, F, HERVE, F. & DRAKE, R. 1991. Paleomagnetism and geochronology of Late Paleozoic granitic rocks from the Lake District of southern Chile: Implications for accretionary tectonics. Geology, 19, 332-335. BENEDETTO, J.L. 1993. La hipotesis de la aloctonia de la Precordillera Argentina: un test estratigrafico y biogeografico. 12th Congreso Geologico Argentino, Actas, 3, Asociacion Geologica Argentina, Buenos Aires, 375-384. BENEDETTO, J.L. 1998. Early Palaeozoic brachiopods and associated shelly faunas from western Gondwana: their bearing on the geodynamic history of the pre-Andean margin. In: PANKHURST, R.J. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 57-83. BENEDETTO, J.L. (ed.) 2003. Ordovician fossils of Argentina. Secretaria de Ciencias y Tecnologia, Universidad Nacional de Cordoba, Argentina. BENEDETTO, J.L. & SANCHEZ, T.M. 1996. Paleobiogeography of brachiopod and molluscan faunas along the South American margin of Gondwana during the Ordovician. In: BALDIS, B.A. & ACENOLAZA, F.G. (eds) Early Paleozoic Evolution in NW Gondwana, Serie Correlation

323

Geologica, San Miguel de Tucuman, Argentina, 12, 23-38. BENEDETTO, J.L., SANCHEZ, T.M., CARRERA, M.G., BRUSSA, E.D. & SALAS, MJ. 1999. Paleontological constraints on succesive paleogeographic positions of Precordillera terrane during the Early Paleozoic. In: RAMOS, V.A. & KEPPIE, J.D. (eds) Laurentia-Gondwana connections before Pangea. Geological Society of America Special Paper, 336, 21-42. BLASCO, G, VILLAR, L. & ZAPPETTINI, E.O. 1996. El complejo ofiolitico desmembrado de la Puna argentina. Provincias de Jujuy, Salta y Catamarca. 13th Congreso Geologico Argentino & III Congreso de Exploration de Hidrocarburos, Actas, 3, Asociacion Geologica Argentina, Buenos Aires, 653-667. BOND, G.C., NICKESON, PA. & KOMINZ, M.A. 1984. Breakup of a supercontinent between 625 and 555 Ma: New evidence and implications for continental histories. Earth and Planetary Science Letters, 70, 325-345. BRIDEN, J.C., MCCLELLAND, E. & REX, D.C. 1993. Proving the age of a paleomagnetic pole: the case of the Ntonya ring structure, Malawi. Journal of Geophysical Research, 98,1743-1749. BRITO NEVES, B.B., CAMPOS NETO, M.C. & FUCK, R.A. 1999. From Rodinia to Western Gondwana: An approach to the Brasiliano-Pan African cycle and erogenic collage. Episodes, 22 (3), 155-166. CANAS, F.L. 1999. Facies and sequences of the Late Cambrian-Early Ordovician carbonates of the Argentine Precordillera: A stratigraphic comparison with Laurentian platforms. In: RAMOS, V.A. & KEPPIE, J.D. (eds) LaurentiaGondwana connections before Pangea. Geological Society of America Special Paper, 336, 43-62. CASQUET, C., BALDO, E., PANKHURST, R.J., RAPELA, C.W., GALINDO, C., FANNING, C.M. & SAAVEDRA, J. 2001. Involvement of the Argentine Precordillera terrane in the Famatinian mobile belt: U-Pb SHRIMP and metamorphic evidence from the Sierra de Pie de Palo. Geology, 29, 703-706. CAWOOD, PA., MCCAUSLAND, P.J.A. & DUNNING, G.R. 2001. Opening lapetus: Constraints from the Laurentian margin in Newfoundland. Geological Society of America Bulletin, 113, 443-453. CHERNICOFF, C.J. & ZAPPETTINI, E.O. 2003. Evidencias geofisicas para la delimitation de terrenos pre Gondwanicos en la region centro-austral argentina. 10th Congreso Geologico Chileno, Conception, Actas (CD), Universidad de Concepcion, Concepcion. COIRA, B. 1973. Resultados preliminares sobre la petrologia del ciclo eruptivo concomitante con la sedimentation de la Formation Acoyte en la zona de Abra Pampa, prov. de Jujuy. Revista de la Asociacion Geologica Argentina, 27, 85-90. COIRA, B. & KOUKHARSKY, M. 1991. Lavas en almohadillas ordovicicas en el cordon de Escaya, Puna Septentrional, Argentina. 6th Congreso Geologico Chileno, 1, Servicio Nacional de Geologia y Mineria, Chile, 674-678. COIRA, B., DAVIDSON, J., MPODOZIS C. & RAMOS V.

324

A. E. RAPALINI

1982. Tectonic and magmatic evolution of the Andes of northern Argentina and Chile. Earth Science Reviews, 18, 303-332. COIRA, B., PEREZ, B., FLORES, P., MAHLBURG KAY, S., WOLL, B. & MANNING, M. 1999. Magmatic sources and tectonic setting of Gondwana margin Ordovician magmas, northern Puna of Argentina and Chile. In: RAMOS, V. & KEPPIE, D. (eds) Laurentia-Gondwana connections before Pangea. Geological Society of America Special Paper, 336, 143-170. CONTI, CM., RAPALINI, A.E., COIRA, B. & KOUKHARSKY, M. 1996. Paleomagnetic evidence of an Early Paleozoic rotated terrane in NW Argentina. A clue for Gondwana-Laurentia interaction? Geology, 24, 953-956. CORDANI, U.G., SATO, K., TEIXEIRA, W., TASSINARI, C.C.G. & BASEI, M.A.S. 2000. Crustal evolution of the South American platform. In: CORDANI, U.G., MILANI, E.J., THOMAZ FILHO, A. & CAMPOS, D.A. (eds) Tectonic Evolution of South America, 31st International Geological Congress, Rio de Janeiro, Brazil, 19-40. CORDANI, U.G., D'AGRELLA, M.S.F., BRITO-NEVES, B.B. & TRINDADE, I.F. 2003. Tearing up Rodinia: the Neoproterozoic paleogeography of South America cratonic fragments. Terra Nova, 15, 350-359. CURTIS, M.L. 2001. Tectonic history of the Ellsworth mountains, West Antarctica: reconciling a Gondwana enigma. Geological Society of America Bulletin, 113, 939-958. D'AGRELLA, M.S.F. & PACCA, I.G. 1988. Paleomagnetism of the Itajai, Castro and Bon Jardim Groups form southern Brazil. Geophysical Journal, 93, 365-376. DALLA SALDA, L., CINGOLANI, C. & VARELA, R. 1990. The origin of Patagonia. Comunicaciones, una revista de Geologia Andina, 41, 55-64. DALLA SALDA, L., CINGOLANI, C. & VARELA, R. 1992. Early Paleozoic erogenic belt of the Andes in southwestern South America: Result of Laurentia-Gondwana collision? Geology, 20, 617-620. DALMAYRAC, B., LANCELOT, J.R. & LEYRELOUP, A. 1977. Two billion year granulites in the late Precambrian metamorphic basement along the southern Peruvian coast. Science, 198, 49-51. DALMAYRAC, B., LAUBACHER, G, MAROCCO, R., MARTINEZ, C. & TOMASI, P. 1980. La Chaine Hercynienne d'Amerique du Sud. Structure et evolution d'une orogene intracratonique. Geologische Rundschau, 69, 1-21. DALZIEL, I.W.D. 1993. Tectonic tracers and the origin of the proto-Andean margin. 12th Congreso Geologico Argentino, Actas, 3, Asociacion Geologica Argentina, Buenos Aires, 367-374. DALZIEL, I.W.D. 1997 Neoproterozoic-Paleozoic geography and tectonics: Review, hypothesis, environmental speculation. Geological Society of America Bulletin, 109, 16-42. DAVIS, J.S., ROESKE, S.M., MCCLELLAND, WC. & SNEE, L.W. 1999. Closing the ocean between the Precordillera terrane and Chilenia: Early Devonian ophiolite emplacement and deformation in the

southwest Precordillera. In: RAMOS, V.A. & KEPPIE, ID. (eds) Laurentia-Gondwana connections before Pangea. Geological Society of America Special Paper, 336, 115-138. ESCAYOLA, M.P., RAME, GA. & KRAEMER, PE. 1996. Caracterizacion y significado geotectonico de las fajas ultramaficas de Cordoba. 13th Congreso Geologico Argentino, Buenos Aires, 3, Asociacion Geologica Argentina, Buenos Aires, 421-438. FINNEY, S., GLEASON, J., GEHRELS, G, PERALTA, S. & ACENOLAZA, G. 2003. Early Gondwanan connection for the Argentine Precordillera terrane. Earth and Planetary Science Letters, 205,349-359. FINNEY, S., GLEASON, I, GEHRELS, G, PERALTA, S. & ACENOLAZA, G. 2004. Corrigendum to 'Early Gondwanan connection for the Argentine Precordillera terrane'. Earth and Planetary Science Letters, 219, 413. FORSYTHE, R. 1982. The late Paleozoic to early Mesozoic evolution of southern South America: A plate tectonic interpretation. Journal of the Geological Society, London, 139, 671-682. FORSYTHE, R., DAVIDSON, J., MPODOZIS, C. & JESINKEY, C. 1993. Lower Paleozoic relative motion of the Arequipa block and Gondwana: paleomagnetic evidence from Sierra de Almeida of northern Chile. Tectonics, 12, 219-236. FRAGOSO-CESAR, A.R.S. 1980. O craton do Rio de la Plata e o Cinturao Dom Felicianono Escudo Uruguayo-sul Riograndense. Congreso Brasileiro de Geologia, Anais, SBG, 31, Camboriu, 5, Sociedade Brasileira de Geofisica, Sao Paulo, 2879-2891. GANGUI,A. & GOETZE,H.J. 1996. The deep structures of the northern Argentina: constraints from 2D seismic data and 3D gravity modeling. 13th Congreso Geologico Argentino & 3rd Congreso de Exploracion de hidrocarburos, Buenos Aires, 2, Asociacion Geologica Argentina, Buenos Aires, 545-565. GONZALEZ, P.D., POIRE, D.G & VARELA, R. 2002. Hallazgo de trazas fosiles en la Formation El Jaguelito y su relation con la edad de las metasedimentitas, Macizo Norpatagonico Oriental, provincia de Rio Negro. Revista de la Asociacion Geologica Argentina, 57, 35-44. GONZALEZ BONORINO, F. & GONZALEZ BONORINO, G. 1991. Precordillera de Cuyo y Cordillera Frontal en el Paleozoico temprano: terrenes bajo sospecha de ser autoctonos. Revista Geologica de Chile, 18, 97-107. GOTZE, H. J. & KIRCHNER, A. 1997. Gravity field at the South American active margin (20° to 29° S). Journal of South American Earth Sciences, 2, 179-188. GRUNOW, A., HANSON, R. & WILSON, T. 1996. Were aspects of Pan-African deformation linked to lapetus opening? Geology, 24,1063-1066. HALLER, M. & RAMOS, V. 1984. Las ofiolitas famatinianas (Eopaleozoico) de las provincias de San Juan y Mendoza. 9th Congreso Geologico Argentino, S. C. Bariloche, 2, Asociacion Geologica Argentina, Buenos Aires, 66-83.

ACCRETION HISTORY OF SOUTHERN S. AMERICA HOFFMAN, RE 1991. Did the breakout of Laurentia turn Gondwanaland inside-out? Science, 252, 1409-1412. JEZEK, P., WILNER, A.P., ACENOLAZA, E & MILLER, H. 1985. The Puncoviscana trough - a large basin of Late Precambrian to Early Cambrian age on the Pacific edge of the Brasilian shield. Geologische Rundschau, 74, 573-584. KELLER, M., BUGGISCH, W. & LEHNERT, O. 1998. The stratigraphical record of the Argentine Precordillera and its plate-tectonic background. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 35-56. KELLOGG, K.S. & BECKMANN, G.EJ. 1983. Paleomagnetic investigations of Upper Proterozoic rocks in eastern Arabian Shield, Kingdom of Saudi Arabia. Bulletin of the Faculty of Earth Sciences, King Abdullaziz University, 6, 483-500. KEMPF, O., KELLERHALS, P., LOWRIE, W. & MATTER, A. 2000. Paleomagnetic directions in Late Precambrian glaciomarine sediments of the Mirbat Sandstone formation, Oman. Earth and Planetary Science Letters, 175, 181-190. KEPPIE, J.D. & BAHLBURG, H. 1999. Puncoviscana formation of northwestern and central Argentina: passive margin or foreland basin deposit? In: RAMOS, V. & KEPPIE, D. (eds) Laurentia-Gondwana connections before Pangea. Geological Society of America Special Paper, 336, 139-143. KEPPIE, J.D. & RAMOS, V.A. 1999. Odyssey of terranes in the lapetus and Rheic oceans during the Paleozoic. In: RAMOS, V. & KEPPIE, D. (eds) Laurentia-Gondwana connections before Pangea. Geological Society of America Special Paper, 336, 267-276. KOUKHARSKY, M. & MiRRE, J.C. 1974. Nuevas evidencias del vulcanismo ordovicico en la Puna. Revista de la Asociacion Geologica Argentina, 29, 128-134. KRAEMER, P.E., ESCAYOLA, M.P & MARTINO, R.D. 1995. Hipotesis sobre la evolucion tectonica neoproterozoica de las Sierras Pampeanas de Cordoba (30°40'-32°40'). Revista de la Asociacion Geologica Argentina, 50, 47-59. KRONER, A. & CORDANI, U. 2003. African and South American cratons were not part of the Rodinia supercontinent: evidence from field relationships and geochronology. Tectonophysics, 375, 325-352. LORK, A. & BAHLBURG, H. 1993. Precise U-Pb ages of monazites from the Faja Eruptiva de la Puna Oriental and the Cordillera Oriental, NW Argentina. 12th Congreso Geologico Argentino & II Congreso de Exploration de Hidrocarburos, Mendoza, 4, Asociacion Geologica Argentina, Buenos Aires, 1-6. LOTTES, A.L & ROWLEY, D.B. 1990. Reconstruction of the Laurasian and Gondwanan segments of Permian Pangaea. In: MCKERROW, W.S. & SCOTESE, C.R. (eds) Palaeozoic Palaeogeography and Biogeography. Geological Society, London, Memoirs, 12, 383-395. LUCASSEN, E, BECCHIO, R., WILKE, H., FRANZ, G.,

325

THIRLWALL, M., VIRAMONTE, J. & WEMMER, K. 2000. Proterozoic-Paleozoic development of the Central Andes (18-26 S) - a mobile belt of the South American craton. Journal of South American Earth Sciences, 13, 697-715. LUCASSEN E, BECCHIO R., HARMON R. ET AL. 2001. Composition and density model of the continental crust in an active continental margin - the Central Andes between 18° and 27°S. Tectonophysics, 341, 195-223. MAHLBURG KAY, S., RAMOS, V.A. & KAY, R. 1984. Elementos mayoritarios y trazas de las vulcanitas ordovicicas de la Precordillera Occidental: basaltos de rift oceanico temprano (?) proximos al margen continental. 9th Congreso Geologico Argentino, S.C. Bariloche, 2, Asociacion Geologica Argentina, Buenos Aires, 48-65. MAHLBURG KAY, S.M., ORRELL, S. & ABRUZZI, J.M. 1996. Zircon and whole rock Nd-Pb isotopic evidence for a Grenville age and Laurentian origin for the basement of the Precordilleran terrane in Argentina. Journal of Geology, 104, 637-648. MANNHEIM, R. 1993. Genesis de las volcanitas eopaleozoicas del sistema de Famatina. 12th Congreso Geologico Argentino, Mendoza, 4, Asociacion Geologica Argentina, Buenos Aires, 147-155. MARTINEZ, C. 1980. Geologic des Andes boliviennes. Travaux et documents de 1'Ostrom No. 119, Paris. MARTINO, R., KRAEMER, P., ESCAYOLA, M., GIAMBASTIANAI, M. & ARNOSIO, M. 1995. Transecta de las Sierras Pampeanas de Corodoba a los 32°S. Revista de la Asociacion Geologica Argentina, 50, 60-77. MCELHINNY, M.W. & MCFADDEN, PL. 2000. Paleomagnetism. Academic Press, New York, San Diego, California. MEERT, J.G. 2001. Growing Gondwana and rethinking Rodinia: a paleomagnetic perspective. Gondwana Research, 4, 279-288. MEERT, J.G. 2003. A synopsis of events related to the assembly of Eastern Gondwana. Tectonophysics, 362, 1-40. MEERT, J.G. & TORSVIK, T.H. 2003. The making and unmaking of a supercontinent: Rodinia revisited. Tectonophysics, 375, 261-288. MEERT, J.G. & VAN DER Voo, R. 1996. Paleomagnetic and 40Ar/39Ar study of the Sinyai dolerite, Kenya: implications for Gondwana assembly. Journal of Geology, 104, 131-142. MEERT, J.G., VAN DER Voo, R. & PAYNE, T.W. 1994. Paleomagnetism of the Catoctin volcanic province: a new Vendian-Cambrian apparent polar wander path for North America. Journal of Geophysical Research, 99, 4625-4641. MILLER, H., HOCKENREINER, M. & SOLLNER, F. 2003. Opening and closure of an island arc-back arc system in the Early Paleozoic at the Gondwana margin: Famatinian system and Las Termas belt, NW Argentina. 10th Congreso Geologico Chileno, Conception, Adas (CD), Universidad de Concepcion, Concepcion. MOREL, P. 1981. Paleomagnetism of a Pan-African diorite: a Late Precambrian pole for western

326

A. E. RAPALINI

Africa. Geophysical Journal of the Royal Astronomical Society, 65, 493-503. MOSHER, S. 1998. Tectonic evolution of the southern Laurentian Grenville erogenic belt. Bulletin of the Geological Society of America, 110, 1357-1375. MURTHY, G, GOWER, C., TUBRETT, M. & PATZOLD, R. 1992. Paleomagnetism of Eocambrian Long Range dykes and Double Mer formation from Labrador, Canada. Canadian Journal of Earth Sciences, 29, 1224-1234. OMARINI, R., GOTZE, H.J., SUREDA, R.J., SEILACHER, A. & PFLUGER, F. 1999. Puncoviscana folded belt in NW Argentina: testimony of Late Proterozoic Rodinia fragmentation and pre-Gondwana collisional episodes. Geologische Rundschau, 88, 76-97. PALMA, M.A., PARICA, P.D. & RAMOS, V.A. 1986. El granito Archibarca: su edad y significado tectonico, provincia de Catamarca. Revista de la Asociacion Geologica Argentina, 41, 414-419. PANKHURST, R.J. & RAPELA, C.A. (eds) 1998. The Proto-Andean margin of Gondwana. Geological Society, London, Special Publications, 142. PANKHURST, R.J., RAPELA, C.W., SAAVEDRA, I, BALDO, E., DAHLQUIST, J., PASCUA, I. & FANNING, CM. 1998. The Famatinian magmatic arc in the central Sierras Pampeanas: an Early to Mid-Ordovician continental arc on the Gondwana margin. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 343-367. PASSCHIER, C.W., TROUW, R.A.J., RIBEIRO, A. & PACIULLO, F.V.P. 2002. Tectonic evolution of the southern Kaoko belt, Namibia. Journal of African Earth Sciences, 35, 61-75. POWELL, C.McA. & PISAREVSKY, S.A. 2002. Late Neoproterozoic assembly of East Gondwana. Geology, 30, 3-6. POWELL, C.McA., Li, Z.X., MC£LHINNY, M.W., MEERT, J.G. & PARK, J.K. 1993. Paleomagnetic constraints on timing of the Neoproterozoic breakup of Rodinia and the Cambrian formation of Gondwana. Geology, 21, 889-892. PRAVE, A.R. 1996. Tale of three cratons: Tectonostratigraphic anatomy of the Damara orogen in northwestern Namibia and the assembly of Gondwana. Geology, 24, 1115-1118. RAMOS, V.A. 1984. Patagonia: Un continente paleozoico a la deriva? 9° Congreso Geologico Argentina, S.C. Bariloche, 2, Asociacion Geologica Argentina, Buenos Aires, 311-325. RAMOS, V. 1986. El diastrofismo Ocloyico: un ejemplo de tectonica de colision durante el Eopaleozoico en el Noroeste Argentine. Revista del Instituto Geologico y Minero, Jujuy, Argentina, 6, 13-28. RAMOS, V.A. 1988. Late Proterozoic-Early Paleozoic of South America - a collisional history. Episodes, 11, 168-174. RAMOS, V.A. 1995. Sudamerica: un mosaico de continentes y oceanos. Ciencia Hoy, Buenos Aires, 6, 24-29. RAMOS, V.A. 2000. The Southern Central Andes. In:

CORDANI, U.G, MlLANI, E J., THOMAZ FlLHO, A.

& CAMPOS, D.A. (eds) Tectonic Evolution of South America, 31st International Geological Congress, Rio de Janeiro, Brazil, 561-604. RAMOS, V.A. 2002. Evolution tectonica. In: HALLER, M.J. (ed.) Geologia y Recursos Naturales de Santa Cruz. Asociacion Geologica Argentina, Buenos Aires, Relatorio, 15th Congreso Geologico Argentine, 365-387. RAMOS, V.A. & BASEI, M. 1997. The basement of Chilenia: an exotic continental terrane to Gondwana during the Early Paleozoic. Symposium on Terrane dynamics '97, New Zealand, Christchurch, New Zealand, 140-143. RAMOS, V. & KEPPIE, D. (eds) 1999. LaurentiaGondwana connections before Pangea. Geological Society of America Special Paper, 336. RAMOS, V.A., JORDAN, T.E., ALLMENDINGER, R.W., MPODOZIS, C., KAY, S.M., CORTES, J.M. & PALMA, M. 1986. Paleozoic terranes of the Central Argentine-Chilean Andes. Tectonics, 5, 855-880. RAMOS, V.A., VUJOVICH, G, MAHLBURG KAY, S. & McDoNOUGH, M. 1993. La orogenesis de Grenville en las Sierras Pampeanas Occidentales: la Sierra de Pie de Palo y su integration al supercontinente proterozoico. 12th Congreso Geologico Argentino, Mendoza, 3, Asociacion Geologica Argentina, Buenos Aires, 343-357. RAMOS, V.A., DALLMEYER, R.D. & VUJOVICH, G. 1998. Time constraints on the Early Paleozoic docking of the Precordillera, central Argentina. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 143-158. RAMOS, V.A., ESCAYOLA, M., MUTTI, D.I., VUJOVICH, G.I. 2000. Proterozoic-Early Paleozoic ophiolites of the Andean basement of southern South America. In: DILEK, Y., MOORES, E.M., ELTHON, D., & NICOLAS, A. (eds) Ophiolites and Oceanic Crust: New insights from field studies and Ocean Drilling program. Geological Society of America Special Paper, 349, 331-349. RAPALINI, A.E. 1993. Evidencias de un evento de remagnetizacion regional en Precordillera asociado a la Fase San Rafael. 12th Congreso Geologico Argentino, Mendoza, 3, Asociacion Geologica Argentina, Buenos Aires, 415-422. RAPALINI, A.E. 1998. Syntectonic magnetization of the Mid-Paleozoic Sierra Grande Formation. Further constraints for the tectonic evolution of Patagonia. Journal of the Geological Society, London, 155, 105-114. RAPALINI, A.E. 2003. New Late Proterozoic paleomagnetic data from the Rio de la Plata craton: implications for Pannotia and Gondwana. 10th Congreso Geologico Chileno, Concepcion, Adas (CD), Universidad de Concepcion, Concepcion. RAPALINI, A.E. & ASTINI, R.A. 1998. Paleomagnetic confirmation of the Laurentian origin of the Argentine Precordillera. Earth and Planetary Science Letters, 155, 1-14. RAPALINI, A.E. & ASTINI, R.A. 2005. La Remagnetizacion Sanrafaelica de la Precordillera: Nuevas

ACCRETION HISTORY OF SOUTHERN S. AMERICA Evidencias. Revista de la Asociacion Geologica Argentina, 60, in press. RAPALINI, A. E. & CINGOLANI, C.A. 2004. First Late Ordovician paleomagnetic pole for the Cuyania (Precordillera) terrane of western Argentina: A microcontinent or a Laurentian plateau? Gondwana Research, 7, 1089-1104. RAPALINI, A.E. & TARLING, D.H. 1993. Multiple magnetizations in the Cambrian-Ordovician carbonate platform of the Argentine Precordillera and their tectonic implications. Tectonophysics, 227, 49-62. RAPALINI, A.E. & VILAS, IF. 1991. Preliminary paleomagnetic data from the Sierra Grande Formation: tectonic consequences of the first Middle Paleozoic paleopoles from Patagonia. Journal of South American Earth Sciences, 4, 25-41. RAPALINI, A.E., VILAS, J.F., BOBBIO, M.L. & VALENCIO, D.A. 1989. Geodynamic interpretations from paleomagnetic data of Late Paleozoic rocks in the Southern Andes. In: HILLHOUSE, J.W. (ed.) Deep structure and past kinematics of accreted terranes. American Geophysical Union, Geophysical Monograph Series, 50, 41-57. RAPALINI, A.E., TARLING, D.H., TURNER, P., FLINT, S. & VILAS, IF. 1994. Palaeomagnetism of the Carboniferous Tepuel Group, central Patagonia, Argentina. Tectonics, 13, 1277-1294. RAPALINI, A.E., ASTINI, R.A. & CONTI, CM. 1999. Paleomagnetic constraints on the tectonic evolution of Paleozoic suspect terranes from southern South America. In: RAMOS, V. & KEPPIE, D. (eds) Laurentia-Gondwana connections before Pangea, Geological Society of America Special Paper, 336, 171-182. RAPALINI, A.E., BORDONARO, O. & BERQUO,T.S. 2000. Paleomagnetic study of Cambrian-Ordovician rocks in the Eastern Precordillera of Argentina: some constraints on the Andean uplift of this block. Tectonophysics, 326,173-184. RAPALINI, A.E., VELASCO, M.S. & KOUKHARSKY, M. 2002. New paleomagnetic data from the Western Puna of Argentina: some tectonic speculations on its Early Paleozoic evolution. 5th International Symposium on Andean Geodynamics, Toulouse, France, Extended Abstracts, Institut de recherche pour le developpement, Paris, 505-508. RAPELA, C.W. & PANKHURST, RJ. 2002. Eventos tectonomagmaticos del Paleozoico Inferior en el margen proto-atlantico del sur de Sudamerica. 15th Congreso Geologico Argentino, El Calafate, 1, Asociacion Geologica Argentina, Buenos Aires, 24-29. RAPELA, C.W., PANKHURST, R.J., CASQUET, C, BALDO, E., SAAVEDRA, I, GALINDO, C. & FANNING, C.M. 19980. The Pampean orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de Cordoba. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 181-217. RAPELA, C.W., PANKHURST, R.J., CASQUET, C., BALDO, E., SAAVEDRA, I & GALINDO, C. 19986. Early

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evolution of the Proto-andean margin of South America. Geology, 26, 707-710. RAPELA, C.W, PANKHURST, R.I, FANNING, C.M. & GRECCO, L.E. 2003. Basement evolution of the Sierra de la Ventana fold belt: new evidence for Cambrian continental rifting along the southern margin of Gondwana. Journal of the Geological Society, London, 160, 613-628. SAAVEDRA, I, TOSELLI, A., Rossi, I, PELLITERO, E. & DURAND, F. 1998. The Early Paleozoic magmatic record of the Famatina System: a review. In: PANKHURST, R.I & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 283-295. SANCHEZ BETTUCCI, L. & RAPALINI, A.E. 2002. Paleomagnetism of the Sierra de las Animas complex, southern Uruguay: its implications in the assembly of Western Gondwana. Precambrian Research, 118, 243-265. SARADETH, S., SOFFEL, H.C., HORN, P., MULLERSOHNIUS, D. & SCHULT, A. 1989. Upper Proterozoic and Phanerozoic pole positions and potassium-argon (K-Ar) ages from the East Sahara craton. Geophysical Journal, 97, 209-221. SATO, A.M., TICKYJ, H., LLAMBIAS, E.S. & SATO, K. 2000. The Las Matras tonalitic-trondjhemitic pluton, central Argentina: Grenvillian-age constraints, geochemical characteristics and regional implications. Journal of South American Earth Sciences, 13, 587-610. SCHWARTZ, J.I & GROMET, L.P. 2004. Provenance of a late Proterozoic-early Cambrian basin, Sierras de Cordoba, Argentina. Precambrian Research, 129,1-21. SHACKLETON, R.M., RIES, A.C., COWARD, M.P. & COBBOLD, PR. 1979. Structure, metamorphism and geochronology of the Arequipa Massif of coastal Peru. Journal of the Geological Society of London, 136,195-214. THOMAS, W & ASTINI, R.A. 1996. The Argentine Precordillera: a traveler from the Ouachita embayment of North American Laurentia. Science, 273, 752-757. THOMAS, W. & ASTINI, R.A. 1999. Simple-shear conjugate rift margins of the Argentine Precordillera and the Ouachita embayment of Laurentia. Geological Society of America Bulletin, 111, 1069-1079. THOMAS, W. & ASTINI, R.A. 2003. Ordovician accretion of the Argentine Precordillera terrane to Gondwana: a review. Journal of South American Earth Sciences, 16, 67-79. THOMAS, W, ASTINI, R.A. & BAYONA, G. 2002. Ordovician collision of the Argentine Precordillera with Gondwana, independent of Laurentian Taconic orogeny. Tectonophysics, 345, 131-152. TORSVIK, T.H., CARTER, L.M., ASHWAL, L.D., BHUSHAN, S.K., PANDIT, M.K. & JAMTVEIT, B. 2001. Rodinia refined or obscured: paleomagnetism of the Malani igneous complex (NW India). Precambrian Research, 108, 319-333. TOSELLI, A., DURAND, F, Rossi DE TOSELLI, I & SAAVEDRA, I 1996. Esquema de evolution

328

A. E. RAPALINI

geotectonica y magmatica del eopaleozoica del sistema de Famatina y sectores de Sierras Pampeanas. 13th Congreso Geologico Argentina, Buenos Aires, 5, Asociacion Geologica Argentina, Buenos Aires, 443-462. TROMPETTE, R. 1997. Neoproterozoic (-600 Ma) aggregation of Western Gondwana: a tentative scenario. Precambrian Research, 82, 101-112. TRUCO, S. & RAPALINI, A.E. 1996. New evidence of a widespread Permian remagnetizing event in the central Andean zone of Argentina. Third International Symposium on Andean Geodynamics, St. Malo, France, Extended Abstracts, ORSTOM editions, Paris, 799-802. VACCARI, N.E., BENEDETTO, J.L. WAISFELD, B.G. & SANCHEZ, T. 1992. La fauna de Neseuretus en la Formation Suri (oeste de Argentina): edad y relaciones paleobiogeograficas. Revista Espahola de Paleontologia, 8, 185-190. VAN DER Voo, R. 1990. The reliability of paleomagnetic data. Tectonophysics, 184, 1-9. VILLAR, L. 1975. Las fajas y otras manifestaciones ultrabasicas en la Republica Argentina y su significado metalogenetico. 2nd Congreso IberoAmericano de Geologia Economica, Buenos Aires, 3, 135-155. VON GOSEN, W. 2002. Polyphase structural evolution in the northeastern segment of the North Patag-

onian Massif (southern Argentina). Journal of South American Earth Sciences, 15, 591-623. VUJOVICH, G & MAHLBURG KAY, S. 1998. A Laurentian ? Grenville-age oceanic arc/back-arc terrane in the Sierra de Pie de Palo, Western Sierras Pampeanas, Argentina. In: PANKHURST, R.J. & RAPELA, C.W (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142,159-179. WEIL, A.B., VAN DER Voo, R., MAC NIOCALL, C. & MEERT, J.G 1998. The Proterozoic supercontinent Rodinia: paleomagnetically derived reconstructions for 1100 to 800 Ma. Earth and Planetary Science Letters, 154, 13-24. WINDHAUSEN, A. 1931. Geologia Argentina. II parte. Geologia Historica y regional del territorio argentino. Editorial Peusser, Buenos Aires. WINTER, H. DE LA R. 1984. Tectonostratigraphy as applied to analysis of South African Phanerozoic basins. Transactions of the Geological Society of South Africa, 87, 169-179. ZIMMERMANN, U. & BAHLBURG, H. 1999. The evolution of the Ordovician southern Puna retro-arc basin (NW Argentina): provenance analysis and paleotectonic setting. 4th International Symposium of Andean Geodynamics, Goettingen (Germany), Abstracts, Institut de recherche pour le developpement, Paris, 830-831.

Geochronology of Proterozoic basement inliers in the Colombian Andes: tectonic history of remnants of a fragmented Grenville belt U. G. CORDANI1, A. CARDONA1, D. M. JIMENEZ1'2, D. LIU3 & A. P. NUTMAN4 ^Institute of Geoscience, USP, Rua do Lago 562, Cidade Universitdria, 05508-080 Sao Paulo, Brazil (e-mail: [email protected]) 2 INGEOMINAS, Diagonal 53, No. 3453, Bogota, Colombia 3 Chinese Academy of Geological Sciences, Beijing, China ^Australian National University, Canberra, ACT 0200, Australia Abstract: Basement inliers of high-grade metamorphic rocks within the eastern Colombian Andes record a Grenvillian history. Among them, the Garzon Complex and the Dibulla, Bucaramanga and Jojoncito gneisses were studied using different geochronological methods to produce better correlations in the context of the reconstruction of the Grenville belt and of the supercontinent of Rodinia. The dynamic evolution of all of these units includes a final collisional event with exhumation of high-grade rocks. Such a tectonic history bears strong similarities with the Grenville Province in Canada and seems to confirm that these domains took part in the aggregation of Rodinia. Mesoproterozoic U-Pb zircon ages indicate heritage from magmatic protoliths, and the Sm-Nd model ages, as well as the e^ values, suggest derivation from an evolved continental domain, such as the Amazonian craton, with some mixing with juvenile Neoproterozoic material. When these continental fragments are correlated with similar terrains in Mexico and the Central Andes, a large crustal fragment is implied; very probably it made up the southern portion of the Grenville belt within Rodinia, which was disrupted when Laurentia separated from Gondwana forming the lapetus Ocean, leaving behind cratonic fragments that were later accreted to the South American Platform.

A single supercontinent, Rodinia, was conceived by McMenamin & McMenamin (1990) to include all continental masses existing at the surface of the Earth at about 1000 Ma. In his reconstruction of Rodinia, Hoffman (1991) used correlation of mobile belts of the Grenvillian orogenic cycle to position the different cratonic fragments (Fig. 1). In this reconstruction and many others (i.e. Dalziel 1997; Weil et al 1998), the present-day western margin of the Amazonian craton is placed against the eastern margin of Laurentia, by correlating the Sunsas and Grenville orogenic belts (Fig. 1). The tectonic framework of the Grenville belt in Laurentia includes the presence of a series of accretionary orogens and a late thickened continental crust related to continental collision (McLelland et al 1996; Rivers 1997; Wasteneys et al 1999). However, its Amazonian counterpart, the Sunsas belt, lacks the accretionary record and the high-grade metamorphic units (Kroner & Cordani 2003). These contrasting features have been related by Sadowsky & Bettencourt (1996) to a transpressional orogen in which Amazonia collided against Laurentia obliquely, yielding transtensional features mainly characterized by intra-continental basins

Fig. 1. Hoffman's (1991) reconstruction of Rodinia. The Grenvillian belts (c. 1000 Ma) are enhanced and the Andean inliers are also included. SB, Sunsas belt; G, Grenville belt; c, Colombian Massifs; m, Mexican Massifs; a, Arequipa-Antofalla basement; cu, Cuyania terrane.

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 329-346. 0305-8719/$15.00 © The Geological Society of London 2005.

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Fig. 2. Reconstruction of the relative position of Laurentia and Gondwana at 550 Ma, modified from Cawood etal. (2001). A A, Arequipa-Antofalla; AC, Amazonian craton; AN, Arabian-Nubian shield; ANT, Antarctica; AU, Australia; AV, Avalonia; C-SF, Congo-San Francisco; C, Colombian terranes; Cu, Cuyania; IN, India; K, Kalahari; LA, Laurentia; MB, Mozambique belt; RP, Rio de La Plata; U-N, Uweinat-Nile; WA, Western Africa. and a poorly evolved oceanic basin in the former, and a high-grade collisional belt in the latter. Much later, in the latest Neoproterozoic (c. 570 Ma), Laurentia separated from what was then West Gondwana (Powell et al 1993). Break-out occurred along the structural weakness of the Grenville/Sunsas belt, giving birth to the Palaeozoic lapetus Ocean (Cawood et al. 2001). Figure 2 illustrates a possible relationship between Laurentia and Gondwana at 550 Ma, just after the opening of the lapetus Ocean. East and West Gondwana are shown together, after the closure of the Mozambique Ocean at about 550 Ma (Meert 2003). In this figure a few allochthonous terranes are shown, formed by disruption of the supercontinent, and bearing a Grenvillian basement. Later, in Palaeozoic times, they may have collided back to the western margin of Gondwana. The Cuyania terrane in the southern Andes is the largest and most widely studied of these, with both a basement of Grenvillian age and an early Palaeozoic cover that records the different phases of its tectonic evolution (Keller 1999, and references therein). Several other Proterozoic inliers have been

identified along the South American Andean Chain (see Wasteneys et al. 1995; Tosdal 1996; Restrepo-Pace etal 1997; Worner etal. 2000, for recent reviews). Due to their age, high-grade metamorphic character and tectonic position, they have been considered fragments of the disrupted Grenville belt. The geological and geochronological evolution of these units thus provides an important clue for possible correlations between the terranes with Grenvillian basement in western South America, and for understanding the formation and later closure of the lapetus Ocean, as well as formation and disintegration of Pangaea. Along the eastern Colombian Andes, some basement inliers of high-grade metamorphic rocks are described, whose broad geochronological constraints record a Grenvillian history (Kroonemberg 1982; Restrepo-Pace et al. 1997; Ordonez 2001). However, the details of their tectono-metamorphic episodes are not yet clear and their regional correlations with broader tectonic provinces are still poorly understood. In order to contribute to a better knowledge of their geological history, new geochronological data are presented on key samples from the Santa Marta, Guajira, Santander and Garzon

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Fig. 3. Terrane map of the Colombian Andes adapted from Toussaint (1993). Sampled Proterozoic inliers considered in this work: GC, Garzon Complex; BG, Bucaramanga Gneiss; DG, Dibulla Gneiss; JG, Jojoncito Gneiss.

inliers (Fig. 3). The results have been obtained through a variety of different dating methods (U-Pb SHRIMP on single zircon, Sm-Nd garnet-whole-rock isochrons and Ar/Ar step-

heating analyses) which, combined with the previously available data, enable confirmation and, to some extent, refinement of the regional correlations. They also enhance knowledge of

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the tectonic evolution of the studied rocks, providing means to a better understanding of their crustal evolution.

Radiometric methods U-Pb determinations from six samples, made on 90 single zircon crystals domains, were carried out on sensitive high-resolution ion microprobes (SHRIMP), either SHRIMP I at the Australian National University, or SHRIMP II of the Chinese Academy of Geological Sciences. Zircons were hand-picked and mounted in an epoxy resin for the isotopic measurements. Because of effects such as the differential yield of metal and oxide species between elements during sputtering, interelement ratios are calibrated with a standard, where the ratios are known by isotope dilution thermal ionization mass spectrometry (IDTIMS). Details of analytical procedures, including calibration methods, are presented fully by Williams (1998) and Stern (1998). 206 Pb-238U ratios have an error component (typically 1.5-2.0%) from calibration of the measurements using the standard zircons. U abundance and U-Pb ratios were calibrated against 238U using fragments of the single crystal SL13 zircon standard (572 Ma). All errors also take into account non-linear fluctuations in ioncounting rates beyond that expected from counting statistics (e.g. Stern 1998). Pooled dates calculated in this paper are weighted mean 207Pb/206Pb and 206Pb_238U dates? with errors at the two sigma level, rounded to the nearest million years, following correction for common Pb based on measured 204Pb. The analytical results are included in Table 1, and shown in the concordia diagrams of Figure 4. Sm-Nd whole rock and garnet analyses from 17 samples were made at the Center of Geochronological Research of the University of Sao Paulo (CPGeo-USP), according to the procedures described by Sato et al (1995). 143 Nd/144Nd isotopic ratios were obtained in a multi-collector mass spectrometer, with analytical precision of 0.0014% (2a). Experimental error for the 147Sm-144Nd ratios is of the order of 0.1%. La Jolla and BCR-1 standards yielded i43Nd/i44Nd = 0.511849 ± 0.000025 (la) and 0.512662 ± 0.000027 (la), respectively, during the period in which the analyses were performed. Due to the presence of 147Sm-144Nd fractionation patterns in some samples, related to the main metamorphic event, double-stage model ages were calculated after De Paolo et al. (1991). The Sm-Nd analytical data are presented in Table 2.

Ar/Ar laser step-heating analyses of micas (16) and amphiboles (8) where carried out following the standard procedures of the Ar/Ar laboratory of the CPGeo-USP (Vasconcelos et al. 2002). Several grains of each sample were irradiated in the nuclear reactor IEA-R1 of the Brazilian Institute of Nuclear Research (IPEN), together with Fish Canyon sanidine standards. Later, two to three of the irradiated grains for each of the samples were selected for the Ar/Ar analyses. During successive incremental heating steps, the released gas was purified in an ultravacuum system and the 40Ar/39Ar ratios were measured in a high-sensitivity MAP-215-50 mass-spectrometer. Table 3 gives a summary of the apparent ages as interpreted from the plateau obtained from each spectrum, and the calculated integrated ages. All figures of the Ar/Ar age spectra for the analysed samples, as well as the complete analytical data, are available online at http://www.geolsoc.org.uk/ SUP18226. A hardcopy can be obtained from the Society Library. For all radiometric data, the decay constants employed are from Steiger & Jager (1977). The U-Pb and Sm-Nd isochron ages were calculated using the programme Isoplot/Ex 2.1 of Ludwig (1999).

Regional geological setting The Colombian Andes are formed by three main mountain ranges that cross-cut the country in a NE-SW direction. Geologically they are formed by several tectonostratigraphic terranes, the most relevant of these are indicated in Figure 3. The westernmost is a composite oceanic-derived terrane whose Meso-Cenozoic accretionary history has been fundamental for the understanding of the Andean Orogeny (see Ramos & Aleman 2000 for a review). The eastern domain is formed by at least three terranes with continental character (Tahami, Chibcha and Andaqui), which include several tectonic inliers of Precambrian age (see Toussaint (1993) for a review). Dispersed fragments of Proterozoic highgrade metamorphic rocks characterize the basement of the Andaqui and Chibcha terranes. However, they show significant differences in their Palaeozoic evolution. Whereas undeformed Early Palaeozoic sedimentary rocks cover the Andaqui terrane, the Chibcha terrane is characterized by a strong Early Palaeozoic tectonomagmatic episode, followed by an extensive Late Palaeozoic sedimentation (Toussaint 1993). Both terranes exhibit Cambro-Ordovician faunas with Gondwana affinities (Velez & Villaroel 1993;

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Fig. 4. Concordia diagrams for U-Pb SHRIMP analyses of zircons: (a) Guapoton-Moncagua Gneiss, sample V-198; (b) Las Margaritas Gneiss, sample Gr-15; (c) Vergel Granulite, sample Gr-29; (d) Dibulla Gneiss, sample A-49; (e) Bucaramanga Gneiss, sample PCM-1105; (f) Jojoncito Gneiss, sample Jojon-1. (MSWD = Mean Squares Weighted Deviation.)

Restrepo-Pace et al 1997). On the other hand, their Meso-Cenozoic tectonic evolution is quite similar and includes the installation of a Jurassic continental magmatic arc and associated basins, followed by an extensive Cretaceous transgres-

sion, and a tectonic inversion in the Cenozoic, producing the Andean mountains, The Tahami terrane, whose basement records a Late Palaeozoic-Triassic tectonic evolution (Toussaint 1993; Vinasco et al 2001), is related

Table 1. U-Pb zircon SHRIMP data Labels/ Sample

U

Th

V-198, Guapoton-Moncagua Gneiss 1 38 22 2 338 89 3 174 61 4 99 38 5 227 90 6 387 171 7 80 34 8 319 64 9 197 45 10 495 209 11 202 68 12 411 140 Gr-29, Vergel Granulite 1 1066 185 2 1214 332 3 1312 290 4 877 72 5 1280 378 6 1611 159 7 1500 304 8 898 40 9 1256 248 10 869 40 11 564 136 12 634 160 13 745 66 14 1077 197 Gr-15, Las Margaritas Gneiss 1 431 208 2 655 321 3 637 269 4 470 486 5 897 312 6 429 254 7 484 563 8 415 360 9 264 159 10 460 219 11 377 216 12 447 335 A49, Dibulla Gneiss 1 324 91 2 551 165 3 1331 207 4 223 81 5 511 139 6 870 371 7 595 347

Th/U

204

Pb/206Pb

I38\jpa6\j

207Pb/206pb

207Pb/235u

207Pb/206Pb age (Ma ± la)

206Pb/238u age (Ma ± la)

% Disc

0.59 0.27 0.36 0.40 0.41 0.46 0.44 0.21 0.24 0.44 0.35 0.35

0.00098 ±44 0.00011 ±5 0.00027 ±8 0.00001 ± 9 0.00007±6 0.00001 ± 3 0.00058 + 21 0.00048 ±29 0.00018 ± 9 0.00008 ± 3 0.00006 ± 1 0.00014 + 5

5.520 ±130 5.920 ±110 4.889 + 90 5.390 ±110 5.047 ±90 5.166 ±90 5.870 ±130 5.317 + 100 5.880 ±110 5.679 ±100 5.000 ±110 5.802 ±100

0.0640 ± 83 0.0724 + 94 0.0749 + 97 0.0824 + 107 0.0769 ± 100 0.0794 ± 103 0.0732 + 95 0.0696 + 90 0.0706 + 92 0.0732 ± 95 0.0774 + 101 0.0719 ± 94

1.60 ± 176 1.687 ± 040 2.113 + 059 2.109 ± 067 2.100 ±048 2.120 ± 042 1.718 ± 086 1.810 ± 119 1.654 ±048 1.778 ± 036 2.134 + 051 1.709 ±036

1073 + 23 1007 + 18 1200 ± 20 1098 ± 21 1165 + 19 1141 ± 18 1014 ± 18 1111 ± 18 1012 ± 18 1045 ± 16 1175 ±23 1025 ± 16

743 ± 220 996 + 27 1067 + 42 1255 ±48 1118 + 30 1183 + 20 1019 ± 93 918 ± 130 946 ±44 1020 ± 19 1132 ±22 984 ± 26

44 1 12 12 4 4 1 21 7 2 4 4

0.18 0.28 0.23 0.08 0.30 0.10 0.21 0.05 0.20 0.05 0.25 0.26 0.09 0.19

0.00004 ±1 0.00002 + 1 0.00002 ± 1 0.00004 ±2 0.00008 + 1 0.00007 ±2 0.00029 + 5 0.00001 + 1 0.00004 ± 1 0.00006 ±2 0.00038 + 16 0.00002 ± 1 -0.00001 ± 0 0.00002 ± 1

5.656 + 73 6.430 ±83 5.561 ±72 6.697 ± 87 6.681 ±99 9.007 ± 129 8.005 + 150 4.399 ± 57 5.989 ±95 8.074 + 105 6.579 ±88 4.567 ±60 6.611 ± 86 5.393 + 150

0.0758 + 3 0.0731 + 6 0.0764 + 3 0.0725 ± 6 0.0870 + 6 0.0792 + 6 0.0883 ±43 0.0889 + 6 0.0758 + 4 0.0785 + 6 0.0785 + 23 0.0899 + 6 0.0737 ±4 0.0777 ± 5

1.849 ±26 1.567 ±22 1.895 ± 25 1.493 ± 21 1.796 ±29 1.213 ± 30 1.521 ±76 2.786 ± 36 1.744 + 30 1.341 ± 19 1.646 ± 53 2.714 ±38 1.538 ± 22 1.986 ± 34

1050 ± 12 932 + 11 1066 ± 13 897 ± 11 899 ±12 679 + 9 759 ± 9 1320 ± 15 995 ± 15 753 ± 9 912 ± 11 1276 ± 15 908 ±11 1097 ± 16

1091 + 9 1017 + 9 1106 ± 8 1000 ± 13 1361 ± 13 1178 ±41 1390 ± 93 1402 ± 7 1089 ± 11 1160 + 12 1160 ± 59 1423 ± 9,6 1035 ± 11 1139 ± 12

4 8 4 10 34 42 45 6 9 35 21 10 12 4

0.50 0.51 0.44 1.07 0.36 0.61 1.20 0.90 0.62 0.49 0.59 0.77

0.00007+ 0.00002 + 1 0.00001 ± 1 0.00002 ± 1 0.00001 ± 1 0.00000 + 1 0.00005 + 2 0.00002 + 1 -0.00003 ± 1 0.00003 + 2 0.00007 + 2 0.00005 ±2

6.137 ± 150 6.087 ± 148 6.085 ±079 6.071 ± 150 6.059 ±110 6.040 ±150 6.157 ±111 6.137 ± 150 6.088 ±82 5.978 ±79 6.102 ±81 5.988 ±81

0.0724 + 6 0.0727 ± 5 0.0735 + 6 0.0727 + 5 0.0734 ± 6 0.0738 ±5 0.0727 + 6 0.0724 + 5 0.0737 ± 6 0.0731 + 5 0.0726 ±6 0.0726 + 5

1.627 ± 42 1.646 ±41 1.665 + 23 1.651+25 1.670 ±32 1.685 ± 25 1.628 ±31 1.628 ± 24 1.669 + 27 1.687 + 25 1.640 ± 26 1.671 ±25

973 + 22 981 ± 22 981 + 12 983 + 12 985 ± 17 988 ± 12 970 ± 16 973 ± 12 980 ± 12 997 ± 12 978 ± 12 996 ± 12

0.29 0.31 0.16 0.37 0.28 0.44 0.60

0.00018 ± 4 0.00001 + 1 0.00010 + 3 -0.00005 + 3 <0.00001 0.00002+1 0.00002 + 1

4.491 ± 61 4.073±54 4.961±64 4.241±66 4.311±57 6.020 ±79 5.460 ±78

0.0880 ± 10 0.0881± 15 0.0814 + 5 0.0878 ±7 0.0874 ± 4 0.0727 ±4 0.0751 ± 6

2.700 + 48 2.984 + 66 2.263 ± 32 2.853 ± 51 2.795 + 39 1.664 ±23 1.896 ± 30

1296 ± 16 1415 ± 17 1184 ±14 1365 ± 19 1345 ± 16 991 ± 12 1081 ± 14

998 + 17 1005 ± 14 1028 + 11 1006 + 14 1024 ± 10 1036 ± 13 1006 ± 15 999 ± 15 1034 + 18 1017 + 14 1003 + 17 1002 ± 15 1381 ±22 1385 ± 33 1231 ± 12 1377 ± 16 1369 ± 9 1004 ± 11 1071 ± 15

2 2 5 2 4 5 4 3 5 2 2 1 6 -2 4 1 2 1 -1

Table 1. (continued) Labels/ Sample

U

Th

A49, Dibulla Gneiss continued 8 317 98 9 352 79 10 193 55 11 2071 263 12 1009 141 13 1183 150 14 1560 97 15 55 18 PCM-1105, Bucaramanga Gneiss 1 44 22 2 213 81 3 453 180 4 92 41 5 418 79 6 77 49 7 291 100 8 165 52 9 223 66 10 158 44 11 361 115 12 213 63 13 175 85 14 187 74 15 86 33 16 711 138 Jojon-1, Jojoncito Gneiss 1.1 226 45 2.1 1062 105 3.1 691 141 4.1 584 83 5.1 654 203 5.2 1184 214 6.1 1107 108 6.2 567 111 7.1 553 150 8.1 197 127 8.2 512 130 9.1 700 151 9.2 783 159 10.1 595 172 11.1 568 74 12.1 593 148 13.1 350 111 14.1 1099 109 15.1 542 121 16.1 318 94 17.1 575 134

Th/U

204

Pb/206Pb

238Tj/206u

207Pb/206Pb

207Pb/235u

207

Pb/206Pb age (Ma ± la)

206

Pb/238U age (Ma ± la)

% Disc

0.32 0.23 0.29 0.13 0.14 0.13 0.06 0.34

0.00007 ± 3 0.00008 ± 3 0.00044 ± 6 0.00003 + 4 0.00001 +1 0.00001 ±1 0.00004 ± 1 0.00053 + 21

5.085 ± 74 4.312 ± 61 4.171 ± 64 7.477 ± 95 5.891+ 76 5.436 ± 74 5.198 ±68 5.437 ± 91

0.0776 ± 9 0.0885 ± 9 0.0912 ± 12 0.0710 + 3 0.0805 + 3 0.0780 ± 3 0.0800 ± 3 0.0791 + 32

2.105 ± 39 2.831 + 49 3.014 ± 60 1.309 + 17 1.885 ± 24 1.978 + 27 2.121 ±29 2.007 + 88

1157 ± 16 1344 ± 17 1385 ± 19 809 + 10 1011 ± 12 1089 + 14 1134 ±14 1088 + 17

1138 ±24 1394 ± 20 1450 ±24 957 ± 8 1210 ±8 1147 ±7 1196 ±8 1176 ± 81

-2 4 4 15 16 5 5 9

0.52 0.39 0.41 0.46 0.20 0.66 0.36 0.32 0.31 0.29 0.33 0.30 0.50 0.41 0.40 0.20

0.00022 ±19 -0.00013 ±5 0.00005 ± 3 0.00012 ±7 -0.00004 ± 2 -0.00002+1 0.00003 + 2 -0.00002 ± 2 0.00011 ± 5 -0.00015 ± 7 0.00009 + 3 0.00001 ± 2 0.00013 ± 4 0.00006 + 4 0.00038 ± 12 <0.0001

5.008 ± 87 5.926 + 80 6.704 ± 89 5.089 ± 74 7.961 + 106 5.358 ± 80 3.774 ± 52 3.999 ± 55 5.381 ± 73 6.269 + 87 3.761+50 6.087 ± 83 6.812 ± 94 6.016 ± 82 6.747 +101 5.613 ± 73

0.0806 ± 31 0.0744 ±11 0.0711 + 6 0.0796 ± 14 0.0700 ± 6 0.0821 ± 11 0.0956 + 6 0.0915 ± 7 0.0766 ± 9 0.0752 ± 12 0.0965 ± 6 0.0751 + 7 0.0679 ± 13 0.0738 ± 9 0.0667 ± 20 0.0776 + 5

2.218 ± 93 1.732 + 34 1.462 + 23 2.156 ± 49 1.213 + 19 2.112 + 42 3.493 + 52 3.153 + 50 1.964 ±35 1.655 + 34 3.539 + 53 1.701 + 28 1.374 + 32 1.691 + 32 1.362 ± 46 1.907 + 26

1211 ± 76 1053 ± 29 960 ± 19 1186 ± 35 930 ± 17 1247 + 26 1540 ± 12 1456 ± 15 1112 ±24 1075 ± 33 1558 ± 12 1071 ± 20 865 ± 40 1035 ± 26 827 ± 63 1138 ±13

1173 + 19 1005 + 13 896 ± 11 1156 ±15 763 ± 10 1103 ±15 1515 ± 18 1439 ± 18 1099+14 954 ± 12 1520 ± 18 980 ±12 883 ± 11 991 ± 13 891 ± 13 1057 ± 13

3 5 7 3 18 12 2 1 1 11 2 8 -2 4 -8 7

0.20 0.10 0.20 0.14 0.31 0.18 0.10 0.20 0.27 0.65 0.25 0.22 0.20 0.29 0.13 0.25 0.32 0.10 0.22 0.30 0.23

0.00024 ± 8 0.00002 + 1 0.00001 ±3 0.00006 ± 3 0.00001 ±2 0.00001 ± 2 0.00003 ±1 0,00007 ± 3 0.00003 ± 5 0.00024 ±11 0.00001 ± 5 0.00002 ±3 <0.0001 0.00006 ± 3 <0.0001 0.00009 ± 3 0.00002 ±3 0.00003 ±3 0.00006 ± 3 0.00004 + 3 0.00007 + 4

3.741 ± 143 3.773± 163 4.063± 148 3.497+180 4.787+139 5.113± 154 4.730± 133 6.173+241 5.809± 187 4.374± 201 6.017+255 6.263± 307 6.058± 174 5.257+220 6.144± 185 5.988±228 3.714± 132 4.305± 167 5.245 + 184 3.781 ± 134 6.045± 198

0.09550 ± 21 0.08764 ±25 0.09268 ± 30 0.09119 ± 34 0.07885 ± 10 0.08326 + 15 0.08159 ± 7 0.07073 ± 15 0.07467 ± 16 0.08346 ± 19 0.06883 ± 25 0.07089 + 12 0.06913 + 7 0.07879 ± 7 0.06954 ± 5 0.06938 ± 10 0.09459 ± 20 0.08621 ± 11 0.07872 + 15 0.09510 +16 0.07318 ± 12

3.519 + 164 3.203 + 174 3.144 + 161 3.595 + 241 2.271+75 2.245 + 82 2.378 + 72 1.579 + 73 1.772 + 72 2.631 + 141 1.577 + 93 1.561 + 84 1.574 + 49 2.066 + 90 1.561 ±49 1.598 + 67 3.512 + 153 2.762 + 116 2.069 + 86 3.468 +142 1.669 + 64

1470 + 51 1471 + 57 1369 ± 45 1568 ± 72 1197 ±32 1122 + 31 1207 + 31 956 ± 35 1007 ± 30 1293 + 54 981+39 943 ± 43 974 ± 26 1098 ± 42 961 ± 27 984 ± 35 1480 ±47 1307 ± 46 1101 + 36 1456 + 46 972 ± 30

1538 ± 43 1375 ± 55 1481 ± 62 1451 + 73 1168 + 25 1275 ± 35 1236 ± 16 950 ± 43 1060 ±44 1280 ±46 894 ± 76 954 ± 35 903 ± 20 1167 ± 17 915 + 13 910 ±29 1520 ±41 1343 ±25 1165 ± 38 1530 ±32 1019 ± 35

1 10 4 12 5 7 0 4 4 11 0 9 4 6 9 1 0 6 1 4 2

All ratios corrected for 204Pb/206Pb with204Pb method.

Table 2. Nd analytical data Sample Guapoton-Moncagua Gneiss, V-198 Vergel Granulite, D-982 Vergel Granulite, V-332 Vergel Granulite, V-314 Vergel Granulite, Gr-29 Vergel Granulite, Gr-29b Vergel Granulite, Gr-29d Las Margaritas Gneiss C 32 Las Margaritas Gneiss, Gr-15p Las Margaritas G^neiss GT 15A. Dibulla Gneiss, A49 Dibulla Gneiss, A51 Dibulla Gneiss, A58 Anorthosite -Santa Marta-A5, All Jojoncito Gneiss, Jojon-1 Bucaramanga Gneiss, PCM-1105 Bucaramanga Gneiss, PCM-1178 (Gabro) * Garnet-whole-rock isochron. Double-stage, second event at 1000 Ma.

t

Material

Sm (ppm)

Nd (ppm)

147Sm_144Nd

143Nd/144Nd

Whole rock Whole rock Garnet Whole rock Garnet Whole rock Whole rock Whole rock Whole rock Whole rock Garnet Whole rock Garnet Whole rock Whole rock Whole rock Whole rock Whole rock Whole rock Whole rock Whole rock

6.324 0.994 12.526 2.540 17.679 17.843 2.714 5.362 5.095 10.547 15.338 8.149 2.423 2.985 19.94 10.747 3.689 0.194 11.493 6.996 10.145

34.019 7.866 3.955 11.060 11.579 67.973 13.302 17.726 20.418 54.574 17.112 26.339 1.256 18.195 84.283 41.708 25.546 1.31 57.651 38.524 32.879

0.1124 ± 4 0.0764 ± 4 1.9150 ± 69 0.1389 ± 5 0.9233 ± 32 0.1587 ± 6 0.1234 ± 4 0.1829 ± 6 0.1509 ± 5 0.1169 ± 4 0.5420 ± 21 0.1871 ± 7 1.1665 ± 48 0.0992 ± 3 0.1431 ± 5 0.1558 ± 5 0.0873 ± 3 0.0895 ± 3 0.1205 ± 5 0.1098 ± 4 0.1866 ± 7

0.511948 ± 1 0.511733 ± 2 0.522892 ± 7 0.512019 ± 1 0.516829 ± 1 0.512371 ± 1 0.511994 ± 1 0.512381 ± 1 0.512144 ± 6 0.511875 ± 1 0.514638 ± 1 0.512035 ± 1 0.518680 ± 2 0.511800 ± 2 0.512214 ± 1 0.51239 ± 1 0.511755 ± 2 0.511865 ± 1 0.511967 ± 1 0.511938 ± 2 0.512415 ± 1

£

Nd(0)

-12.6 11 f\

8

Nd(1140 Ma)

T(Ma)* TDm(Ma)t

-1.3 -0.4

1.79 1 7">

925

-16.1

-3.6

-10.8 -13.4 -14.3 -14.5 15 0

0.5 -1.9 -2.7 -2.8 33

935

1.96 1.65 1.84 1.90 1.91 1 94

000

-22.8

2.40

-10.1

1034 - 14 5 11 ^ -9.6 -12.9 10 7

-13.5 -12.4 -14.1

23 -0.4 1.2 -1.5 0.3 -2.6 -1.1 -2.6

1 87 1 7^>

1.59 1.81 1 67

1.85 1.78 1.89

GEOCHRONOLOGY OF COLOMBIAN BASEMENT

337

Table 3. Summary of obtained Ar/Ar data Sample Guapoton-Moncagua Orthogneiss (V-198)

Material

Biotite Biotite Biotite Hornblende Hornblende Biotite Vergel Granulite, charnockite (V-332) Biotite Biotite Biotite Vergel Granulite, enderbite (Gr-29) Biotite Biotite Biotite Vergel Granulite, mafic granulite (D-986) Biotite Biotite Biotite Vergel Granulite, biotite-garnet-gneiss (D-982) Biotite Biotite Vergel Granulite, mafic granulite (V-309) Biotite Biotite Hornblende Hornblende Hornblende Biotite Margaritas Gneiss, biotite gneiss (C-302) Biotite Biotite Biotite Margaritas Gneiss, enderbite (C-271) Biotite Biotite Margaritas Gneiss, biotite-pyroxene-hornblende gneiss (C-299) Biotite Biotite Hornblende Hornblende Biotite Margaritas Gneiss, mafic granulite (C-32) Biotite Hornblende Hornblende Biotite Margaritas Gneiss, biotite gneiss (Gr-15p) Biotite Biotite Biotite Dibulla Gneiss, biotite gneiss (A-49) Biotite Biotite Biotite Dibulla Gneiss, biotite-amphibolite (A-59) Biotite Hornblende Hornblende Biotite Bucaramanga Gneiss, hornblende-biotite gneiss (PCM-815) Biotite Biotite Hornblende Hornblende Hornblende Bucaramanga Gneiss, hornblende-biotite gneiss (PCM-1102) Biotite Biotite Biotite Hornblende Hornblende Hornblende Bucaramanga Gneiss, biotite gneiss (PCM-1105) Biotite Biotite Biotite N.D, Plateau not defined.

Plateau age (Ma)

Integrated age (Ma)

N.D N.D N.D N.D N.D 919.7 ± 0.8 N.D 907.4 ±1.0 914.3 ± 1.3 913.6 ± 1.1 N.D N.D N.D N.D N.D N.D 928.2 ± 1.2 976 ± 0.9 964 ±1.0 977.7 ± 1.6 981 ± 1.6 N.D. 1007.3 ± 0.9 956.1 ±1.1 1011.6 ±1.1 1044.7 ± 1.1 1017.8 ±1.2 N.D. 977.3 ± 1.9 981.1 ± 1.2 N.D N.D 954.6 ±1.3 948.3 ± 1.1 982.3 ± 1.6 N.D 964 ± 1.1 974.2 ±1.1 962 ±1.1 N.D N.D N.D N.D N.D 929.6 ± 1.5 N.D 208.6 ± 0.5 203.8 ± 0.7 202.0 ± 0.6 201.9 ± 0.5 198.2 ± 0.6 200.1 ± 0.6 193.6 ± 0.5 160.7 ± 1.0 154.5 ± 1.6 212.7 ± 0.5 N.D N.D 200.2 ± 0.4 199.1 ± 0.3 199.2 ± 0.4

479.0 ± 0.5 417.0 ± 0.4 298.8 ± 0.4 644.7 ± 0.7 670 ± 0.8 918.8 ± 0.8 919.2 ± 0.8 904.2 ± 0.8 914.3 ± 1.3 912.1 ± 1.1 937.1 ± 1.2 782.3 ± 0.8 863.6 ± 0.8 834 ± 0.8 857.4 ± 0.7 859.8 ± 0.8 930.4 ± 0.9 974.8 ± 0.9 963 ±1 1002 ±1.3 999.5 ± 1.5 1044 ±1.6 1006.6 ± 0.8 947.3 ± 0.9 1007.6 ± 0.9 1034.1 ± 0.9 1015.2 ± 1.2 1028.2 ±1.0 972.6 ± 1.1 980.1 ± 0.8 1076.5 ± 1.1 1067.5 ± 1.2 954.8 ± 0.8 945.2 ± 0.8 992.2 ±1.0 1001. 2 ±0.9 960.9 ±1.0 966.3 ± 0.9 957 ± 1.1 558.5 ± 0.8 717.6 ± 1.0 511.8 ±0.8 61 1.8 ±0.7 664.4 ± 0.6 928.2 ±1.3 802.3 ± 1.3 200.7 ± 0.4 193.0 ± 0.4 192.7 ± 0.7 196.2 ± 0.7 199.1 ± 0.6 204.0 ± 0.4 152.0 ± 2 138.9 ± 1.4 199.0 ± 0.5 206.4 ± 0.5 193.8 ± 0.4 181.8 ±1.2 201.0 ± 0.3 197.7 ± 0.3 199.0 ± 0.3

338

U. G. CORDANI ETAL.

to the formation of Pangaea and it is not discussed in this work.

Results and discussion In this chapter, a short description of the geological setting of each of the Proterozoic inliers is reported, followed by some comments on the significance of the geochronological data acquired in this work.

Garzon Complex The Garzon Complex (Fig. 3) in southeastern Colombia is by far the most extensive and continuous exposure of Proterozoic crust within the northern Andes. Following Kroonemberg (1982), a three-fold tectonostratigraphic subdivision for the complex has been assumed, from west to east: (1) Guapoton-Moncagua Gneiss, consisting of hornblende-biotite augen orthogneisses; (2) Vergel Granulites, comprising quartzfeldspathic gneisses, and charnockite to enderbite garnet granulites, with intercalations of mafic, ultramafic and calc-silicate rocks; (3) Las Margaritas Gneiss, a unit of predominantly metasedimentary character, composed of banded biotite-garnetsillimanite gneiss and related rocks, injected by granitic leucosomes formed by partial fusion, with local high temperature mylonitic zones. For the entire complex, the petrological evidence suggests prograde metamorphic conditions at about the amphibolite-granulite facies transition, with associated partial fusion related to biotite breakdown. A geothermobarometric study on rocks of the Las Margaritas Gneiss indicated a near-isothermal decompression (ITD) path, with pressures between 8 kbar and 4 kbar and temperatures between 800 °C and 630 °C. In contrast, the Vergel Granulites showed a counterclockwise path along 6-7 kbar and 700-733 °C. The previously available geochronological data from the Garzon Complex include a U-Pb zircon age of 1098 ± 9 Ma on the GuapotonMancagua Gneiss. In addition, there are some Rb-Sr whole-rock ages indicating a regional event around 1180 Ma and several K-Ar ages of 975-915 Ma (Priem et al 1989; Restrepo-Pace et al 1997, and references therein). The new U-Pb SHRIMP ages for the Garzon Complex are reported in Figures 4a, b, c. Figure 4a shows the data obtained on sample V-198 of the Guapoton-Mancagua orthogneiss.

In this rock, the zircons are elongated prisms, weakly rounded. Cathodo-luminescence (CL) images show core-border patterns, the former with oscillatory zoning of magmatic zircons and the latter clearly related to a metamorphic origin. The twelve analyses can be grouped in two main age populations: an older one with a 206Pb_238Tj age Of 1158 ± 23 Ma (Mean Squares Weighted Deviation (MSWD) - 0.79) defined by three cores, and a younger one of 1000 ± 25 Ma (MSWD - 0.76) on five metamorphic overgrowths. The older age is considered to be related to the crystallization of the granitic protolith and the younger one corresponds most probably to the high-grade metamorphic event responsible for the fabric of the rock. The protolith age overlaps within error with the conventional U-Pb ages of 1098 ± 9 Ma presented by Restrepo-Pace et al (1997). Figure 4b illustrates the U-Pb SHRIMP results obtained from sample Gr-15, a discordant leucosome from a migmatite of the Las Margaritas Gneiss. The zircons exhibit welldeveloped prismatic faces and all the analyses are nearly concordant, with a mean xnpbfxtopfo age of 1015 ± 8 Ma (MSWD - 0.76). This age is interpreted as the crystallization of the leucosome, formed by partial fusion in a magmatic event associated with decompression at the beginning of regional exhumation. Figure 4c shows the U-Pb SHRIMP results on zircons from the enderbite Gr-29 belonging to the Vergel Granulites. Two main zircon populations were identified, one with strongly rounded crystals and a second with elongated prisms. The zircon crystals have high U contents, and metamict zones are common. Most apparent ages are discordant and the points follow different trends of Pb loss. Two of the analyses suggest the presence of material older than 1300 Ma. A few of the prismatic zircons define approximately an upper intercept around 1100 Ma, whereas three rounded zircons seem to indicate a different intercept close to 1000 Ma. Although these data cannot be considered definitive, it is possible to interpret them in terms of a protolith formed at about 1100 Ma, by crustal derivation, and a metamorphic age around 1000 Ma. Four Sm-Nd whole-rock-garnet two-point isochrons on rocks from the Garzon Complex were obtained (see Table 1). Two high-grade gneisses from the Las Margaritas Gneiss yielded 1034 ± 6 Ma and 990 ± 8 Ma, representing cooling below 600 °C, after the main metamorphism, the younger age apparently recording a superimposed mylonitic overprint. A charnockite and a garnet gneiss of the Vergel Granulites

GEOCHRONOLOGY OF COLOMBIAN BASEMENT

yielded ages of 935 ± 5 Ma and 925 ± 7 Ma, also corresponding to cooling ages below 600 °C. A concordant Rb-Sr whole-rock isochron of 959 ± 28 Ma was obtained on a similar rock of the Vergel Granulites, confirming that a major metamorphic episode may have occurred at a temperature of about 500-600 °C, producing Sr isotopic homogenization of the mineral phases. The apparent ages around 1100 Ma for the high-grade metamorphic rocks of the Garzon Complex are synchronous with the igneous activity recorded in the Guapoton-Moncagua Gneiss, and are consistent with the interpretation of all three units belonging to the same orogenic belt (Kroonemberg 1982). Biotites and amphiboles from different units of the Garzon Complex show variable Ar/Ar ages (Table 2). Five biotites and one hornblende from the Vergel Granulites yielded plateau ages between 860 Ma and 980 Ma. In contrast, five biotites and two amphiboles from the Las Margaritas Gneiss show older ages of 950-1000 Ma, with one of the ages close to 1032 Ma. The different cooling ages obtained for the Vergel Granulites and the Las Margaritas Gneiss by both Sm-Nd garnet-whole-rock and Ar/Ar methods confirm a different thermal evolution for these two metamorphic units of the Garzon Complex. It is considered that the easternmost tectonic block, where the Las Margaritas Gneiss is found, was exhumed earlier in the history than the western block of the Vergel Granulite. The Guapoton-Moncagua Orthogneiss yielded very complex Ar/Ar spectra, with both amphibole and biotite showing irregular and staircase patterns (Table 2). They are due most probably to partial loss of Ar during younger Phanerozoic events related to Andean tectonics that, however, did not affect other units located towards the east - the Vergel Granulites and the Las Margaritas Gneiss.

Dibulla Gneiss In the Sierra Nevada de Santa Marta, northern Colombia (Fig. 3) a series of high-grade metamorphic rocks occur, including banded felsic and intermediate gneisses, pelitic paragneisses, hornblende-pyroxene gneisses, calc-silicate rocks, as well as metamorphosed anorthositic stocks and dykes (Tschanz et al 1974). These rocks are considered as derived from a riftrelated volcano-sedimentary sequence. The Dibulla Gneiss is part of the northeastern exposure of this complex (MacDonald & Hurley 1969) and consists of centimetric- to

339

decimetric-banded sequences of amphibolite and quartz-feldspathic rocks with associated migmatitic bands. Amphibole-plagioclase thermobarometry on amphibolites indicates minimum metamorphic conditions of 6.0-7.6 kbar and 760-810 °C, within the amphibolite-granulite facies transition. Previous geochronological data on the Dibulla Gneiss and other rocks from Santa Marta have shown metamorphic cooling ages around 970 Ma (Sm-Nd whole-rock-garnet isochron, and K-Ar mineral ages), as well as complex U-Pb zircon ages between 1000 Ma and 1500 Ma (Tschanz et al 1974; RestrepoPace et al 1997; Ordonez 2001). Rb-Sr wholerock data have yielded apparent ages around 1300-1400 Ma (MacDonald & Hurley 1969). A biotite gneiss (sample A-49) was analysed by the U-Pb SHRIMP method on zircon crystals that are elongated, with rounded edges. They are fragmented frequently and many exhibit metamict zones. The apparent ages are distributed between 1400 Ma and 980 Ma (Fig. 4d), with some of the analyses showing Pb loss. Five concordant zircons with a remarkably constant Th/U ratio (around 0.3), typical of magmatic zircons, yielded a strongly coherent 207pg/206p|-) age of 1374 ± 13 Ma (MSWD - 0.51). Two other nearly concordant zircons yielded a 207p|3/206p|:) age of 1145 ± 14 Ma, and two more concordant grains presented 207Pb/206U ages of 1081 ± 14 Ma and 991 ± 12 Ma. The discordant zircons show more intense metamictization and Pb loss. The c. 1370 Ma. age can be attributed to the magmatic crystallization of the zircons within a magmatic protolith. The zircon ages around 1140 ± 14 Ma seem to be geologically meaningful and can be related to a strong metamorphic event. The 991 ± 12 Ma age may be related to a younger metamorphic event, possibly of highgrade, considering the thermobarometric evidence given above. Ar/Ar ages from one amphibole and two biotites from rocks of the Dibulla Gneiss show variable behaviour (Table 2). Whereas the amphibole preserves older apparent 40Ar/39Ar ages around 930 Ma, which can be related to cooling after the last metamorphic event, biotite spectra do not define plateau ages, and the apparent ages of the incremental steps vary between 500 Ma and 700 Ma, suggesting partial Ar loss during younger Phanerozoic events.

Bucaramanga Gneiss The Bucaramanga Gneiss, within the Chibcha terrane (Fig. 3), is affected strongly by widespread Palaeozoic tectonomagmatic episodes.

340

U.G.CORDANIETAL.

Lithologically it corresponds to a metamorphic sequence of sillimanite, cordierite and garnet paragneisses, with intercalated amphibolites, marbles and calc-silicate rocks. Mineral parageneses suggest metamorphic conditions within the low-pressure-type amphibolite facies. Preliminary plagioclase-hornblende thermobarometry on amphibolites indicates pressures of 4-6 kbar and temperatures of 600-800 °C. Previous geochronological data are scanty. A K-Ar apparent age on biotite was about 920 Ma (Ward et al 1973), but Ar/Ar analyses on amphiboles have shown a more complex thermal evolution, pointing to an Early Jurassic partial resetting episode related to Mesozoic Andean tectonomagmatic activity (RestrepoPace et al 1997). Figure 4e shows the analytical data of sixteen U-Pb SHRIMP determinations carried out on zircon crystals from biotite gneiss. The zircons are prismatic, with rounded edges. Several points are nearly concordant and are distributed between 1550 Ma and 900 Ma. Three grains exhibited ages older than 1430 Ma. Other grains are scattered between 1200 Ma and 900 Ma and may indicate metamorphic resetting and/or mixed domains accidentally selected for analysis during the SHRIMP measurements. Three weakly discordant zircons can be grouped at a 207Pb/206Pb age of 1057 ± 28 Ma (MSWD - 0.61) whereas an isolated concordant zircon yielded an age of 1112 ± 24 Ma. The younger group presented a mean 207Pb/206Pb age of 864 ± 66 Ma (MSWD = 0.26), considered to be related to a late metamorphic episode. Three biotites and two amphiboles from three gneissic rocks clearly exhibit completely reset Ar/Ar spectra, with well-defined plateau ages in the 190-200 Ma time interval (Table 3). This is similar to a few U-Pb ages obtained on closely related granitic rocks from the same region (Dorr et al. 1995), indicating that the pervasive Jurassic magmatic episode had a significant influence on the thermal evolution of the basement units.

Jojoncito Gneiss The Jojoncito Gneiss (Alvarez 1967) represents the northernmost exposure of Proterozoic rocks in South America. It is a fine- to mediumgrained quartz-feldspathic rock exposed within the pre-Mesozoic basement of the Guajira Peninsula in the Caribbean region of Colombia. Stratigraphic relations are predominantly tectonic, but gradational contacts with metasedimentary rocks were observed. Previous geochronological dating include a Pb-a age of

1250 Ma (Banks 1975). Moreover, gneissic boulders from a Cenozoic conglomerate on Bonaire that can be correlated with rocks of the Guajira Peninsula have given U-Pb zircon ages of about 1160 Ma (Priem et al 1986). The zircon crystals separated from the Jojoncito Gneiss show variable external morphology, from prismatic, isometric to ovoid, with dimensions between 50 um and 150 um. Their internal structure in CL images is characterized by very clear core-border patterns. Most cores show oscillatory zoning typical of magmatic zircons, while borders are usually diffuse and homogeneous, typical of metamorphic overgrowth. Figure 4f illustrates the U-Pb SHRIMP results for the Jojoncito Gneiss. The analytical points of cores with oscillatory zones are mainly concordant and spread along the Concordia with 207pb/206pb ages between 1200 Ma and 1550 Ma. Three main groupings with similar ages and Th-U ratios can be suggested, with apparent ages of 1529 ± 43 Ma (MSWD - 0.05), 1342 ± 25 Ma and 1236 ± 16 Ma. The overgrowth domains have Th-U ratios of 0.1-0.3, typical of high-grade metamorphic zircon (Vavra et al 1999). Six of them define a 207Pb/206Pb age of 916 ± 19 Ma (MSWD - 0.48) and two others yielded individual ages of c. 1165 Ma. Such U-Pb determinations confirm a complex geological history for the Jojoncito Gneiss. The apparent ages are interpreted as related to a sedimentary parental rock that received detrital zircon from several different Mesoproterozoic sources. Moreover, two superimposed metamorphic events may have occurred at about 1165 Ma and 916 Ma. Therefore, the time of deposition of the sedimentary rock that was later transformed into the Jojoncito Gneiss is bracketed by the oldest metamorphism at 1165 Ma and by the youngest detrital zircon at 1260 Ma.

Neodymium isotopic constraints The sixteen samples of gneissic rocks from the studied Proterozoic inliers analysed by the Sm-Nd method (Table 2) are plotted in Figure 5. All Sm-Nd TDM model ages are approximately between 1.6 Ga and 2.4 Ga and, taking into account the zircon evidence demonstrating a strong Proterozoic heritage, such model ages may be considered to represent mean crustal residence ages (Arndt & Goldstein 1987). The Nd isotopic evolution lines shown in Figure 5 exhibit strong similarities. Negative to weakly positive eNd calculated at 1140 Ma (-3.6 to +1.2), as well as the overall Sm-Nd model age

GEOCHRONOLOGY OF COLOMBIAN BASEMENT

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Fig. 5. Nd isotope evolution diagram for basement rocks of the northern Andes, Amazonian craton and Mexican terranes. Envelopes include the Nd data presented by Restrepo-Pace et al. (1997), Lawlor et al (1999), Weber & Kohler (1999), Cordani et al. (2000) and Ordonez (2001).

pattern, suggest some mixture between Mesoproterozoic juvenile material and older crustal components for the source material of the highgrade metamorphic rocks. The existence of premetamorphic zircon populations confirmed the action of crustal reworking processes and, therefore, the presence of an ancient continental basement. These Nd isotopic constraints are similar to data already reported for the same units and for other Proterozoic exposures in Colombia (Restrepo-Pace et al. 1997; Ordonez 2001). Such isotopic coherence for the basement domains of the Chibcha and Andaqui terranes indicates that they have probably been subjected to a common crustal evolution.

Tectonic evolution The crustal derivation for their protoliths, their high-grade metamorphic character, and the

apparent synchronism of their metamorphic evolution, suggest that all the studied Colombian Proterozoic domains can be correlated regionally. Due to the numerous data from the Garzon Complex, a reasonable interpretation of its tectonic evolution can be attempted. A magmatic arc is envisaged at 1145 ± 14 Ma, established over an older continental margin. A first metamorphic event took place between 1100 Ma and 1040 Ma, followed by a younger one at about 1000 Ma. The latter was possibly due to a continental collision that could have produced large overthrusts, and may also have produced a doubling of crustal thickness, inducing the observed clockwise and anticlockwise paths for the metamorphism in the Margaritas Gneiss and Vergel Granulites. The former path could be linked directly to the collisional history, and the latter path may be related to the thickening of a previously heated continental

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Fig. 6. Distribution of concordant U-Pb zircon ages from Colombian Andes inliers and the Orinoco River sands (Goldstein et al. 1997).

crust during the formation of the magmatic arc. The Ar/Ar systematics described earlier confirm the differential exhumation rates of the two tectonic blocks. Within the northern units (Bucaramanga, Dibulla and Jojoncito gneisses) the tectonic environment may have been that of a continental rift within a previous continental mass. The maximum sedimentation age may be envisaged somewhere between 1200 Ma and 1300 Ma, and the metamorphic evolution would include a first event between 1140 and 1190 Ma, and a clearly defined later high-grade event with an age corresponding to the continental collision described for the Garzon Massif at about 1000 Ma. The presence of an active magmatic arc at 1145 ± 14 Ma in the Garzon Complex (the Guapoton-Moncagua Orthogneiss), very likely synchronous with the first metamorphic episode in the other tectonic domains, suggests that a Cordilleran-type tectonic setting was in existence at that time. Moreover, the younger metamorphic events at about 950-1000 Ma affecting all the studied terranes can be associated with a strong collisional episode involving the entire region. The orogenic activity would have ceased at 950-920 Ma, as suggested by several Ar/Ar cooling ages on minerals from rocks of the Garzon Complex and the Dibulla Gneiss.

Subsequent thermal events related to Mesozoic Andean magmatism have affected some of the inliers, as seen by the partial and/or total resetting of the amphibole and biotite Ar/Ar geochronometers. The importance of the described geochronological domains in the northern Andes is enhanced when detrital zircon data from the Orinoco River area (Goldstein et al 1997) and the inherited Grenvillian zircons from Palaeozoic magmatic rocks in the Merida Andes (Burkley 1976) are taken into account. In the Orinoco River work, 30% of 49 analysed zircons are within the 1.0-1.4 Ga time interval. Considering that the great majority of the sediments of the Orinoco come from the Andes (Johnson 1991), and that the other zircon ages can be found in the neighbouring cratonic provinces (see Tassinari et al 2000), the zircons with Grenvillian ages may well have been derived from the erosion of the exposed Proterozoic inliers of the Eastern Cordillera of Colombia. Figure 6 is a histogram that includes all the concordant (±5%) zircon ages determined in this work, together with those from Goldstein et al (1997) for the time interval between 800 and 1700 Ma. The matching of the two populations is quite remarkable. Of course, the Orinoco zircon population also includes younger grains derived from the Andean

GEOCHRONOLOGY OF COLOMBIAN BASEMENT region, as well as grains derived from the Archaean and Early Proterozoic provinces of the Amazonian craton. In addition, the histogram of Figure 6 emphasizes the statistical importance of the zircon population at about 1000 Ma, as well as other smaller peaks at 1100 Ma and 1200 Ma. Regarding provenance, the possibility of a terrane transfer between the Amazonian craton and Laurentia and/or other terranes during Late Mesoproterozoic continental collision has been commonly suggested (Dalziel 1997). In this respect, the older U-Pb zircon apparent ages from the Colombian Proterozoic rocks, roughly between 1300 Ma and 1600 Ma, can be traced to sources that can be found easily within the Rio Negro-Juruena, Rondonian and Sunsas provinces that occur at the southwestern corner of the Amazonian craton (Tassinari et al 2000). Arc magmatism within the Garzon Complex also compares well with the evolution of the Sunsas belt. In addition, the Nd systematics shown by the Colombian basement inliers (Fig. 5) are similar to those of the SW part of the Amazonian craton (Cordani et al 2000). According to Tosdal & Bettencourt (1994) and Ruiz et al (1999), the available data indicate that the Pb isotope composition is very similar for the basement rocks of southern Mexico, the Amazonian craton and the Colombian Andes, a situation also seen in the Nd isotope systematics (Fig. 5). Consequently, if one looks for largescale correlations, the Proterozoic rocks south of the Ouachita suture in Mexico can be correlated with the Andean inliers, as postulated by Ortega-Gutierrez et al (1995), Keppie & Ortega-Gutierrez (1999) and Ruiz et al (1999). The tectonic history of southern Mexico includes an older crust of 1400-1500 Ma, an arc-rift tectonic environment between 1280 Ma and 1110 Ma, a poorly constrained tectonometamorphic event at 1100 Ma and a granulite facies metamorphism between 990 Ma and 974 Ma (Keppie et al 2003; Solari et al 2003, and references therein). Such tectonic history is completely compatible with the Colombian domains. Moreover, within the South-Central Andes of South America a series of high-grade metamorphic rocks occurs, forming the basement of the Arequipa-Antofalla terrane in southern Peru, northern Chile and Bolivia. This domain has also yielded magmatic and metamorphic ages of 1.20-0.98 Ga (Wasteneys et al 1995; Tosdal 1996; Worner et al 2000). Further south, basement rocks from the Cuyania terrane also show a similar tectonomagmatic evolution (Kay et al 1996; Sato et al 2000). Orogenic events in the 1.2-1.0 Ga time

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interval have been related to continental collisional processes that eventually produced the Rodinia supercontinent (Hoffman 1991; Fig. 1). The Grenville Province of Laurentia is the most representative orogenic belt of this age and it is characterized by high-grade metamorphism and three successive tectono-metamorphic events (Rivers 1997): the older at about 1190-1140 Ma (Elzevirian pulse), an intermediate and more widespread one at 1080-1020 Ma (Ottawan), and a younger one at 1000-980 Ma (Rigolet). The Elzevirian episode has been related to back-arc closure and arc accretion, and the younger episodes have been related to continental collision (Rivers 1997; Wasteneys et al 1999). As demonstrated in this work, the ages and the metamorphic evolution found in the Colombian Proterozoic rocks (plus their correlative terranes in Mexico and in the Central Andes) and also the tectonic heritage from their protoliths, show strong similarities with the material and the evolution of the Grenville belt of Laurentia. The common geochronological evidence and the tectonic features encountered in all of these continental masses strongly suggest that they are fragments of the once coherent high-grade collisional belt that participated in the aggregation of Rodinia. For a palinspastic reconstruction, and following to some extent the ideas put forward by Hoffman (1991), when the tectonic basement inliers of Colombia are joined with the correlative Mexican and Central Andean domains and are placed near the northeastern part of Laurentia to form an extensive Grenville belt, a huge Meso-Neoproterozoic continental mass is formed (Fig. 1). Such a cratonic nucleus records a continental magmatic arc growth related to the consumption of a large Mesoproterozoic ocean between Laurentia, Baltica and the Amazonian craton (Starmer 1996; Kroner & Cordani 2003) and can be considered a prototype for the process of aggregation of the Rodinia supercontinent. After final formation of Rodinia around 980 Ma, and without discarding possible alternative reconstructions, this domain, including the whole of Laurentia plus a large part of what later would form West Gondwana, must have remained united. When Laurentia separated from Gondwana forming the lapetus Ocean at about 580 Ma (Cawood et al 2001; see Fig. 2), it is very likely that many fragments were left behind, but remained close to the western margin of Gondwana. The allochthonous terranes, such as the Colombian basement inliers, and possibly also Cuyania, ArequipaAntofalla and Chilenia, as suggested by

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different authors (Ramos & Aleman 2000), were mobilized later during Palaeozoic orogenic pulses, and accreted to the western margin of the South American Platform during the process of aggregation of Pangaea. The authors wish to thank the support received by the Brazilian Ministry of Science and Technology (PRONEX 41.96.0899.00) and the Sao Paulo State Foundation of Research Support (FAPESP 00/09695-1 and 01/08940-5). W. MacDonald (State University of New York, USA) and C. Garcia (Universidad Industrial de Santander, Colombia) kindly provided samples from the Jojoncito and Bucaramanga gneisses. Careful review and suggestions by S. R. Noble, W. McCourt, as well as the helpful editorial comments by R. J. Pankhurst, are gratefully acknowledged.

References ALVAREZ, W. 1967. Geology of the Simarua and Carpintero area, Guajira Peninsula, Colombia. PhD thesis, Princeton University, NJ. ARNDT, NT. & GOLDSTEIN, S.L. 1987. Use and abuse of crust formation ages. Geology, 15, 893-895. BANKS, P. 1975. Basement rocks bordering the Gulf of Mexico and the Caribbean Sea. In: NAIRN, A.E. & STEHLI, F.G. (eds) The Ocean basins and margins: 3 The Gulf of Mexico and the Caribbean. Plenum Press, New York, 191-199. BURKLEY, L.A. 1976. Geochronology of the Central Venezuelan Andes. PhD thesis, Case Western Reserve University. CAWOOD, PA., MCCAUSLAND, PJ.A. & DUNNING, G.R. 2001. Opening lapetus: constraints from the Laurentian margin in New Foundland. Geological Society of America Bulletin, 113, 443-453. CORDANI, U.G., SATO, K., TEIXEIRA, W., TASSINARI, C.C.G & BASEI, M.A.S. 2000. Crustal evolution of the South American Platform. In: CORDANI, U.G., MILANI, EJ.,THOMAZ-FILHO, A. & CAMPOS, D. A. (eds) Tectonic evolution of South America. 31st International Geological Congress, Rio de Janeiro 19-40. DALZIEL, I.W. 1997. Neoproterozoic-Paleozoic geography and tectonics: review, hypothesis, and environmental speculation. Geological Society of America Bulletin, 109, 16-42. DE PAOLO, D.J., LINN, A.M. & SCHUBERT, G. 1991. The continental crustal age distribution: Methods of determining mantle separation ages from Sm-Nd isotopic data and application to the southwestern U.S. Journal of Geophysical Research, 96, 2071-2088. DORR, W., GROSSER, J.R. & RODRIGUEZ, G.I. 1995. Zircon U-Pb age of the Paramo Rico tonalitegranodiorite, Santander Massif (Cordillera Oriental, Colombia) and its geotectonic significance. Journal of South American Earth Sciences, 8, 187-194. GOLDSTEIN, S.L., ARNDT, NT. & STALLARD, R.F. 1997. The history of a continent from U-Pb ages of

zircons from Orinoco River sand and Sm-Nd isotopes in Orinoco river sediments. Chemical Geology, 139, 269-284. HOFFMAN, PA. 1991. Did the birth of North America turn Gondwana inside out? Science, 252, 1409-1411. JOHNSON, MJ. 1991. Controls on the composition of fluvial sands from tropical weathering environment; sands of the Orinoco River drainage basin, Venezuela and Colombia. Geological Society of America Bulletin, 103, 1622-1647. KAY, S.M., ORRELL, S. & ABBRUZZI, J.M. 1996. Zircon and whole rock Nd-Pb isotopic evidence for a Grenville age and a Laurentian origin for the basement of the Precordillera in Argentina. Journal of Geology, 104, 637-648. KELLER, M. 1999. Argentina Precordillera: sedimentary and plate tectonic history of a Laurentian crustal fragment in South America. Geological Society of America Special Papers, 341. KEPPIE, J.D. & ORTEGA-GUTIERREZ, F. 1999. Middle American Precambrian basement: a missing piece of the reconstructed 1 Ga orogen. In: RAMOS, V.A. & KEPPIE, J.D. (eds) LaurentiaGondwana Connections before Pangea. Geological Society of America Special Papers, 336,199-210. KEPPIE, J.D., DOSTAL, J.D., SOLARI, L.A., ORTEGAGUTIERREZ, F. & LOPEZ, R. 2003. Geochronology and Geochemistry of Grenvillian igneous suites in northern Oaxacan Complex, Southern Mexico: tectonic implications. Precambrian Research, 112, 165-181. KROONEMBERG, G.S. 1982. A Grenvillian granulite belt in the Colombian Andes and its relations to the Guiana Shield. Geologie en Mijnbouw, 61, 325-333. KRONER, A. & CORDANI, U.G. 2003. African, southern Indian and South American cratons were not part of the Rodinia supercontinent: Evidence from field relationships and geochronology. Tectonophysics, 375, 325-352. LAWLOR, P.J., ORTEGA-GUTIERREZ, F., CAMERON, K.L.,

OCHOA-CARAMILLO, H., LOPEZ, R. &

SAMPSON, D.E. 1999. U-Pb geochronology, geochemistry and provenance of the Grenvillian Huiznopala Gneiss of Eastern Mexico. Precambrian Research, 94, 73-99. LUDWIG, K.R. 1999. Isoplot/Ex version 2.10. A Geochronological Toolkit for Microsoft Excel. Special Publication No. la. Berkeley Geochronology Center Berkeley, California, USA. MACDONALD, W.D. & HURLEY, P.M. 1969. Precambrian Gneisses from Northern Colombia, South America. Geological Society of American Bulletin, 80, 1867-1872. MCLELLAND, J., DALY, S. & MCLELLAND, J.M. 1996. The Grenville Orogenic cycle (1350-1000 Ma): an Adirondack perspective. Tectonophysics, 265, 1-28. MCMENAMIN, M.O. & MCMENAMIN, D.L.S. 1990. The Emergence of Animals: the Cambrian Breakthrough. Columbia University Press, New York. MEERT, J.G. 2003. A synopsis of events related to the

GEOCHRONOLOGY OF COLOMBIAN BASEMENT assembly of eastern Gondwana. Tectonophysics, 362, 1-40. ORDONEZ, O. 2001. Caracterizaqdo isotopica Rb-Sr e Sm-Nd dos principals eventos magmdticos nos Andes Colombianos. Dr Thesis, University of Brasilia, Brazil. ORTEGA-GUTIERREZ, F., Ruiz, J. & CENTENO-GARCIA, E. 1995. Oaxaquia - A Proterozoic microcontinent - accreted to North America during Late Paleozoic. Geology, 23, 1127-1130. POWELL, C.McA., Li, Z.X., MC£LHINNY, M.W., MEERT, J.G. & PARK, J.K. 1993. Paleomagnetic constraints on timing of the Neoproterozoic breakup of Rodinia and the Cambrian formation of Gondwana. Geology, 21, 889-892. PRIEM, H.N.A., BEETS, DJ. & VERDURMEN, E.A. TH. 1986. Precambrian rocks in an Early Tertiary conglomerate in Bonaire, Netherlands Antilles (southern Caribbean Borderland): evidence for a 300 Km eastward displacement relative to the South American mainland? Geologie en Mijnbouw, 65, 35-40. PRIEM, H.N.A., KROONEMBERG, S.B., BOELRIJK, N.A.I. & HEBEDA, E.H. 1989. Rb-Sr and K-Ar evidence of 1.6 Ga basement underlying the 1.2 Ga Garzon-Santa Marta granulitic belt in the Colombian Andes. Precambrian Research, 43, 315-324. RAMOS, V. & ALEMAN, A. 2000. Tectonic evolution of the Andes. In: CORDANI, U.G., MILANI, E.J., THOMAZ, A. & CAMPOS, D.A. (eds) Tectonic Evolution of South America. 31st International Geological Congress, Rio de Janeiro, 635-685. RESTREPO-PACE, PA., Ruiz, J., GEHLERS, G. & COSCA, M. 1997. Geochronology and Nd isotopic data of Grenville-age rocks in the Colombian Andes: new constraints for Late Proterozoic-Early Paleozoic paleocontinental reconstructions of the Americas. Earth and Planetary Science Letters, 150, 427-441. RIVERS, T. 1997. Lithotectonic elements of the Grenville Province: review and tectonic implications. Precambrian Research, 86, 117-154 Ruiz, X, TOSDAL, R.M., RESTREPO-PACE, PA. & MuRiLLO-MuNETON, G. 1999. Pb Isotope evidence for Colombian-southern Mexico connections in the Proterozoic. In: RAMOS, V.A. & KEPPIE, J.D. (eds) Laurentia and Gondwana Connections before Pangea. Geological Society of America Special Papers, 336, 183-198. SADOWSKY, G.R. & BETTENCOURT, J.S. 1996. Mesoproterozoic tectonic correlations between eastern Laurentia and the western border of the Amazon Craton. Precambrian Research, 76, 213-227. SATO, K., TASSINARI, C.C.G., KAWASHITA, K. & PETRONILHO, L. 1995. O Metodo Geocronologico Sm-Nd no IG/USP e suas Aplica9oes. Anais da Academia Brasileira de Ciencias, 67, 313-336. SATO, A.M., TICKYJ, H., LLAMBIAS, EJ. & SATO K. 2000. The Las Matras tonalitic-trondhjemitic pluton, central Argentina: Grenvillian-age constraints, geochemical characteristics, and regional implications. Journal of South American Earth Sciences, 13, 587-610. SOLARI, L.A., KEPPIE, J.D., ORTEGA-GUTIERREZ, F.,

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CAMERLON, H.L., LOPEZ, R. & HAMES, E. 2003. 990 Ma and 1100 Ma Grenvillian tectonothermal events in the northern Oaxacan Complex, Southern Mexico: roots of an orogen. Tectonophysics, 356, 257-282. STARMER, I.C. 1996. Accretion, rifting and collision in the North Atlantic supercontinent, 1700-950 Ma. In: BREWER, T.S. (ed.) Precambrian Crustal Evolution in the North Atlantic Region. Geological Society, London, Special Publications, 112, 219-248. STEIGER, R.H. & JAGER, E. 1977. Subcommission on geochronology: convention on the use of decay constants in geo- and cosmochronology. Earth and Planetary Science Letters, 36, 359-362. STERN, R.A. 1998. High resolution SIMS determination of radiogenic trace-isotope ratios in minerals. In: CABRI, LJ. & VAUGHAN, DJ. (eds) Modern Approaches to Ore and Environmental Mineralogy. Mineralogical Association of Canada Short Courses Series, 27, 241-268. TASSINARI, C.C.G., BETTENCOURT, J.S., GERALDES, M.C., MACAMBIRA, M. & LAFON, J.M. 2000. The Amazonian Craton. In: CORDANI, U.G., MILANI, E.J., THOMAZ, A. & CAMPOS, D.A. (eds) Tectonic Evolution of South America. 31st International Geological Congress, Rio de Janeiro, 41-96. TOSDAL, R. 1996. The Amazon-Laurentian connection as viewed from the Middle Proterozoic rocks in the Central Andes, western Bolivia and northern Chile. Tectonics, 15, 827-842. TOSDAL, R.M. & BETTENCOURT, J.S. 1994. U-Pb zircon and Pb isotopic compositions of Middle Proterozoic Rondonian Massifs, southwestern margin of the Amazon Craton, Brazil. 7° Congreso Geologico Chileno, Concepdon, Chile, Adas, 1538-1541. TOUSSAINT, J.F. 1993. Evolucion Geologica de Colombia. Precdmbrico-Paleozoico. Universidad Nacional de Colombia, Medellin, Colombia. TSCHANZ, C.M., MARVIN, R.F. & CRUZ, J. 1974. Geologic evolution of the Sierra Nevada de Santa Marta area, Colombia. Geological Society of America Bulletin, 85, 273-284. VAVRA, G., SCHMID, R. & GEBAUER, D. 1999. Internal morphology, habit and U-Th-Pb microanalysis of amphibolite-to-granulite facies zircons: geochronology of the Ivrea Zone (Southern Alps). Contributions to Mineralogy and Petrology, 134, 380-404. VASCONCELOS, P.M., ONOE, A.T., KAWASHITA, K., SCARES, AJ. & TEIXEIRA, W. 2002. 40Ar/39Ar geochronology at the Institute de Geociencias, USP: Instrumentation, analytical procedures, and calibration. Anais da Academia Brasileira de Ciencias, 74, 297-342. VELEZ, M.I. & VILLAROEL, C. 1993. La fauna de trilobites de la Formacion el Higado (Ordoviciano Medio), aflorante en la Serrania de las Minas (Huila-Colombia). VI Congreso Colombiano de Geologia. Memorias, 2, 169-180. Medellin, Colombia. VINASCO, C.J., CORDANI, U.G. & VASCONCELOS, P. 2001. 40Ar/39Ar dates in the Central Cordillera of

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Colombia: evidence for an Upper Triassic regional tectono-magmatic event. /// South American Symposium of Isotope Geology, Pucon, Chile. Actas (CD). WARD, D.E., GOLDSMITH, R. & CRUZ, J. 1973. Geologia de los cuadrdngulos H-12 Bucaramanga y H-13 Pamplona, Departamento de Santander. Boletin Geologico Ingeominas, 21. WASTENEYS, H.A., CLARK, A.H., FARRAR, E. & LANGRIDGE, RJ. 1995. Grenvillian granulitefacies metamorphism in the Arequipa Massif, Peru: a Laurentia-Gondwana link. Earth and Planetary Science Letters, 132, 63-73. WASTENEYS, H., MCLELLAND, J. & LUMBERS S. 1999. Precise zircon geochronology in the Adirondacks Lowlands and implications for revising platetectonic models of the Central Metasedimentary belt and Adirondack Mountains, Grenville Province, Ontario and New York. Canadian Journal of Earth Sciences, 36, 967-984. WEBER, B. & KOHLER, H. 1999. Sm-Nd, Rb-Sr and

U-Pb geochronology of a Grenville Terrane in Southern Mexico: origin and geologic history of the Guichicovi Complex. Precambrian Research, 96, 245-262. WEIL, A.B., VAN DER Voo, R., MAC NICOAILL, C. & MEERT, J.G. 1998. The Proterozoic supercontinent Rodinia: paleomagnetically derived reconstructions for 1100 to 800 Ma. Earth and Planetary Science Letters, 154,13-24. WILLIAMS, I.S. 1998. U-Th-Pb geochronology by ion microprobe. In: McKiBBEN, M.A., SHANKS III, W.C.P. & RIDLEY, WI. (eds) Applications of Micro analytical Techniques to Understanding Mineralizing Processes. Reviews in Economic Geology, 7, 1-35. WORNER, G, LEZAUN, I, BECK ETAL. 2000. Precambrian and early Paleozoic evolution of the Andean Basement at Belen (northern Chile) and Cerro Uyarani (western Bolivia Altiplano). Journal of South American Earth Sciences, 13, 717-737.

Archaeocyathan limestone blocks of likely Antarctic origin in G oncl wan an tillite from the Falkland Islands P. STONE1'2 & M. R. A. THOMSON3 British Geological Survey, Murchison House, West Mains Road, Edinburgh EH9 3LA, UK (e-mail: psto@bgs. ac. uk) ^Department of Mineral Resources, Ross Road, Stanley, Falkland Islands ^School of Earth Sciences, University of Leeds, Leeds LS2 9JT, UK (e-mail: m. thomson@stone-house. demon, co. uk)

l

Abstract: Cambrian limestone clasts containing a rich, well-preserved archaeocyathan fauna have been recovered from the late Carboniferous Fitzroy Tillite Formation of the Falkland Islands. Since neither Cambrian strata nor limestone are present anywhere in the indigenous rock succession, the clasts are regarded as exotic erratics introduced during the Permo-Carboniferous Gondwana-wide glaciation. Most recent reconstructions of Gondwana rotate the Falklands into proximity with the Eastern Cape, South Africa and the Ellsworth Mountains, Antarctica. In both of these areas, Permo-Carboniferous diamictites correlated with the Fitzroy Tillite Formation also contain rare, exotic clasts of archaeocyathan limestone. The Transantarctic Mountains seem the most likely source for all of these unusual erratics. This interpretation sustains the requirement for substantial rotation of the Falklands microplate into Gondwana reconstructions and illustrates the extent of the late Carboniferous ice sheet. Apparent differences in the tillite clast assemblages between East and West Falkland suggest variable provenance within the regional ice-flow regime.

Archaeocyaths are abundant, varied and wellpreserved in clasts of massive, grey and white limestone found in late Carboniferous tillite at Port Purvis and Hill Cove, West Falkland (Fig. 1); fragmentary archaeocyaths also occur in clasts of limestone breccia. In East Falkland, rare clasts of sparsely archaeocyathan limestone have been found in the tillite near Mount Pleasant Airport. Archaeocyaths are exclusively Cambrian (Rozanov & Debrenne 1974) but the Falkland Islands sedimentary sequence is entirely post-Cambrian, whilst the entire succession is devoid of limestone (Aldiss & Edwards 1999). Hence, the limestone clasts are likely to be far-travelled erratics introduced into the Falklands area during the late Carboniferous glaciation that affected wide areas of the Palaeozoic Gondwana supercontinent prior to its Mesozoic fragmentation. This paper provides an introduction to the archaeocyathan fauna but concentrates on the significance of its likely provenance in the context of Gondwanan glaciation and palaeogeography. In reconstructions of Gondwana the Falklands microplate is generally rotated through almost 180° and placed off the east coast of South Africa (e.g. Adie 1952; Marshall 1994; Storey et al. 1999; Trewin et al 2002) in proximity to East Antarctica and the Ellsworth

Mountains microplate. One such reconstruction is shown in Figure 2 in a polar projection after Powell & Li (1994). There is also a consensus that the sedimentary sequence preserved in the Falkland Islands is representative of those adjacent and once-continuous parts of the Gondwana supercontinent now fragmented into South America, Africa and Antarctica. Particularly close stratigraphical similarities are seen between the Falkland Islands and the Cape Fold Belt and Karoo Basin of Southern Africa (e.g. Aldiss & Edwards 1999; Trewin et al. 2002; Hunter & Lomas 2003). In the Falkland Islands, conjecturally Silurian to unequivocally Devonian, marine but nearshore clastic rocks of the West Falkland Group (Aldiss & Edwards 1999; Hunter & Lomas 2003) form the lower part of the sedimentary succession and rest unconformably on a Precambrian crystalline basement, the Cape Meredith Complex. The West Falkland Group is succeeded by the Carboniferous to Permian, mainly marine to lacustrine clastic strata of the Lafonia Group (Aldiss & Edwards 1999; Trewin et al. 2002). At the base of the Lafonia Group is a glacigenic unit, the Fitzroy Tillite Formation. Limestone is notably absent from the entire succession.

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 347-357. 0305-8719/$15.00 © The Geological Society of London 2005.

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Fig. 1. Outline geology of the Falkland Islands, showing the outcrop of the Fitzroy Tiilite Formation, after Aldiss & Edwards (1999).

The Fitzroy Tiilite Formation The Fitzroy Tiilite Formation (the Lafonian Tiilite in older literature) is the representative in the Falkland Islands of the late Carboniferous to early Permian glacigenic diamictite that is widespread across the fragments of Gondwana, recording a glacial episode about 290 Ma ago

(Veevers & Powell 1987). Gondwana began to break up about 200 Ma ago and continental fragments were dispersed around the Southern Hemisphere, each fragment with its own part of the once-contiguous tillite succession: the Dwyka Tillite in South Africa, the Sauce Grande Formation in Ventania, Argentina, and the Whiteout Conglomerate in the Ellsworth Mountains, Antarctica. The broad correlation between these different sequences is well established (e.g. Caputo & Crowell 1985; Matsch & Ojakangas 1992). The Fitzroy Tillite Formation of the Falkland Islands (Fig. 1) has been reviewed recently and defined formally by Aldiss & Edwards (1999). A comprehensive earlier study by Frakes & Crowell (1967) related lithological differences between East and West Falkland to different depositional environments, as follows. In the west, a brown, sandy mudstone matrix contains and supports a variety of exotic rock clasts, mostly as small pebbles but ranging up to

boulders 7 m across; this is regarded as a terrestrial, sub-glacial deposit. In the east, a dark grey and fine-grained muddy matrix contains a sparse assemblage of clasts that tend to be smaller and of a more restricted lithological range than is seen in the west; the East Falkland tillite was probably deposited in marine conditions under a floating ice sheet. Whatever the precise depositional process, the clasts in the Fitzroy Tillite Formation were derived from a wide range of rock types. They were carried into the Falklands area of Gondwana by ice, having been eroded from original sources that may have been some considerable distance away. The most common clast types are quartzite, sandstone, various granites, quartz and shale; previously reported accessory types include gneiss, dolerite, slate, porphyry, limestone, chert and conglomerate (Halle 1912; Baker 1924; Frakes & Crowell 1967; Aldiss & Edwards 1999). From recent work, this list can be expanded to include garnet-mica schist, ignimbrite, banded ironstone, and a much wider range of hypabyssal igneous types than was indicated by the previous designation of 'porphyry'. Not all clasts are necessarily exotic and some Skolithosbearing sandstone blocks could have been derived from the underlying Port Stephens Formation.

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Fig. 2. Reconstruction of Gondwana for the late Carboniferous, in polar projection after Powell & Li (1994). Ice-flow directions from Frakes & Crowell (1967) for the Falkland Islands, and from Crowell & Frakes (1972) for South Africa.

The limestone clasts contained in the tillite are relatively rare overall but, in West Falkand, appear to be concentrated at certain localities. In the Hill Cove section, for example, limestone clasts were found only at the eastern end; there they are relatively abundant with 22 of the 25 clasts examined in detail proving to contain archaeocyaths. Further, archaeocyathan limestone clasts appear to be rarer in East Falkland than in West Falkland. Although limestone is a widespread accessory in the East Falkland tillite only two limestone clasts have been found to contain archaeocyaths (out of 15 examined in detail). In the West Falkland clasts the archaeocyaths were commonly visible macroscopically, whereas in the East Falkland clasts the archaeocyaths were discovered only in thin section. Apart from the archaeocyaths described here, the authors know of only one other possible example of a limestone clast containing macrofossils: a cobble of crinoid-bearing limestone, found loose on West Falkland and held privately, probably came from the tillite although its reported location is not close to any known tillite outcrop (Richard Cockwell pers. comm. 2002). If both limestone types did indeed

come from the tillite they demonstrate further the lithological range of its clasts, since the archaeocyathan limestone must be Cambrian whereas crinoids did not appear widely until post-Cambrian times. In East Falkland the only additional, and rather tentative organic traces so far found are ooids in a clast recently collected from Mount Pleasant, East Falkland (BGS specimen LX1003); the ooids may have had an algal component to their formation (Davis et al 1978). Other qualitative differences are also apparent between the tillites of East and West Falkland, quite apart from the well-defined facies contrast: the lithological and size range of the clasts is much greater in the west than in the east; clasts in the west are generally more rounded than those in the east; in the west there is a much higher proportion of red granite and quartzite, relative to white varieties. Additionally, the apparent concentration of archaeocyathan limestone clasts at a few localities in West Falkland, rather than their being randomly distributed, is a feature that may have broad stratigraphical and palaeogeographical implications.

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Fig. 3. A wind-abraded limestone clast from the Fitzroy Tillite Formation at Port Purvis, showing transverse sections through archaeocyaths. The two-pence coin has a diameter of 2.5 cm. Department of Mineral Resources, Stanley, specimen number PS218.

The archaeocyaths Archaeocyaths are extinct organisms with a long history of phylogenetic uncertainty but a current consensus that they were related closely to the sponges (Hill 1972; Rowland 2001). They had a calcareous skeleton commonly consisting of two cones, one inside the other and connected by a variety of vertical and horizontal lamellar structures. They lived in a tropical, shallow water (< 100 m) environment. Fossil archaeocyaths typically range in size, in both diameter and length, from less than 1 mm up to about 5 cm, although lengths of over 50 cm have been recorded. The first examples appeared about 530 Ma ago during the Early Cambrian and archaeocyaths then diversified rapidly into hundreds of species that were important contributors to the construction of early marine reefs. Despite their success they were a relatively short-lived group and were extinct before the end of the Cambrian, less than 25 Ma after their first appearance. In contrast, other families of sponges that also appeared in the Cambrian are still represented today, over 500 Ma later. The most striking of the archaeocyathan limestone clasts was also the first to be found, at Gladstone Bay, an inlet at the western end of

Port Purvis (Fig. 1). Its discovery early in 2002 by Judith Clay and Sue Macaskill, whilst they were on holiday from the UK, was entirely fortuitous (Stone & Rushton 2003), but Clay and Macaskill recognized its appearance as highly unusual and arranged for it to be sent on to the Department of Mineral Resources in Stanley for identification. It was found loose but resting on outcrop of the Fitzroy Tillite Formation, and wind-abraded surfaces reveal the archaeocyaths in great detail (Fig. 3). This clast now forms part of the display collection of the Department of Mineral Resources, Stanley, Falkland Islands (Specimen number PS218). The original discovery provoked a systematic search by one of the authors (PS) with more than 20 archaeocyathan limestone clasts subsequently collected in situ from the tillite at Fox Point, to the east of Hill Cove, West Falkland, and two collected in situ from the tillite at Frying Pan Quarry, near Mount Pleasant Airport, East Falkland. These specimens (numbers prefixed PS303 and PS304) are currently held by the British Geological Survey for further study and, hopefully, formal identification. The Port Purvis clast (Fig. 3) is crammed with fragments and apparently more or less whole

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specimens of archaeocyaths in a finely crystalline limestone matrix. Cups are of solitary individuals, slenderly conical to cylindro-conical in form, and up to a little over 2.5 cm in diameter. Transverse sections are circular or suboval; all visible cups are double-walled, the intervallar spaces have radial septa and the central cavities appear to be empty. No lateral outgrowths are visible. The presence and arrangement of mural pores cannot be determined, but there are ragged fragments of test on the clast showing densely packed alternating rows of pores. At least three different species appear to be present and there are numerous small transverse sections with few septa that might be distinct species, or merely represent the early stages of the larger forms. A preliminary description and further illustrations of the Port Purvis archaeocyaths, based entirely on the naturally occurring sections, form Appendix A. In addition to the archaeocyaths, intricately curved shell fragments seen in cross-section are almost certainly derived from trilobites.

Clast provenance Archaeocyaths are restricted to Cambrian limestones (Rozanov & Debrenne 1974). There are no rocks of that age or that lithology within the Falkland Islands rock succession. The archaeocyathan limestone clasts in the Fitzroy Tillite are, therefore, exotic to the Falklands and were transported into the area by ice during the latest Carboniferous to earliest Permian glacial episode that produced the tillite deposits. Frakes & Crowell (1967) derived an overall ice flow direction of west to east (in modern coordinates) from lateral facies changes within the Fitzroy Tillite Formation. This becomes approximately south to north after rotating the Falklands microplate into a Permo-Carboniferous reconstruction of Gondwana (Fig. 2) where it is in line with ice-flow patterns deduced from South Africa (Crowell & Frakes 1972) and is indicative of ice flow from what is now East Antarctica. In East Antarctica, archaeocyathan limestone is present in situ at several localities in the Transantarctic Mountains (Laird & Bradshaw 1982; Debrenne & Kruse 1986; 1989; Buggisch & Henjes-Kunst 1999) and as loose material in recent moraines at Whichaway Nunataks (Hill 1965); illustrated species appear similar to those in the Falklands clasts. The specimens collected at Whichaway Nunataks (during the 1957 TransAntarctic Expedition) are now housed in The Natural History Museum, London and show a

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macroscopic similarity between their limestone lithology and that of some white limestone clasts from Hill Cove. In West Antarctica, limestone blocks containing archaeocyaths (albeit much smaller varieties than are seen in the Falklands example) have been recovered from the Whiteout Conglomerate of the Ellsworth Mountains (Debrenne 1992). That conglomerate is the local representative of the PermoCarboniferous, Gondwanan tillite and hence the time-stratigraphical equivalent of the Fitzroy Tillite Formation. Although archaeocyaths are also known from in situ Cambrian limestone in the Ellsworth Mountains (Buggisch & Webers 1992), they are somewhat unusual irregularians that persisted apparently into the Late Cambrian (Debrenne et al. 1984) and are very different species to those recovered from the Whiteout Conglomerate. Thus, the examples from the Whiteout Conglomerate most probably have an exotic source. A closer analogy to the Falkland Islands' situation is seen in the Dwyka Tillite, South Africa, where rare archaeocyathan limestone clasts have also been recorded and Antarctica proposed as the most likely provenance (Cooper & Oosthuizen 1974; Debrenne 1975; Oosthuizen 1981; Visser at al 1986) since archaeocyaths are otherwise unknown in the region. Curiously, one of the examples illustrated by Visser et al. (1986, fig. 6b) from the Dwyka Tillite has the broad central cavity and narrow rim of widely spaced septa, somewhat reminiscent of one of the examples from Port Purvis illustrated in Figure 3. So far as the authors are aware, there are no records of archaeocyaths from the Argentine representative of the Late Carboniferous Gondwana tillite, the Sauce Grande Formation of the Sierras Australes, nor any substantiated reports of in situ archaeocyaths from Cambrian strata in Argentina. The Transantarctic Mountains area (Fig. 2) would seem to be the most likely source of the archaeocyathan limestone clasts in the Fitzroy (Falkland Islands) and Dwyka (South Africa) tillites, and also of those in the stratigraphicallyequivalent Whiteout Conglomerate (Ellsworth Mountains, West Antarctica). Archaeocyathan limestone has been described from several different localities and stratigraphical levels (Debrenne & Kruse 1989; Wood et al. 1992) with the fauna from the Shackleton Limestone formally described by Hill (1964) and by Debrenne & Kruse (1989). The Shackleton Limestone host lithology is described by Hill as 'massive grey limestone' and among the taxa

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present is Thalamocyathus, which is also well represented in erratic faunas from Antarctica and South Africa and is possibly also present in the Port Purvis grey limestone clast (see Appendix A). The Transantarctic Mountains would also seem to be the most likely original source of the Whichaway Nunataks moraine samples and of the ice-rafted archaeocyathanlimestone clasts found at the Atlantic periphery of Antarctica: by dredging in the Weddell Sea (Gordon 1920) and on King George Island, South Shetlands, in Oligocene tillites (Troedson & Riding 2002; Troedson & Smellie 2002) and in recent moraines (Morycowa et al. 1982; Wrona & Zhuravlev 1996). Regional significance The discovery of archaeocyathan limestone clasts in the Fitzroy Tillite Formation, Falkland Islands, bears significantly on several aspects of the regional geology. Archaeocyath palaeobiology Archaeocyaths are not common in the Southern Hemisphere, with extensive faunas described only from Antarctica and Australia. Since the likely source of the Falkland Islands clasts is the Transantarctic Mountains, the fossils contained in them may expand the Antarctic archaeocyathan faunal record once the systematic description of the specimens from Hill Cove is complete. Extent of the Permo- Carboniferous ice sheet The Falklands clasts are found mostly in a terrestrial glacigenic lithofacies (Frakes & Crowell 1967). With a likely source in the Transantarctic Mountains (Fig. 2), this would indicate that the Late Carboniferous, Gondwana ice sheet flowed from the contemporary polar region to almost 60° south, at least, before becoming glacio-marine. Whilst a large ice sheet with these dimensions is accepted in most palaeogeographical proposals (e.g. Scotese et al. 1999), it is not supported by recent work in the Transantarctic Mountains themselves (Isbell et al. 2003). Further work on the Falklands clasts may help resolve this apparent contradiction. Glacial fades relationships Facies interpretations of the successions in the Ellsworth Mountains (Matsch & Ojakangas 1992, p. 59), suggest that those of the Whiteout

Conglomerate represent environments extending from the ice-sheet grounding line (Meyer Hills) into the open sea (Sentinel Range). Significantly, the glacio-marine facies is represented in the Falkland Islands by that part of the Fitzroy Tillite Formation that is furthest from Antarctica in the Figure 2 reconstruction of Gondwana, implying a terrestrial to glaciomarine transition across the Falklands microplate. In this respect the South African record is complicated by migration of multiple ice dispersal centres and consequential likelihood of clast re-working (Crowell & Frakes 1972; Caputo & Crowell 1985; Visser et al. 1986). One hitherto unexplored feature of the facies relationships between the Fitzroy Tillite Formation in East and West Falkland is the difference in the clast assemblages. This includes the apparently greater absolute number of archaeocyathan limestone clasts in the terrestrial facies of West Falkland relative to the glacio-marine facies of East Falkland, and the higher proportion of clasts carrying archaeocyaths within the total limestone population in West Falkland relative to East Falkland. If the terrestrial, late Carboniferous ice sheet did flow from the Transantarctic Mountains to the Falklands areas of Gondwana, it may then have merged with a glacio-marine ice sheet sourced in an entirely different area. Gondwana reconstructions In a reconstruction of Gondwana (Fig. 2), the coincidence of ice flow patterns between the Falkland Islands and South Africa (Adie 1952; Frakes & Crowell 1967) is now strengthened by the occurrence of the unusual, archaeocyathan limestone clasts in both areas. This adds further credibility to those break-up interpretations that require a substantial rotation of the Falklands microplate (e.g. Marshall 1994; Storey et al. 1999; Trewin et al. 2002), as incorporated in Figure 2. The apparent contrast between East and West Falkland in Fitzroy Tillite clast assemblages may also be relevant to discussion of the areas' structural assembly. Several authors (e.g. Marshall 1994; Thomas et al. 1997) have suggested that East and West Falkland were juxtaposed by major strike-slip movement along a large fault coincident with Falkland Sound. This would allow the contrasting glacigenic facies, with their different erratic clast assemblages, to be brought together tectonically (though Curtis & Hyam (1998) have argued from structural evidence against such large-scale movement), but would

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also preclude the use of their present relationships in any large-scale reconstructions. Thanks are due to the following for discussion and advice on various aspects of this paper and for reviews of its earlier manifestations: Don Aldiss, Peter Flood, John Isbell, Bob Thomas, Nigel Trewin, Adrian Rushton and an anonymous referee, clearly an archaeocyath specialist, whose comments were most helpful. In addition, Adrian Rushton kindly arranged access to the Whichaway Nunataks samples in The Natural History Museum, London. The authors thank Dra Susana Damborenea for advice on aspects of Argentine geology. PS publishes by permission of the Executive Director, British Geological Survey, NERC, and thanks Mrs P. Rendell, Director, Department of Mineral Resources, Stanley, for her support.

Appendix A: Descriptions of archaeocyaths from a limestone clast in the Fitzroy Tillite Formation, Port Purvis, Falkland Islands The following descriptions are based on examination of the hand-specimen only. The identifications suggested are preliminary and are based only on general similarities to illustrated specimens. Informally, three species are recognized.

Species 1 Transverse sections of species 1 are more or less circular (Fig. Ala, b) and up to 17.3 mm in diameter. They are characterized by a wide intervallar space with a proportionally narrow central cavity (about half of the cup diameter), and numerous thin radial septa in the intervallar space. The appearance of densely packed septa is strengthened because of the insertion of septa from the outer wall (Fig. Ale). Septal counts for a half circle on the best specimen (Fig. Alb) are approximately 35 on the inner wall and 48 on the outer wall, or 70 and 96 per full circle, respectively. From the sections available for study, it is unclear whether or not the septa are porous, although this would normally be expected to be the case. The specimen seen in longitudinal section (Fig. A2a) probably also belongs to species 1. The section lies slightly to one side of the median plane. It is approximately 80 mm long, with a maximum width of 19 mm. Septa are closely spaced and crossed by a series of numerous closely spaced tabulae, giving a gridlike pattern to the internal mould. The tabulae are more or less flat, although there are indications in places that some curve slightly against

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the internal wall and all curve more strongly against the external wall. The skeleton of the present archaeocyaths has been weathered preferentially compared to the matrix and a toothed structure to the infill between the tabulae (Fig. A2e) probably represents pore infills, although it is not possible to discern the arrangement of the pores. A close match with species 1 is yet wanting. However, because the tabulae appear to curve down to form the outer wall (Fig. A2e), it is possible that the species is a member of the Coscinocyathina.

Species 2 The second species is distinguished by having a relatively narrower intervallum and a correspondingly wider central cavity than species 1 (Figs Ala, and A2a, b, c). Available transverse sections are typically sub-oval, although this might possibly be because they are slightly oblique to the axis of the cup. In addition, the septa have the appearance of being more widely spaced than in species 1, but a septal count (ri) of 120 for the largest specimen with a maximum diameter (£>) of 22 mm actually gives almost the same septal index (nID = 5.5) as species 1 (nID = 5.6). Septal pores are visible in a few places and these seem to be aligned between adjacent septa. Oblique sections (Aid, e) of mediumsized specimens that probably belong to this species suggest that tabulae are not present, or at least are not developed to the same degree as in species 1. In places, on the individual in Fig. A2b, where alternate septa meet the internal wall, short hair-like spinose structures, protruding from the inner wall into the central cavity, are visible. In all cases these are deflected sharply in the same direction on either side of the central cavity to extend at a shallow angle to the wall; at the 'top' and 'bottom' of the transverse section there is a thin line of test concentric with the inner wall. Some of the 'spines' have short beak-like projections on the outer angle of the bend. On the largest specimen (Fig. Ala) hook-like 'spines' appear to emanate from a position corresponding approximately to the intersection of alternate septa with the inner wall, but on a narrowly ovate specimen (Fig. A2c, d) probably representing a more oblique section, they are proportionately shorter and more numerous, with one per septum and another one or even two between the 'septal' ones. At the 'top' of the oval section, several lines of concentric structure are visible. All of these 'spine'-like structures and accompanying lines of concentric structure, on

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Fig. Al. Archaeocyaths in a limestone clast, Fitzroy Tillite Formation, Port Purvis, West Falkland (a) General view of clast, showing sections through species 1 and 2 (X 2); (b) transverse section of species 1 (X 2 5)- (c) detail of (b), to show septal insertions (X 7); (d) longitudinal section of species 2; (X 2)- (e) longitudinal section through another specimen of species 2 (X 2); (f) oblique transverse section through species 3 (X 3V (g) oblique transverse section through another specimen of soecies 3 (x 3V

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Fig. A2. Archaeocyaths in a limestone clast, Fitzroy Tillite Formation, Port Purvis, West Falkland, (a) General view of another face of the clast, showing a longitudinal section through species 1 (right) and a transversesection through species 2 (top left) (X 2); (b) transverse section of species 2 (X 2.5); (c) oblique section through species 2, showing annulae (X 4); (d) detail of (c) to show annuli (X 8); (e) detail of longitudinal section of species 1, to show porous tabulae (X 2).

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the inner face of the inner wall, appear to correspond to a variety of sections through annulae on the inner wall. There are several specimens that seem to be almost longitudinal sections through smaller examples of species 2, the best being those in Figs Aid, e. The example in Fig. Aid is a section through a narrow cup with a curved axis. The lower part is close to a true transverse section and shows radial septa around a central cavity, the middle part passes along the intervallum and shows longitudinal lines marking the septa, whereas the top part is narrowly oblique to the median plane of the cup, clearly showing the arrangement of septa and annuli. The example in Fig. Ale, however, is more difficult to interpret. The lower part is much abraded but the upper part compares closely with Fig Aid and appears to be a section through a cup with a curved or undulating axis. The sharp demarcation between the lower and upper halves could suggest that the subject is either a branched cup or a narrow cup with an acute bend in the middle. At first sight, the largest specimen shown in Figure Ala and another enlarged in Figure Alb bear a superficial resemblance to Ladaecyathus fortiseptatus (Hill 1965, pi. V, fig. 2) from Whichaway Nunataks, Antarctica, but neither that species nor the genus are characterized by the annulae of the present specimens. A more likely affinity is with species of Thalamocyathus, described and illustrated widely in situ from the Lower Cambrian of South Australia and Shackleton Limestone of Antarctica (Hill 1964), from 'recent' erratic blocks from the Weddell Sea (Gordon 1920, pi. II, fig. 18) and Whichaway Nunataks (Hill 1965, pi. VII, especially figs 2, 3) and from erratic blocks in the Carboniferous Dwyka Tillite of South Africa (Debrenne 1975, fig. Ib, c).

Species 3 At least two oblique sections (Fig. Alf, g) through a possible third species appear on the reverse, more degraded side of the clast. Both are relatively small (10 mm or less) and characterized by an intervallar space broken up into narrowly ovate or tear-shaped sections. Narrowly ovate sections extend more or less radially from the inner to the outer wall, but where tear-shaped sections are present, they are two or three deep between the two walls. Both specimens are rather badly worn and it is difficult to be sure what the structure is. Two possibilities come to mind. One is that these represent undulose oblique cross-sections through species 1, i.e. a section that varies

between those seen in Figures Alb and A2a. The second is that these are sections through forms with taeniae, rather than septa, such as Flindersicyathus Bedford & Bedford, well illustrated from Whichaway Nunataks, Antarctica by Hill (1965, pi. XI, especially figs la, 2a, b, 7a, 8a) or even unrelated forms, such as Syringocnema (cf. Hill 1965, pi. XI, figs 17, 18).

References ADIE, RJ. 1952. The position of the Falkland Islands in a reconstruction of Gondwanaland. Geological Magazine, 89, 401-410. ALDISS, D.T. & EDWARDS, EJ. 1999. The Geology of the Falkland Islands. British Geological Survey Technical Report, WC/99/10. BAKER, HA. 1924. Final Report on Geological Investigations in the Falkland Islands, 1920-1922. Government Printer, Stanley. BUGGISCH, W. & HENJES-KUNST, F. 1999. Stratigraphy, facies and provenance analyses of the Lower Cambrian Mount Wegener Formation of the Shackleton Range, Antarctica. Terra Antartica, 6, 211-228. BUGGISCH, W. & WEBERS, GF. 1992. Facies of Cambrian carbonate rocks, Ellsworth Mountains, West Antarctica. In: WEBERS, GF, CRADDOCK, C. & SPLETTSTOESSER, J.F (eds) Geology and Paleontology of the Ellsworth Mountains, West Antarctica. Geological Society of America Memoir, 170, 81-89. CAPUTO, M.V. & CROWELL, 1C. 1985. Migration of glacial centers across Gondwana during Paleozoic Era. Geological Society of America Bulletin, 96, 1020-1036. COOPER, M.R. & OOSTHUIZEN, R. 1974. Archaeocyathid-bearing Erratics from Dwyka Subgroup (Permo-Carboniferous) of South Africa, and their importance to Continental Drift. Nature, 247, 396-398. CROWELL, J.C. & FRAKES, L.A. 1972. Late Palaeozoic Glaciation: Part V, Karroo Basin, South Africa. Geological Society of America Bulletin, 83, 2887-2912. CURTIS, M.L. & HYAM, D.M. 1998. Late Palaeozoic to Mesozoic structural evolution of the Falkland Islands: a displaced segment of the Cape Fold Belt. Journal of the Geological Society, London, 155, 115-129. DAVIS, ID., BUBELA, B. & FERGUSON, I 1978. The formation of ooids. Sedimentology, 25, 703-730. DEBRENNE, F. 1975. Archaeocyatha provenant de blocs erratiques des tillites de Dwyka (Afrique de Sud). Annals of the South African Museum, 67, 331-361. DEBRENNE, F. 1992. The archaeocyathan fauna from the Whiteout Conglomerate, Ellsworth Mountains, West Antarctica. In: WEBERS, GF, CRADDOCK, C. & SPLETTSTOESSER, IF. (eds) Geology and Paleontology of the Ellsworth Mountains, West Antarctica. Geological Society of America Memoir, 170, 279-284.

FALKLANDS ARCHAEOCYATHS DEBRENNE, F & KRUSE, P.D. 1986. Shackleton Limestone archaeocyaths. Alcheringa, 10, 235-278. DEBRENNE, F. & KRUSE, P.D. 1989. Cambrian Antarctic archaeocyaths. In: CRAME, J.A. (ed.) Origins and Evolution of the Antarctic Biota. Geological Society, London, Special Publications, 47, 15-28 DEBRENNE, F, ROZANOV, A.Yu. & WEBERS, G.F. 1984. Upper Cambrian Archaeocyatha from Antarctica. Geological Magazine, 121, 291-299. FRAKES, L.A. & CROWELL, J.C. 1967. Facies and paleogeography of Late Paleozoic Diamictite, Falkland Islands. Geological Society of America Bulletin, 78, 37-58. GORDON, W.T. 1920. Scottish National Antarctic Expedition, 1902-4: Cambrian organic remains from a dredging in the Weddell Sea. Transactions of the Royal Society of Edinburgh, 52, 681-714. HALLE, T.G 1912. On the geological structure and history of the Falkland Islands. Bulletin of the Geological Institution of the University of Uppsala, 11,115-229. HILL, D. 1964. Archaeocyatha from the Shackleton Limestone of the Ross System, Nimrod Glacier area, Antarctica. Transactions of the Royal Society of New Zealand, 2, 137-146. HILL, D. 1965. Archaeocyatha from Antarctica and a review of the phylum. Trans-Antarctic Expedition Scientific Reports, 10. HILL, D. 1972. Treatise on Invertebrate Paleontology, Part E (revised), 1 Archaeocyatha. C. Teichert (ed.), 2nd edition. The Geological Society of America and The University of Kansas, Boulder, Colorado and Laurence, Kansas. HUNTER, M.A. & LOMAS, S.A. 2003. Reconstructing the Siluro-Devonian coastline of Gondwana: insights from the sedimentology of the Port Stephens Formation, Falkland Islands. Journal of the Geological Society, London, 160, 459-476. ISBELL, J.L., LENAKER, PA., ASKIN, R.A., MILLER, M.F. & BABCOCK, L.E. 2003. Reevaluation of the timing and extent of late Paleozoic glaciation in Gondwana: Role of the Transantarctic Mountains. Geology, 31, 977-980. LAIRD, M.G. & BRADSHAW, J.D. 1982. Uppermost Proterozoic and Lower Palaeozoic Geology of the Transantarctic Mountains. In: CRADDOCK, C. (ed.) Antarctic Geoscience. The University of Wisconsin Press, Madison, 525-533. MARSHALL, J.E.A. 1994. The Falkland Islands: a key element in Gondwana palaeogeography. Tectonics, 13, 499-514. MATSCH, C.L. & OJAKANGAS, R.W. 1992. Stratigraphy and sedimentology of the Whiteout Conglomerate: an Upper Paleozoic glacigenic unit, Ellsworth Mountains, West Antarctica. In: WEBERS, G.F., CRADDOCK, C. & SPLETTSTOESSER, J.F. (eds) Geology and Paleontology of the Ellsworth Mountains, West Antarctica. Geological Society of America Memoir, 170, 37-62. MORYCOWA, E., RUBINOWSKI, Z. & TOKARSKI, A.K.

1982. Archaeocyathids from a moraine at Three Sisters Point, King George Island (South Shetland Islands, Antarctica). Studia Geologica Polonica, 74, 73-80.

357

OOSTHUIZEN, R.D.F 1981. An attempt to determine the provenance of the southern Dwyka from palaeontological evidence. Palaeontologica Africana, 24, 27-29. POWELL, C.McA. & Li, Z.X. 1994. Reconstruction of the Panthalassan margin of Gondwanaland. In: VEEVERS, J.J. & POWELL, C.McA. (eds) Permian-Trias sic Pangean basins and foldbelts along the Panthalassen margin of Gondwanaland. Geological Society of America Memoir, 184, 5-9. ROWLAND, S.M. 2001. The Archaeocyathan Enigma Solved? Journal of Paleontology, 75,1065-1078. ROZANOV, A.Yu. & DEBRENNE, F. 1974. Age of Archaeocyathid assemblages. American Journal of Science, 274, 833-848. SCOTESE, C.R., BOUCOT, AJ. & MCKERROW,W.S. 1999. Gondwanan palaeogeography and palaeoclimatology. Journal of African Earth Sciences, 28, 99-114. STONE, P. & RUSHTON, A.WA. 2003. Some new fossil records and notabilia from the Falkland Islands. Falkland Island Journal, 8 (2), 1-10. STOREY, B.C., CURTIS, M.L., FERRIS, J.K., HUNTER, M.A. & LIVERMORE, R.A. 1999. Reconstruction and break-out model for the Falkland Islands within Gondwana. Journal of African Earth Sciences, 29, 153-163. THOMAS, R.J., JACOBS, J. & WEBER, K. 1997. Geology of the Mesoproterozoic Cape Meredith Complex, West Falkland. In: RICCI, C.A. (ed.) The Antarctic Region: Geological Evolution and Processes. Terra Antartica Publication, 21-30. TREWIN, N.H., MACDONALD, D.I.M. & THOMAS, C.G.C. 2002. Stratigraphy and sedimentology of the Permian of the Falkland Islands: lithostratigraphic and palaeoenvironmental links with South Africa. Journal of the Geological Society, London, 159, 5-19. TROEDSON, A.L. & RIDING, IB. 2002. Upper Oligocene to lowermost Miocene strata of King George Island, South Shetland Islands, Antarctica: stratigraphy, facies analysis, and implications for the glacial history of the Antarctic Peninsula. Journal of Sedimentary Research, 72, 510-523. TROEDSON, A.L. & SMELLIE, J.L. 2002. The Polonez Cove Formation of King George Island, Antarctica: stratigraphy, facies and implications for mid-Cenozoic cryosphere development. Sedimentology, 49, 277-301. VEEVERS, J.J. & POWELL, C.McA. 1987. Late Palaeozoic glacial episodes in Gondwanaland reflected in transgressive-regressive depositional sequences in Euramerica. Geological Society of America Bulletin, 98, 475-487. VISSER, J.N.J., HALL, K.J. & LOOCK, J.C. 1986. The application of stone counts in the glacigene Permo-Carboniferous Dwyka Formation, South Africa. Sedimentary Geology, 46, 197-212. WOOD, R.A., EVANS, K.R. & ZHURAVLEV, A. Yu. 1992. A new post-early Cambrian archaeocyath from Antarctica. Geological Magazine, 129, 491-495. WRONA, R. & ZHURAVLEV, A.Yu. 1996. Early Cambrian archaeocyaths from glacial erratics of King George Island (South Shetland Islands), Antarctica. Palaeontologia Polonica, 55, 9-36.

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Lithospheric mantle domains beneath Antarctica P. T. LEAT1, A. A. DEAN1,1. L. MILLAR2, S. P. KELLEY3, A. P. M. VAUGHAN1 & T. R. RILEY1 1 British Antarctic Survey, High Cross, Madingley Road, Cambridge CBS GET, UK (e-mail: p. leat@bas. ac. uk) 2 British Antarctic Survey, c/o NIGL, British Geological Survey, Keyworth, Nottingham NG12 5GG, UK 3 Department of Earth Sciences, The Open University, Milton Keynes MK7 6A A, UK Abstract: The chemistry of mafic volcanic rocks and minor intrusions erupted on continents can be used to define sub-continental asthenospheric and lithospheric mantle sources. Data have been collated from Antarctica and the Falkland Islands (adjacent in Gondwana) in order to identify lithospheric mantle sources beneath the continent. The lithosphere-derived magmas include lamproitic and some lamprophyric rocks and endmembers in basaltic suites that are interpreted as mixtures of magmas from lithospheric and asthenospheric sources. The lithosphere-derived mafic rocks from Archaean to Middle Proterozoic cratonic and circumcratonic areas of East Antarctica have time-corrected eNd values of -20 to -3. This demands isolation of the LREE-enriched sources within pockets of stable sub-cratonic lithosphere for more than 1 Ga, consistent with the lithosphere thickness up to 250 km imaged by seismic tomography. In contrast, lithosphere-derived mafic rocks from Middle Proterozoic to Early Palaeozoic areas of West Antarctica, Victoria Land and the Falkland Islands that formed the Gondwana continental margin, have timecorrected 8Nd values of -3.6 to +3.5, implying more recent isolation from asthenosphere. In terms of mantle reservoirs, cratonic and circumcratonic areas trend toward EMI, with EMU possibly being a minor component. In contrast, Gondwana margin areas trend toward EMU, with EMI being, at most, a very minor component.

The Antarctic continent consists of a large number of cratons, orogenic belts and terranes. The crustal structure resulting from these amalgamation processes is known increasingly well (e.g. Stump 1995; Jacobs et al 1998; 2003; Pankhurst et al 1998; Fitzsimons 2000; Boger et al. 2001; Mikhalsky et al 2001). By contrast, the corresponding structure and compositional variations in the lithospheric mantle underlying the crust of Antarctica is known relatively poorly. In this paper, the geochemistry of mafic magmatic rocks is used to attempt to define lithospheric provinces in Antarctica and a discussion follows on how these relate to crustal age and structure.

Mafic magmas as lithospheric probes Mafic (broadly basaltic) magmas are derived by partial melting of mantle. Studies of mafic rocks around the world have shown clearly that many compositionally distinct mantle reservoirs have been tapped. An obvious feature is the difference between most of the lavas erupted in the oceans and many lavas erupted on the continents. The vast majority of basalts erupted at

oceanic spreading centres (mid-ocean ridge basalts - MORB) fall in a restricted compositional range and are interpreted as derived from the part of the upper convecting asthenospheric mantle that had been depleted (ie. lost a fraction of its original incompatible element content) during previous melt-generation events. Other oceanic basalts that occur characteristically on ocean islands (ocean island basalt - OIB) have less depleted and more varied compositions. They are also sourced dominantly in the asthenosphere, but probably contain components from the lower mantle. Such components are thought widely to have been convected upward by upwellings in the mantle, such as mantle plumes. Importantly, for the case developed in this paper, sub-oceanic lithospheric mantle can also be a source for oceanic basalts. Its composition is very close to that of the asthenosphere (from which it was derived directly) and its contribution to oceanic basalts can be inferred only in favourable circumstances (e.g. M. Storey et al 1988; Frey et al 2000). OIB-like (and minor MORB-like) basalts occur on continents as well as in the oceans, notably, but not exclusively, in places

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 359-380. 0305-8719/$15.00 © The Geological Society of London 2005.

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where continents are rifting. The asthenosphere is, therefore, a major source of continental basalts. However, many mafic rocks erupted through continents, including basalts, fall outside the compositional range of oceanic OIB and MORE (e.g. Saunders et al 1992; Hawkesworth et al. 1999). This compositional difference remains after the effects of contamination of the magmas by continental crust on their way to the surface have been quantified and eliminated. As the only remaining reservoir that occurs beneath the continents but not the oceans is sub-continental lithospheric mantle (SCLM), this is assumed to be a significant source of continental mafic magmas. It follows that mafic rocks, including basalts, that have been identified as coming from a lithospheric source can be used as probes of lithosphere composition. The composition of sub-continental lithosphere departs from that of the asthenosphere because lithospheric mantle is frozen to the continental base and does not convect. Influxes of incompatible elements transported by fluids or small-degree melts (Menzies & Hawkesworth 1987; McKenzie 1989) change the composition of continental lithosphere and, in time, these changes are reflected in differences in radiogenic isotopes compared to asthenospheric values. Small-volume mafic magmas, notably potassic rocks such as lamproites and calc-alkaline lamprophyres, which formed by small-degree partial melting of the parts of the continental lithospheric mantle that were enriched in incompatible elements, can have compositions that depart markedly from oceanic basalts (Bergman 1987; Foley et al. 1987; Thompson et al. 1989). These potassic magmas are, by consensus, the least ambiguous melts of sub-continental lithospheric mantle. Where possible, such rocks are used to define mantle sources in the continental lithospheric mantle. A significant problem in using mafic rocks as probes of lithospheric mantle compositions is that the lithosphere-derived melts may mix with asthenosphere-derived basaltic magmas during magma generation or during upward transport to the surface, thus obscuring the original lithosphere melt composition (e.g. Leat et al 1988; Thompson et al 1989; Wilson & Downes 1991). The concept of using compositions of lithosphere-derived mafic magmas as probes of lithospheric mantle structure and evolution has been applied in several continental regions. Leeman (1982) contoured Sr isotope compositions of Cenozoic basalts in the western USA to reveal lithospheric mantle structure and the

idea has been developed by numerous workers. Asthenosphere-derived basalts in the region are restricted to 'oceanic' domains (mostly regions of significant lithospheric extension), whereas lithosphere-derived basalts and potassic rocks are erupted in areas of low lithospheric extension underlain by older, thicker and more stable lithosphere (e.g. Menzies 1989; Thompson et al 1989; Gibson et al 1993). Menzies (1989) divided the western USA into cratonic, circumcratonic and oceanic domains based on the composition of basalts and ages and architecture of the lithosphere. Canning et al (1996) used the compositions of calc-alkaline lamprophyres (minettes) to show a terrane boundary in lithospheric mantle composition beneath the Scottish Caledonide Orogen and Wilson & Downes (1991) found regional differences in compositions of alkaline mafic Cenozoic and Quaternary rocks in mainland Europe that may be related to terranes in the Variscan Orogeny. Similar relationships have been demonstrated in northern China (Menzies et al 1993; Zhang et al 2004).

Antarctica and terranes Antarctica is divided tectonically into the geologically more stable East Antarctica and the more mobile West Antarctica. This division is based mainly on the contrasting tectonic behaviour of the two parts of the continent during and after break-up of Gondwana: East Antarctica behaved as a single continental fragment at this time, while West Antarctica behaved as several small crustal blocks separated by zones of intense (largely extensional) deformation (Dalziel & Elliot 1982; B. Storey et al 1988; 19990; Grunow et al 1991). East Antarctica consists of several Archaean cratons separated by crust formed during Early to Middle Proterozoic orogenies (Fig. 1; Harley et al 1998; Fitzsimons 2000). Late Proterozoic to Early Palaeozoic (c. 700-500 Ma) orogenies occur in two zones - along the Transantarctic Mountains and in a broad zone from the Shackleton Range to near Bunger Hills (Fig. 1) - and they probably record the time of stabilization of East Antarctica. The Ross Orogeny is well understood and is characterized by three major terranes, the Wilson, Bowers and Robertson Bay terranes. These terranes were amalgamated at the Gondwana-palaeo-pacific margin which trended at that time approximately along the line of the Transantarctic Mountains (Stump 1995). This amalgamation represents crustal growth and stabilization along that Gondwana margin. The broad zone of orogenesis in East

ANTARCTIC LITHOSPHERIC DOMAINS

361

Fig. 1. Tectonic sketch map of Antarctica, showing the location of mafic igneous dykes and lavas discussed in the paper. The map is compiled from the following sources: Black et al. (1991), Storey (1991), Sheraton et al. (1993), Stump (1995), Harley et al. (1998), Jacobs et al. (1998, 2003), Pankhurst et al. (1998), Fitzsimons (2000), Luttinen & Furnes (2000), Vaughan & Storey (2000), Boger et al. (2002), Hokada et al (2004). BT, Bowers terrane; CD, Central Domain; ED, Eastern Domain; EWM, Ellsworth-Whitmore Mountains; HN, Haag Nunataks; MBL, Marie Byrd Land; RBT, Robertson Bay terrane; WD, Western Domain; WT, Wilson terrane.

Antarctica is regarded as a continuation of the Pan-African Orogeny in eastern Africa and is mainly characterized by reworking of earlier Proterozoic mobile belts (Fig. 1; Fitzsimons 2000; Jacobs et al 1998; 2003). This orogeny is thought to represent the collision of East and West Gondwana - the final stage of amalgamation in the formation of the supercontinent (Fitzsimons 2000; Boger et al 2002; Jacobs et al 2003). If this interpretation is correct, the

East-West Gondwana suture runs through the zone of Pan-African Orogeny (Fig. 1) and may correlate with an ophiolitic complex in the Shackleton Range (Talarico et al 1999). The Falkland Islands crust was adjacent to Antarctica in Gondwana, within a zone of EarlyMiddle Proterozoic orogenies ('Grenvillian') that extend through South Africa and Dronning Maud Land (Fig. 2). Data from the Falkland Islands are, therefore, included in this paper.

362

P.T. LEATETAL.

Fig. 2. (a) Location of Cape Meredith in the Falkland Islands, (b) Sketch of the Early Jurassic reconstruction of Gondwana, showing the position of the Falkland Islands crustal block adjacent to Antarctica and southern Africa (after Groenewald et al. 1991; Storey et al. 1992; 19990; Jacobs et al 2003). The Falkland Islands block was situated within a c. 1.1 Ga orogen that included the Namaqua-Natal belt of southern Africa and the Maudheim Province of Dronning Maud Land. The basement of the Haag Nunataks and Ellsworth Whitmore Mountains crustal blocks probably also belonged to this province. The ages of the basements of the Filchner and Maurice Ewing blocks are poorly known - gneisses from DSDP Site 330 close to Maurice Ewing Bank yielded a Rb-Sr age of 535 Ma (Beckinsale et al 1976), possible reset during a Pan-Africa event. Abbreviations: GC, Grunehogna craton; Na-Na, Namaqua-Natal belt; MP, Maudheim province; MEB, Maurice Ewing block; FI, Falkland Islands block; FM, Filchner block; HN-EWM, Haag Nunataks-Ellsworth-Whitmore mountains blocks.

Lithospheric magmatism in Antarctica This paper is concerned only with the composition of mafic melts of sub-continental lithospheric mantle. The reasons for allocating

magma compositions to lithospheric sources are dealt with in each example below. The examples have been selected based on availability of Sr and Nd isotopic data and reliability of dating. This is because derived epsilon (e) Sr and Nd values are a function of analysed isotopic ratios and ages of emplacement. £Sr and ENd are a measure of the Rb-Sr and Sm-Nd ratios of the source mantles (and, hence, history of incompatible trace element enrichment or depletion) and of the times since the sources separated from the convecting asthenosphere. Some other examples clearly should be included in the dataset, such as the calc-alkaline lamprophyres of Victoria Land (Wu & Berg 1992), but the authors are unaware of published Sr and Nd isotope data for these rocks. Many of the lithosphere-derived mafic magmas discussed represent small-degree melts and are notable for their high abundances of elements such as K, Rb, Sr and LREE that are strongly incompatible in normal mantle minerals. These melts are, therefore, resistant to contamination by crust during the magmas' uprise, in the sense that very large degrees of contamination must take place before Sr and Nd isotopic values are altered significantly. Moreover, the magmas were small in volume, particularly in the case of the more alkaline types and, hence, carried very little heat for fusion of wall rocks. In most of the cases described, a significant role for such contamination by crust to explain the magmas' trace element and isotopic compositions was specifically ruled out in the original papers to which readers are referred. The possible role for crustal contamination of the magmas described is assessed further in the Results section. Analytical methods 40 Ar/39Ar analyses were carried out at the Open University, using methods detailed by Leat et al (2000). The J value calculated using the biotite standard GA1550 (98.8 Ma; Renne et al 1998) was 0.002940 ± 0.00001. Materials used were biotite (FI.215.1), amphibole (FI.219.1) and plagioclase (Z.736.4 and Z.914.4). 40Ar/39Ar analytical data are presented in Table 1. Sr and Nd isotopes were analysed at the NERC Isotope Geosciences Laboratory at Keyworth. The la errors are: 147Sm/144Nd, 0.3%; 143 Nd/144Nd, 0.003% and 87Sr/86Sr, 0.005%. i43Nd/i44Nd is normalized to 146Nd/144Nd 0.7219. During the period of analysis, the following standard results were obtained: La Jolla, 0.511860 ± 36 (2a); J&M, 0.511123 ± 32 (2o).

363

ANTARCTIC LITHOSPHERIC DOMAINS Table 1. Ar/Ar geochronological data for mafic dykes from the Falkland Islands and the Shackleton Range Step

40

Ar/39Ar

38

Ar/39Ar

Sample FI.215.1 1 127.51 2 115.46 3 123.27 4 280.70 118.52 5 6 117.89 7 117.93 8 104.28 9 123.78 121.67 10 11 140.42 12 115.55 114.34 13

0.032 0.020 0.028 0.127 0.030 0.019 0.026 0.033 0.029 0.026 0.011 0.026 0.021

Sample FL219.1 1 957.27 2 1713.58 3 1421.79 239.09 5 4 117.33 6 141.16 197.52 7 8 138.33 3408.52 9 242850.94 10 112759.14 11

0.454 0.825 0.806 0.178 0.085 0.075 0.151 0.116 2.352 191.976 69.125

Sample Z.736.4 1 2 3 6 4 5 7

37

Ar/39Ar

0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.04 0.03 0.03 0.01 0.04 0.01

36

Ar/39Ar

0.0357 0.0050 0.0198 0.5634 0.0034 0.0059 0.0035 0.0374 0.0378 0.0164 0.0792 0.0177 0.0056 Plateau age

66.21 62.56 50.27 125.88 77.52 90.02 138.52 442.39 503.34 39947.76 8522.64 Lowest

39

Ar %

9.0 13.9 21.6 32.1 40.8 50.1 58.0 62.4 74.2 81.5 82.5 93.2 100.0

0.3363 4.1 5.2 0.5080 0.4008 8.6 0.0082 15.5 0.1227 49.3 0.0173 73.8 0.0299 81.9 -0.0477 97.6 0.5807 99.7 29.7331 99.8 14.5538 100.0 age in the saddle

Age (Ma)

533.2 521.4 535.0 522.3 535.4 530.0 533.0 437.0 515.9 532.6 533.4 506.8 516.3 531.2*

±(Ma) 3.5 5.0 3.7 4.5 3.7 3.9 3.7 5.4 3.2 3.9 20.9 3.5 3.9 6.1

2271.5 3106.9 2841.9 952.8 385.6 607.0 796.1 668.0 4245.0 11790.6 10405.9 385.6

98.6 355.1 115.6 38.7 6.9 10.1 29.9 15.1 214.7 4738.0 2249.4 6.9

75.07 57.66 52.12 59.12 58.04 75.02 76.15

0.073 0.024 0.023 0.015 0.013 0.027 0.013

1.96 2.05 1.56 5.12 4.10 5.82 6.99

0.0479 0.0181 0.0006 0.0084 0.0000 0.0221 0.0068 Minimum age

9.7 26.5 55.9 80.4 62.5 69.0 100.0

297.2 258.1 256.4 277.9 284.2 330.9 355.8 256.4

5.8 3.8 2.2 5.0 8.5 8.5 3.3 2.2

Sample Z.914.4 1 121.05 2 124.22 96.21 3 98.67 5 4 116.73 107.49 6

0.045 0.108 0.059 0.031 0.052 0.021

2.57 9.21 9.11 6.05 9.97 11.13

0.0627 0.1349 0.0495 0.0177 0.0316 0.0109 Minimum age

19.9 29.8 40.1 69.8 48.2 100.0

475.3 399.7 387.9 437.9 495.0 482.4 387.9

11.0 21.6 21.9 9.8 26.2 7.6 21.9

*Ages included in the plateau are indicated in bold.

Elimination of asthenosphere-derived magmas Most of the documented Antarctic magmatism has been eliminated from consideration in this paper because of its asthenospheric or 'oceanic domain' characteristics. An asthenospheric source is determined when magmas have chemical evidence for derivation from depleted

MORB-source mantle (DMM) or an OIBsource mantle. This derivation is based mainly on Sr-Nd (and, to some extent, Pb) isotope compositions. Antarctic magmas with high components of MORB-source mantles normally have e Nd values > +7 and £Sr values < -14.2, and Antarctic magmas dominated by OIB-source mantle normally have eNd values in the range +3.2 to +9 and Esr values < -14.2. It

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P.T. LEAT ETAL.

should be emphasized that these values are approximate and guides only. Evidence from mantle xenoliths shows that there is a large overlap in Sr and Nd isotopic compositions between sub-continental lithospheric mantle and MORE and OIB reservoirs (Pearson & Nowell 2002). Nevertheless, many basaltic magmas in Antarctica can be assigned to asthenospheric sources with some confidence. The most clear-cut case probably is the widespread Cenozoic OIB magmatism in Antarctica that is dominated by the easily recognized HIMU OIB component, a global reservoir distinct in its low ENd and low £Sr values (and high 206Pb/204Pb and 207Pb/204Pb ratios). Using these parameters, most Cenozoic volcanism in Victoria Land (Rocholl et al 1995) and all Cenozoic volcanism in Marie Byrd Land, Thurston Island and the Antarctic Peninsula (Hole etal 1993; Hole & LeMasurier 1994; Hart et al. 1995; 1997; Panter et al 2000) can be assigned to DMM or OIB-source mantle. Note that some authors have suggested that the HIMU source is located in the lithosphere when melting took place, but make it clear that, if this is the case, this HIMU reservoir was recently (i.e. during Mesozoic to Cenozoic times) attached to the lithosphere and is not a structural part of the long-lived lithosphere (Rocholl et al 1995; Hart et al 1997; Panter et al 2000). Ultramafic lamprophyres in the Pensacola Mountains show that the HIMU mantle source was present beneath Antarctica by 183 Ma (Leat et al 2000; Riley et al 2003) and may have arrived at that time in mantle plumes responsible for the Karoo-Dronning Maud LandFerrar large igneous province and the early stages of the break-up of Gondwana (Riley et al 2003; 2005). Other mafic rocks that can be assigned to an asthenospheric source on the basis of OIB-like Sr and Nd isotope compositions or OIB-like incompatible trace element compositions, or both, include most ultramafic lamprophyres from Beaver Lake, Prince Charles Mountains (Andronikov & Foley 2001), Schirmacher Oasis (Hoch 1999) as well as the Pensacola Mountains (Leat et al 2000; Riley et al 2003), alkaline lamprophyres, such as those from the Vestfold Hills (Delor & Rock 1991) and Jetty Peninsula, Prince Charles Mountains (Mikhalsky & Sheraton 1993) and many Palaeozoic alkaline and tholeiitic basaltic dykes. Data for basalts from the Jurassic KarooDronning Maud Land-Ferrar large igneous province have not been included. Geochemical studies show that some of the low-Ti basalt types in this province probably were derived

from, or contained components of, lithospheric mantle sources (Luttinen & Furnes 2000; Hergt & Brauns 2001; Riley et al 2005). However, there is strong evidence that many of the magmas in this province, especially the low-Ti ones most likely to have been derived from the lithosphere, were emplaced by a process of lateral flow over a distance of several thousands of kilometres (Elliot et al 1999; Elliot & Fleming 2000), through conduits in the crust and it is consequently difficult to relate magma composition to lithospheric and crustal age. For each location described, the evidence for the age of the oldest crust, usually the oldest identifiable magmatic event, is described. Ideally, this age would give an approximation to the time when the sub-crustal lithospheric mantle may have become fixed to the continental base. In practice, several of the 'oldest' crustal ages probably are significantly younger than the real age of crustal formation, as discussed below.

Antarctic Peninsula The Antarctic Peninsula is the largest crustal block of West Antarctica (B. Storey et al 1988). It developed along the continental margin of Gondwana from Mesozoic to Cenozoic times. For much of this time, the margin was a subduction zone and the block is dominated by a batholith and associated volcanic and arcrelated sedimentary rocks (Storey & Garrett 1985; Leat etal 1995). The crustal volume along the margin grew during several episodes of accretion in Mesozoic times (Storey & Garrett 1985). Therefore, major terrane boundaries are likely to exist in the Antarctic Peninsula (Vaughan & Storey 2000). The last major terrane accretion event was probably in midCretaceous times, when one or more terranes of western Palmer Land docked with the continental margin (Vaughan et al 20020, b). Mesozoic mafic magmatism is represented by dykes and volcanic rocks with various relationships to episodes of subduction (Scarrow et al 1998). These contrast with Cenozoic, postsubduction mafic volcanic rocks which have compositions that are dominated by asthenospheric sources unmodified by subduction (Hole et al 1993). Leat et al (2002) interpreted the compositions of widespread Cretaceous dykes in the Antarctic Peninsula to indicate the tapping of two distinct mantle sources. Dyke compositions from Oscar II Coast (east coast of Graham Land) were compared to those of dykes of similar age (mid-Cretaceous) from Black Coast

ANTARCTIC LITHOSPHERIC DOMAINS (east coast of Palmer Land). High-Mg Oscar II Coast samples have £ Ndt values from +7.3 to +7.9 that, together with trace element data, indicate derivation from a subduction-modified asthenospheric source. In contrast, most dykes from the Black Coast have much lower £ Ndt values of -2.8 to +3.4 and cannot have been derived from the same asthenospheric source. Their source is interpreted to be in the subcontinental lithospheric mantle (Leat et al 2002), consistent with the earlier work of Scarrow et al (1998) on mafic dykes from the west coast of the Antarctic Peninsula. The dykes are difficult to date, but those from the east coast are thought, on the basis of K-Ar chronology and structural relationships, to have been emplaced at c. 106-126 Ma (Leat et al. 2002). In most of the southern Antarctic Peninsula, the basement appears to be represented by scattered outcrops of Palaeozoic to Early Jurassic high-grade rocks (Millar et al 2002). The oldest known zircon U-Pb protolith age is 435 ± 8 Ma from a gneiss at Mount Eissenger, Palmer Land (Millar et al 2002) and this has been used as the age of the crust.

Thurston Island Thurston Island and the adjacent Jones Mountains are the surface expression of another of the terranes or crustal blocks of West Antarctica. Thurston Island consists of a basement sequence of calc-alkaline igneous rocks recording Pacific margin magmatism of Carboniferous to Late Cretaceous age (White & Craddock 1987; Leat et al 1993; Pankhurst et al 1993). These basement rocks are overlain by dominantly asthenosphere-derived alkali basalts of Miocene age erupted after the cessation of subduction along this part of the margin (Hole et al 1994; Hart et al 1995). It is not possible to identify a distinct lithospheric component in the Thurston Island mafic rocks. All the mafic rocks for which Nd isotopic data exist (Pankhurst et al 1993) have been included in the dataset. These include rocks dated at between 300 Ma and 100 Ma and have £Ndt between -1.1 and +1.7 and eSrt between -1.1 and +19. It is only possible to say that this restricted range of values is probably close to that of local lithospheric mantle melts. The oldest known rocks in the Thurston Island area are orthogneisses at Morgan Inlet dated at 309 ± 5 Ma by Rb-Sr (Pankhurst et al 1993) and 348 ± 3 Ma by zircon U-Pb (Pankhurst pers. comm. 2004). However, Nd model ages in intermediate and silicic rocks are varied and range up to about 1200 Ma, implying

365

the presence of older crustal material (perhaps sediment-derived) (Pankhurst et al 1993). The age of 348 Ma is used as the best estimate of formation of the local crust, with the knowledge that it is, at best, a minimum age.

Marie Byrd Land Marie Byrd Land is one of the largest terranes or crustal blocks of West Antarctica and it may have consisted of two separate crustal blocks before Cretaceous times (DiVenere et al 1995). It was subdivided by Pankhurst et al (1998) into two provinces with different geochronological and geochemical histories, the interior Ross Province and the exterior Amundsen Province. The basement rocks were intruded during rifting of New Zealand from Antarctica by a suite of mid-Cretaceous mafic dykes (107 ± 5 Ma by Ar/Ar; 101 ± 1 Ma by U-Pb) and related silicic intrusions (Weaver et al 1994; Storey et al 19996; Mukasa & Dalziel 2000). All these rocks are overlain locally by extensive, asthenosphere-derived, mafic-silicic, alkalic volcanic sequences of Late Cenozoic age (Hole 6 LeMasurier 1994; Panter et al 1997). The geochemistry of the suite of midCretaceous mafic dykes is complex, with several high- and low-Ti groups. Storey et al (19995) interpreted the suite to represent mixtures between melts of lithospheric mantle and plume mantle. They identified one sample, MB.219.4, as the most likely candidate for the lithospheric end-member. This sample has an £ Ndt value of -0.3 and an £Sr value of +8.7, the most evolved of any known mafic rock in Marie Byrd Land, and these values are used in the dataset. Sample MB.219.4 intrudes the Amundsen Province on the Ruppert Coast. The oldest known rock of the Amundsen Province is an orthogneisss dated at 446 ± 1 6 Ma by Rb-Sr (Pankhurst et al 1998) and 514 ± 5 Ma by zircon U-Pb (Pankhurst pers. comm. 2004). Nd model ages of about 1000-1300 Ma imply older components in the crust, perhaps sedimentderived (Pankhurst et al 1998). However, Re-Os model ages ranging from zero to over 1100 Ma in mantle xenoliths have been used to suggest a Proterozoic age for lithosphere formation, at least in the Ross Province (Handler et al 2003). Therefore, 514 Ma is taken as the crustal age, but is considered a minimum age.

Victoria Land The widespread and voluminous Cenozoic volcanism of the McMurdo Volcanic Group in

366

P.T. LEATETAL.

Victoria Land appears to be related to extension of the Ross Sea rift and a possible underlying mantle plume (Tessensohn & Worner 1991; Worner 1999; Rocchi et al 2002). The products are varied compositionally, with a large number of silicic types. Rocholl et al (1995) and Worner (1999) interpreted compositions of mafic volcanic rocks to represent three mantle sources, depleted MORB-source mantle (DMM), HIMU mantle and enriched mantle (EM). Most of the mafic magmas were dominated by the asthenospheric DMM and HIMU sources. The EM source is interpreted as situated in the lithospheric mantle and is represented by two basaltic samples, MA-9 and MA-117, from the Malta Plateau volcanic complex (Rocholl et al 1995). These samples have 8Ndt values of +1.3 to +3.3 and eSrt values of -10.4 to +7.5. The Malta Plateau volcanic complex overlies the boundary between the Robertson Bay and Bowers terranes. Both terranes consist of Cambrian to Early Ordovician volcano-sedimentary units metamorphosed to low grade (Stump 1995; Tessensohn & Henjes-Kunst 2005). They appear to represent elements of a subducting margin and were juxtaposed and metamorphosed at c. 500 Ma, during the Ross Orogeny. The age of the oldest crust is poorly known for both terranes, but there are no reliable dates earlier than Middle Cambrian (c. 530 Ma) (Stump 1995). Ellsworth Mountains The Ellsworth Mountains form part of the Ellsworth-Whitmore Mountains crustal block within West Antarctica. The mountain range contains a Middle Cambrian to Permian sedimentary succession which was deformed during an Early Mesozoic tectonic event (Curtis & Storey 1996). Stratigraphic, structural and palaeomagnetic data indicate strongly that the crustal block was situated close to Coats Land from at least Middle Cambrian times until the break-up of Gondwana, when it moved away from the Gondwana core to its present position in West Antarctica (Curtis & Storey 1996; Randall & MacNiocaill 2004). The sequence hosts Middle Cambrian to possible basal Upper Cambrian volcanic rocks hosted by the Springer Peak and Liberty Hills sedimentary formations. The sequence is dated at c. 515-500 Ma (Randall et al 2000). The volcanic rocks are dominantly mafic in composition and interpreted by many workers to have erupted in a rift environment (Hjelle et al 1982; Vennum et al 1992; Curtis et al 1999) - although other settings

such as a back-arc basin (Curtis & Storey 1996) and continental arc (Duebendorfer & Rees 1998) have also been suggested. The mafic volcanic rocks range from N-MORB-like to shoshonitic and lamprophyric compositions, with £Nd505 values from -2.0 to +5.1 (Curtis et al 1999). They are interpreted as melts of at least two mantle sources. The MORB-like magmas were erupted from the central part of the rift structure from asthenospheric sources. The more alkaline magmas, including a spessartite calc-alkaline lamprophyre, are interpreted to have been derived from lithospheric sources (Curtis et al 1999). These have £Nd505 values in the range -1.6 to +1.3, and ESr505 in the range -0.1 to +47 (recalculated from Curtis et al (1999) and omitting gabbro sample HN1, which is of uncertain age and provenance, and may have high £Sr as a result of hydro thermal alteration). Pre-Middle Cambrian basement is not exposed on the Ellsworth-Whitmore Mountains block. It is thought, from geophysical evidence, that gneisses similar to those exposed in the Haag Nunataks continue under the EllsworthWhitmore Mountains (Curtis & Storey 1996). The Haag Nunatak gneisses have been Rb-Sr dated as 1176 ± 76 Ma (Grenvillian: Millar & Pankhurst 1987) and this is taken as the crustal age of the Ellsworth Mountains. Falkland Islands The Falkland (Malvinas) Islands consist of Silurian to Permian sedimentary rocks overlying Proterozoic basement (Greenway 1972; Marshall 1994; Hunter & Lomas 2003). Currently part of the South America plate, the crustal block was situated in Gondwana between southeast Africa and the Dronning Maud Land-Ellsworth Mountains margin of Antarctica (Marshall 1994; Curtis & Hyam 1998 and references therein). The Proterozoic basement, exposed at Cape Meredith, West Falkland, is cut by several groups of mafic dykes. One group consists of calc-alkaline lamprophyres (minette, spessartite and vogesite), which are interpreted as partial melts of lithospheric mantle (Thistlewood et al 1997). Three of the lamprophyres were dated by K-Ar at 523 ± 9 Ma (sample FI.204.2), 473 ± 12 Ma (FI.215.1) and 306 ± 8 Ma (FI.219.1) (Thistlewood et al 1997). Thomas et al (1998) redated (K-Ar) the same dyke as sample FI.204.2 and obtained a similar age of 503 ± 6 Ma. They also dated a fourth lamprophyre dyke at 520 ± 5 Ma and suggested that all the lamprophyres were emplaced at c. 520 Ma (Middle-Late Cambrian). The 40Ar/39Ar results here (Table 1)

367

ANTARCTIC LITHOSPHERIC DOMAINS

give an age of 531 ± 6 Ma for sample FI.215.1 (Fig. 3). This confirms a Cambrian emplacement for this dyke and suggests, with the K-Ar data, that three of the four dated lamprophyres are Cambrian. However, sample FI.219.1 gave an 40 Ar/39Ar age of 385 Ma (Devonian; Fig. 3). Although this age is not well-constrained, being based on the lowest age in the saddle in Figure 3, it suggests strongly that at least this single lamprophyre was emplaced later. The lamprophyres are interpreted as smalldegree melts of lithospheric mantle because of their potossic compositions, their high abundances of incompatible trace elements such as Ba, Th, Rb, Sr and LREE, and their Sr and Nd isotopic compositions which lie outside the range of contemporary OIB (Thistlewood et al 1997). They have ENdt values (recalculated, using the new 40Ar/39Ar geochronology, from Thistlewood et al 1997) of -0.4 to -3.2 and eSrt values of +31.3 to + 43.4 (Table 2).

The oldest age obtained from the Proterozoic gneisses intruded by the lamprophyres at Cape Meredith is a U-Pb age of 1118 ± 8 Ma on zircon, interpreted as the age of eruption of volcanic protoliths (Jacobs et al. 1999) and this is used as the crustal age.

Shackleton Range Mafic dykes are abundant in the Shackleton Range and have been divided into several groups with different distributions, ages and compositions (Clarkson 1981; Hotten 1993; Spaeth et al 1993; Techmer et al 1995). The dykes range in age from Late Proterozoic to Jurassic. One group of dykes from Herbert Mountains, Pioneers Escarpment and La Grange Nunataks, all in the north and east of the Shackleton Range, form an alkali-rich, mildly potassic suite including calc-alkaline lamprophyres (Group II of Techmer et al 1995).

Fig. 3. Results of Ar/Ar age measurements on mafic dyke samples from Cape Meredith, Falkland Islands. Box heights are 2a uncertainties.

Table 2. Sr and Nd isotope data for lamprophyric and shoshonitic dykes from the Falkland Islands and Shackleton Range Sample

Age (Ma)

Method

Rb

Sr

FI.204.2 FI.215.1 FI.219.1 Z.628.1 Z.736.4 Z.914.4

523 531 385 466 256 388

K-Ar Ar-Ar Ar-Ar K-Ar Ar-Ar Ar-Ar

90 198 148 63 56 66

887 780 673 472 555 384

87

Sr/86Srm

0.708979 0.711649 0.710603 0.708775 0.709216 0.709738

87

Sr/86Srt

eSrt

0.707347 0.706082 0.707110 0.706208 0.708152 0.706987

41.1 31.3 43.4 32.0 56.1 41.7

147

Sm/144Nd

0.1294 0.1346 0.1270 0.1063 0.1282 0.1253

143

Nd/144Ndm

0.512387 0.512306 0.512302 0.512046 0.512295 0.512238

143

Nd/144Ndt

0.511944 0.511838 0.511982 0.511722 0.512080 0.511920

ENdt

-0.4 -2.3 -3.2 -6.2 -4.5 -4.3

Sr and Nd data for Falkland Islands samples (FI) recalculated from Thistlewood et al. (1997). K-Ar age of sample FI.204.2 from Thistlewood et al. (1997); K-Ar age of sample Z.628.1 recalculated from Rex (1972). Rb and Sr concentrations by XRF, University of Keele, 1993. The Falkland Island samples are described and located by Thistlewood et al. (1997) and the Shackleton Range samples by Clarkson (1981) and Techmer et al. (1995).

368

P.T.LEAT ET AL.

These are interpreted as melts of lithospheric mantle. The ages of the Group II dykes have been uncertain. K-Ar methods have yielded ages from Ordovician to Late Carboniferous (304 ± 12 Ma to 466 ± 18 Ma) (recalculated after Rex 1972; Hofmann et al 1980; Hotten 1993; Techmer et al 1995) that might record emplacement in several episodes or the effects of alteration. The uncertainty is enough to affect time-corrected ENd and eSr values significantly. Two Group II samples were dated by 40 Ar/39Ar geochronology (Table 1). Sample Z.736.4 is a potassic, shoshonitic basalt dyke from La Grange Nunataks. A plagioclase separate produced steps with ages of 256-356 Ma, with a saddle-like spectrum, probably the result of excess Ar. The age of this sample is not well constrained. The minimum of 256 ± 2 Ma is the most likely age. This Ar/Ar age is younger than two K-Ar whole-rock ages, 304 ± 12 Ma and 319 ± 8 Ma, obtained by Rex (1972, recalculated age) and P. Leat (unpublished data, 1994), respectively, on the same sample. Sample Z.914.4 is a potassic, shoshonitic basalt dyke from the Herbert Mountains. It produced a poorly constrained Ar/Ar age, with a saddle-like spectrum, probably a result of excess Ar, with steps ranging from 388 Ma to 495 Ma using a plagioclase separate. The minimum step of 388 ± 22 Ma is the most likely age. The Ar/Ar ages support the interpretation that the Group II, potassic dykes were emplaced over hundreds of millions of years. If the (recalculated) whole-rock K-Ar age of 466 Ma for sample Z.628.1 is accepted, they were emplaced from Ordovician to Permian times. However, there is a small range in timecorrected Nd and Sr isotope values: eNdt values range from -6.2 to -4.2 and eSrt from 32.0 to 56.4 (Table 2), suggesting that the magmas tapped similar mantle sources which probably resided in the lithospheric mantle. The basement of the Shackleton Range consists of several distinct Precambrian units brought together during the early Palaeozoic (c. 500 Ma) Ross/Pan-African Orogeny (Tessensohn et al. 1999). The Herbert Mountains, Pioneers Escarpment and La Grange Nunataks are underlain by thrust structures which bring together Proterozoic basement, Cambrian or older oceanic crust and Cambrian sedimentary units (Talarico et al 1999; Tessensohn et al 1999). The oldest U-Pb date from the Proterozoic rocks is for a granite protolith to a gneiss, emplaced at 2328 ± 47 Ma and affected by migmatization at 1715 ± 6 Ma (Brommer et al 1999). The crustal age associated with the

source of the potassic magmatism is uncertain, but is probably older than the Ross deformation because of the low eNdt values of the mafic dykes and is, therefore, at least Proterozoic.

Dronning Maud Land - Vestfjella The western part of Dronning Maud Land (between 0° W and 16° W) contains abundant dominantly basaltic lavas and igneous intrusions that have been dated as Jurassic in age (mostly in the range 178-198 Ma: Brewer et al 1996; Duncan et al 1997; Fazel et al 2003; Zhang et al 2003; Riley et al 2005). The magmatism was part of the huge Karoo-Ferrar basaltic province that formed during the initial stages of Gondwana break-up between southern Africa and Antarctica (Duncan et al 1997; Elliot & Fleming 2000; Riley & Knight 2001). The basalts have a wide range of compositions that are interpreted as derived from asthenospheric sources, with contributions from lithospheric mantle and crust (Luttinen & Furnes 2000; Riley et al 2005). In the Vestfjella area, a thick sequence of basaltic lavas is intruded by lamproite dykes which have been dated at 158.7 ± 1.6 Ma (Luttinen et al 2002), some 22 million years after the probable peak of Jurassic lava extrusion in the area (Duncan et al 1997). The Kjakebeinet lamproites are interpreted on geochemical and mineralogical grounds as derived by partial melting of trace elementenriched lithospheric mantle in response to either thermal effects of the Karoo mantle plume, or extension related to rifting of Antarctica from southern Africa (Luttinen et al 2002). The lamproites have £Ndi59 values from -6.0 to -6.7 and eSri59 values from +14.3 to +17.5 (Luttinen et al 2002). The age of the continental crust underlying the lamproite location is equivocal, as an unexposed boundary between Archaean craton and Proterozoic belt lies in the vicinity of Vestfjella (Moyes et al 1993). The oldest known rocks in the Archaean Grunehogna craton are granites which are dated by Rb-Sr at 2768 ± 88 Ma, thought to be approximately the granite emplacement age (Barton et al 1987). The closest significant outcrop of the Proterozoic Maudheim Province to Vestfjella is in Heimfrontfjella, where the oldest reported U-Pb zircon ages are 1161 ± 10 Ma (Bauer et al 2003) and 1171 ± 25 Ma (Jacobs et al 2003). The Maudheim Province in the area was deformed and metamorphosed at c. 500 Ma and is interpreted as part of the Pan-African Orogeny (Jacobs et al 2003). The Vestfjella lavas were probably emplaced above the suture between

ANTARCTIC LITHOSPHERIC DOMAINS the crustal provinces (Luttinen & Furnes 2000). This suture may have been important in permitting ascent of the lamproites. It is, therefore, likely that the Archaean lithosphere was involved as a mantle source of the lamproites.

369

Ruker terrane. U-Pb data from zircons have been used to interpret a crystallization age of 3160 Ma for a gneissic granite from this terrane (Boger et al 2001), which is broadly consistent with a wider geochronological database (Mikhalsky et al 2001).

Dronning Maud Land - Schirmacher Oasis Several lamprophyres, including minettes, intrude Proterozoic high-grade rocks at Schirmacher Oasis (Hoch & Tobschall 1998). The minettes are interpreted, on elemental and isotopic grounds, as partial melts of enriched lithospheric mantle (Hoch & Tobscall 1998; D'Souza & Chakraborty 2000; Hoch et al 2001). The dykes are thought, on the basis of Rb-Sr geochronology and regional relationships, to have been emplaced at about 445 Ma (Dayal & Hussain 1997; Hoch et al 2001). The minettes have £Nd445 values that range from -3 to -20, and £51445 values that range from +31 to +89 (Hoch et al 2001). The ages of the high-grade metamorphic rocks at Schirmacher Oasis are poorly known. Grew & Manton (1983) interpreted U-Pb data on zircon and allanite crystals to suggest an age of c. 1500 Ma for initial crystallization of gneisses. However, in a more detailed U-Pb study of zircon from the nearby Orvinfjella area, Jacobs et al. (1998) suggested an age of 1137 ± 21 Ma for the volcanic protoliths and a c. 1190 Ma inherited component. The existing crust was deformed strongly and metamorphosed to granulite facies at c. 530-515 Ma, during the Pan-African event which may record closure of the 'Mozambique Ocean' and amalgamation of East and West Gondwana (Jacobs et al 1998; Fitzsimons 2000). Therefore, 1137 Ma is taken as the crustal age relating to the Schirmacher minettes.

Southern Prince Charles Mountains Lamproites intrude the Ruker terrane of the Southern Price Charles Mountains (Mikhalsky et al 2001). One lamproite dyke at Mount Bayliss was K-Ar dated at 413 ± 10 Ma on Krichterite and 430 ± 12 Ma on K-arfvedsonite (Sheraton & England 1980). A second lamproite dyke from Mount Rubin yielded a mineral Rb-Sr isochron of 461 ± 23 Ma (Mikhalsky et al 1994), interpreted as the age of emplacement (Mikhalsky et al 2001). The dykes probably belong to the same OrdovicianSilurian intrusion event. The Mount Rubin dyke has an £Nd461 value of -8.7, and an £Sf46] value of +171 (Mikhalsky et al 1994). The lamproite dykes intrude the Archaean

Gaussberg Gaussberg is a volcanic cone consisting of lamproite lavas erupted some 56 000 years ago close to the continental margin (Sheraton & Cundari 1980; Collerson & McCulloch 1983; Tingey et al 1983). The lamproites are isolated from other known regions of Phanerozoic magmatism on the Antarctic continent and are interpreted as generated by small degrees of melting of an enriched mantle source (Sheraton & Cundari 1980; Collerson & McCulloch 1983). Some authors (Nelson et al 1986; Williams et al 1992) suggested that this source was situated within the underlying lithospheric mantle, consistent with most interpretations of lamproite genesis. An alternative view, that the lamproite source was subducted sediment within the asthenosphere was proposed by Collerson & McCulloch (1983) and Murphy et al (2002). The lamproites have eNd values of -12.2 to -15.0 and £Sr values of +67 to +77 (Murphy et al 2002). No local basement is exposed at Gaussberg, but it is probable that the local basement is Archaean. The nearest basement outcrops are at Vestfold Hills, to the west, and Hunger Hills, to the east, both of which expose Archaean gneisses (dated at 2526 Ma and 2641 Ma, respectively (Black et al 1991; Sheraton et al 1993). Two crustal xenoliths in the volcanic rocks give Nd model ages (TDM) of 2174 Ma and 2327 Ma (Collerson & McCulloch 1983), supporting the presence of underlying Archaean basement. Results Data for the lithosphere-derived mafic magmas are summarized in Table 3. There is a very strong relationship between eNdt and crustal age (Fig. 4). Two domains can be defined, one being cratonic and circumcratonic (after Menzies 1989), and the other being limited to the protoPacific margin of Gondwana. In this paper, the convention (e.g. Menzies 1989; Pearson & Nowell 2002; Pearson et al 2003) of applying the term 'cratonic' to Archaean (>2500 Ma) terrains is followed, and the term 'circumcratonic' is used to describe Proterozoic areas around the margins of cratons that may contain

370

P. T. LEAT ETAL.

Table 3. Summary of Antarctic lithosphere-derived mafic magmatism Region

Crustal age (Ma)

Mafic lithosphere-derived rocks: ranges Age (Ma)

e

References

Nd t

Antarctic Peninsula

435

120

-2.7 to+3.5

-3.3 to+30.2

Leatetal. (2002), Millar etal (2002)

Thurston Island

348

100-300

-1.1 to 1.7

-1.1 to+19

Pankhurst et al (1993), Pankhurst (pers. comm. 2004)

Marie Byrd Land (Amundsen Province)

514

107

-0.3

+8.7

Pankhurst et al. (1998), Storey et al (1999b)

Victoria Land

530

15

+1.3 to+3.3

-10.4 to+7.5

Ellsworth Mountains

1176

505

-1.6 to+1.3

-0.1 to+47

Millar & Pankhurst (1987), Curtis etal. (1999)

Falkland Islands

1118

522-333

-0.4 to-3.2

+31.3 to+43.4

Thistlewood etal (1997), Jacobs et al (1999), this paper

Shackleton Range

2328

249-466

-4.3 to -6.2

+32.0 to+56.1

Rex (1972), Techmer et al (1995), Brommeretal. (1999), this paper

Dronning Maud Land - Vestfjella

2768

159

-6.0 to -6.7

+14.3 to + 17.5

Barton etal (1987), Luttinen et al (2002)

Dronning Maud Land - Schirmacher Oasis

1137

445

-3.0 to -20

+31 to +89

Jacobs et al (1998), Hoch etal (2001)

Southern Prince Charles Mountains

3160

461

-8.7

Gaussberg

2641

0.056

+171

-12.2 to-15.0

+67 to+77

Rocholletal. (1995), Stump (1995)

Mikhalskygfa/. (1994, 2001), Roger etal (2001) Tingey et al (1983), Sheraton et al (1993), Murphy et al (2002)

ENdt and 8Srt calculated at crystallization age.

tectonically reworked Archaean rocks. The cratonic and circumcratonic domain is characterized by Middle Proterozoic to Archaean crustal ages (1137-3160 Ma) and low ENdt values of -20 to -3 for the lithosphere-derived mafic rocks. The contrasting Gondwana margin domain has much younger, Middle Proterozoic to Palaeozoic crustal ages (1176-309 Ma) and much higher eNdt values of -3.2 to +3.5 for its lithosphere-derived mafic rocks. The domain could be also described as 'non-cratonic' (Pearson et al 2003). The largest source of error in the relationship in Figure 4 is in the crustal ages. The crustal age for each location is the oldest known age of crystalline rocks and so the plotted ages are minima. The relatively young age of basement gneisses at Schirmacher Oasis, within the cratonic and circumcratonic domain, is based on U-Pb analyses of zircons (see above), but Nd isotope data presented by Jacobs et al (1998)

give depleted mantle model ages of 1.07 Ga to 2.05 Ga, implying that there may have been older Proterozoic crust in the region. Likewise, the age of crustal basement beneath much of West Antarctica is uncertain. Millar et al (2002) dated the oldest exposed crust in the Antarctic Peninsula at 435 Ma, although isotope data from granitoids indicate that Proterozoic lower crust may exist in the Antarctic Peninsula (Millar et al 2001). The existence of crust older than 309 Ma beneath the Thurston Island block is likely. In the case of Marie Byrd Land, both Nd model age data and Os isotope data have been used to suggest the existence of Proterozoic lithosphere (Pankhurst et al 1998; Handler et al 2003). A secondary source of error in Figure 4 is the possible effects of contamination of the mafic magmas by crust during their uprise. This error is minor. Contamination by crust will have had least effect on the lamproite magmas of

ANTARCTIC LITHOSPHERIC DOMAINS

Fig. 4. Plot of 8Ndt versus crustal age for mafic dykes and lavas from Antarctica and the Falkland Islands that were derived from lithospheric mantle sources.

Gaussberg, Prince Charles Mountains and Vestfjella. These magma have high contents of Sr and Nd that are between 3 and 14 times the abundances of these elements in bulk and upper continental crust (Rudnick & Fountain 1995). The maximum amount of contamination by crust permissible in these magmas is limited by their lamproitic composition, which is quite unlike that of continental crust. As an example of the effects of such contamination, 20% contamination of Gaussberg lamproite magma (Murphy et al 2002) by typical Archaean gneiss having present day e Nd of as low as -32 (Mikhalsky et al 2001) will result in an ENd shift in the magma of approximately 0.7. This is considerably less than the variation in the Gaussberg suite as a whole (Fig. 4). In the case of the mafic rocks from the Gondwana margin domain, their lower Nd abundances make them more susceptible to contamination by crust. However, late Proterozoic to Early Palaeozoic crustal lithologies in the region commonly have £Nd values of approximately -6 to +1.3 (at the time of eruption of the mafic magmas) (Thistlewood et al 1997; Pankhurst et al 1998; Millar et al 2001 and references therein). The crustal values are, therefore, similar to those of the mafic magmas considered in this study. As the Gondwana margin domain mafic rocks, whose isotope compositions are the more likely to have been affected by contamination by crust, have higher E Ndt than the cratonic and circumcratonic mafic rocks, which were certainly not significantly affected by contamination, it

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follows that the relationships in Figure 4 are not a result of contamination of magmas by crust. The variation in Nd isotope data for the mafic lithosphere-derived melts (Fig. 4) does, however, indicate that there is a substantial difference in lithosphere mantle evolution between the cratonic and circumcratonic and Gondwana margin areas. In particular, the lower 8Nd values in the cratonic and circumcratonic domain suggest that the lithosphere in that region has been enriched metosomatically in light rare earth elements (high Nd/Sm) and isolated from the convecting mantle since at least the Middle Proterozoic. In contrast, lithospheric mantle of the Gondwana margin domain was isolated later, probably during Late Proterozoic to Palaeozoic times. The relationships depicted in Figure 4 are characteristic of other continental areas that include ancient cratons and mafic volcanic rocks that have erupted through them e.g. Scotland, southern Africa, the western USA and western Australia (Fraser et al 1985; Nelson et al 1986; Menzies et al 1987; Menzies 1989). The mafic lithosphere melts from the Antarctic domains are plotted relative to established global mantle reservoirs in an E Ndt versus eSrt diagram (Fig. 5). The enriched mantle endmembers (EMI and EMU) are normally associated with sub-continental lithospheric mantle (although they are also present in sub-oceanic mantle), whereas DMM, HIMU and prevalent mantle (PREMA) are normally asthenospheric (Zindler & Hart 1986). The mafic rocks plot

Fig. 5. Plot of eNdt versus eSrt for mafic dykes and lavas from Antarctica and the Falkland Islands that were derived from lithospheric mantle sources. Symbols as in Figure 4. Global mantle reservoirs are after Zindler & Hart (1986).

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between these end-members and can be assumed to represent gradations between the end-members. The rocks from the Gondwana margin domain have limited ranges in both ENdt and eSrt, but these appear to form a shallow trend between the asthenospheric end-members and EMIL This is interpreted as indicating that the domain is dominated by EMU, although its relatively young lithosphere ages mean that extreme EMU compositions have not developed. Some of the samples plot slightly below this trend, indicating that there could be a very small input from EMI in the domain. EMU originates in the mantle as a result of enrichments in Rb/Sr and LREE, possibly due to metasomatism by hydrous fluids (e.g. Menzies & Hawkesworth 1987), and probably associated with subduction of continental sediments (Zindler & Hart 1986; Dickin 1995). It is suggested here that Antarctic EMU originated during late Proterozoic to Mesozoic subduction beneath Gondwana. This mechanism was proposed for the source of the Jurassic Ferrar magmas by Kyle (1980) and Kyle et al (1983). The mafic rocks from the cratonic and circumcratonic domain form a steep trend in Figure 5, between the asthenospheric mantle end-members and a component near, but to the right of EMI. The single obvious outlier from this trend is the Southern Prince Charles lamproite that plots near, and is apparently dominated by, EMIL The plotted ESrt of +171 for this sample comes from the initial 87Sr/86Sr ratio of 0.716 derived from a mineral isochron (Mikhalsky et al 1994). The steep trend for the rest of the cratonic and circumcratonic samples is taken to indicate that these lithosphere-derived melts are dominated by EMI. This is consistent with the domination of sub-continental mantle by EMI in other Archaean and Early Proterozoic domains such as the Hebridean craton of Scotland (Menzies et al. 1987), the Wyoming craton, USA (Menzies 1989) and the Aldan Shield, Siberia (Mitchell et al 1994). The EMI trend toward low £Ndt and low 8Srt is the result of enrichment in LREE in the absence of enrichment in Rb-Sr, and may be related to metasomatism of the source by small-degree, asthenosphere-derived, anhydrous melts (e.g. Menzies & Hawkesworth 1987). Interestingly, the EMI domain occurs on both sides of the possible Pan-African suture in East Antarctica (Fig. 1), although whether this is coincidental or was caused by a common pre-Pan-African evolution is not obvious.

Discussion Ritzwoller et al (2001) showed that the East Antarctic craton has a very different velocity structure to that of West Antarctica. East Antarctica is characterized by faster (= cooler) than global average seismic velocities to a depth of about 250 km. It has a depth-velocity structure that is comparable to those of stable cratons of the world, but is slower (= hotter) than is normal for Archaean cratons between about 50 km and 200 km depth (Ritzwoller et al 2001). There are no large-scale differences between different parts of East Antarctica and the line of the possible Pan-African suture between East and West Gondwana does not appear to be shown by the seismic data. West Antarctica has a very different seismic velocity structure shown by the same data (Ritzwoller et al 2001). It has significantly slower velocities than East Antarctica and is more similar to oceanic regions - its velocity structure is indistinguishable from that of ocean asthenosphere below about 200 km (Ritzwoller et al 2001). Above 200 km depth, the mantle has an anomalously low velocity structure that seems to be centred on the central part of the Cretaceous-Recent West Antarctic rift system (Behrendt et al 1991). This seismic structure of Antarctica is consistent with the geochemical evidence. East Antarctica appears to consist of an amalgamation of small Archaean cratons and Proterozoic domains that are marginal to those cratons. These elements were brought together and stabilized in their current dispositions during the Pan-African Orogeny (Fitzsimons 2000; Roger et al 2002; Jacobs et al 2003). The lithospheric thickness of 200 km indicated by the seismic evidence is consistent with the presence of Archaean cratonic roots at least locally beneath East Antarctica. Archaean cratons normally have stable lithospheric mantle roots up to about 250 km thick, contrasting with the <150 km lithosphere thicknesses of circumcratonic areas (e.g. Boyd & Gurney 1986; Menzies 1992; Pearson & Nowell 2002). These thick lithospheric roots probably were the source of subsequent lithosphere-derived mafic magmas rich in the EMI component that were erupted in East Antarctica. The slower seismic velocities of East Antarctica relative to stable cratons of the world is consistent with the relatively recent final stabilization of the region in Pan-African times. The low E Ndt values of minettes from Schirmacher Oasis are anomalous relative to the known local crustal age of 1137 Ma. This may indicate that Archaean lithospheric roots

ANTARCTIC LITHOSPHERIC DOMAINS are more extensive in this region than suggested by current surface age data. The seismic data indicate that the thick lithospheric root is absent beneath most of the Transantarctic Mountains. This is consistent with the relatively young lithospheric mantle indicated by the compositions of Cenozoic lavas in Victoria Land (Rocholl et al 1995) and the tectonic evolution of the Ross Orogeny, which occurred on the continental margin and was characterized by accretion of Cambrian terranes. Interestingly, the seismic data indicate no thinning of lithosphere in the western Dronning Maud Land to Coats Land part of East Antarctica, despite the large volume of Jurassic magmas erupted in that area. These volcanic rocks are thought commonly to have been generated from one or more mantle plumes (Storey et al 2001; Riley et al 2003; 2005). If these mantle plumes existed, they clearly did not significantly erode the lithospheric mantle beneath East Antarctica. This is consistent with generation of the majority of the volume of these magmas in a relatively narrow zone of high extension between the rifted margins of southeast Africa (now marked by the Lebombo monocline) and the Explora Escarpment and its lateral equivalents (Sweeney et al 1994; Elliot & Fleming 2000). The absence of seismically discernable lithospheric roots beneath West Antarctica is consistent with the geochemical evidence showing an absence of ancient lithosphere in the magma sources of the region. The oldest known basement rocks in West Antarctica are the late Middle Proterozoic gneisses exposed on Haag Nunataks, which are thought also to underlie the Ellsworth-Whitmore Mountains block (Millar & Pankhurst 1987; Curtis & Storey 1996). This part of West Antarctica must have been in place as part of the Gondwana margin by Middle-Late Cambrian times (Randall & MacNiocaill 2004). Amalgamation of the rest of West Antarctica, probably by a process of terrane accretion, can be assumed to have occurred after the Cambrian Ross Orogeny (Pankhurst et al 1998; Vaughan & Storey 2000) and was completed probably when the most outboard terranes of the Antarctic Peninsula docked with the continental margin in midCretaceous times (Vaughan et al 2002(2, b). The age of lithospheric mantle below West Antarctica is, therefore, late Middle Proterozoic (at oldest) to Phanerozoic (for the most part), and is not sufficiently old to have developed enriched radiogenic isotope values. The anomalously low seismic velocities indicate that lithospheric mantle is unlikely to

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be thicker than about 50 km below West Antarctica (Ritzwoller et al 2001). This is considerably thinner than the c. 80-100 km that is normal for both oceanic or continental plates that have had normal cooling histories for more than 100 Ma (Parsons & Sclater 1977; Jordan 1988; Doin & Fleitout 1996 and references therein). This implies that the West Antarctic lithosphere has been thinned during late Mesozoic to Cenozoic times, and this is supported by passive seismic data from Marie Byrd Land and the adjacent part of the West Antarctic rift system that indicate crustal thicknesses of 25 km and 21 km, respectively (Winberry & Anandakrishnan 2004). Such thinning was probably caused by three events: extension of West Antarctica during the breakup of Gondwana in Middle Jurassic to Cretaceous times (B. Storey et al 1988; Grunow et al 1991); extension of the West Antarctic rift system in Cretaceous to Recent times (Behrendt et al 1991); and the likely impingement of hot mantle plumes on the base of the lithosphere (Hole & LeMasurier 1994; Hart et al 1997; Storey et al 2001). The current configuration of mantle plumes below West Antarctica is uncertain, but it is likely that one currently exists beneath the Mount Erebus area in the Ross Sea and that a separate mantle plume is situated below Marie Byrd Land (Storey et al 19996 and references therein). Indeed, Behrendt et al (1991) suggested that the Moho and the top of the asthenosphere essentially are coincident beneath part of the West Antarctic rift system, suggesting that the lithospheric mantle has been completely removed locally. Comparisons with other well-studied continental areas (e.g. Menzies 1989) suggest that the two domains identified here have different heat flow characteristics. However, there are insufficient heat flow measurements from Antarctica to confirm such a relationship (Pollack etal. 1993). Shapiro & Ritwoller (2004) used seismic data to infer heat flow values for Antarctica. Their method indicated that heat flow is likely to be nearly three times higher in West Antarctica than in East Antarctica. This is consistent with the interpretation that East Antarctica is a stable, dominantly Precambrian region with a thick lithospheric keel, while West Antarctica is more recently tectonically active and does not have a thick lithospheric keel.

Conclusions •

Palaeozoic to Quaternary mafic magmas that were derived from lithospheric mantle

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•

•

•

•

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sources are scattered across West and East Antarctica and the Falkland Islands (adjacent to Antarctica in Gondwana). The mafic rocks reveal systematic Sr and Nd variations that are linked to tectonic development and crustal age. Two domains are identified from the compositions of the lithospheric mantlederived melts: a cratonic and circumcratonic domain and a Gondwana margin domain. The cratonic and cratonic domain is restricted to East Antarctica, and is characterized by Archaean to Middle Proterozoic crustal ages. It is an amalgamation of small Archaean cratons and Proterozoic rocks marginal to the cratons. Seismic data indicate deep lithospheric roots extending to 250 km below the domain, consistent with at least local preservation of thick Archaean lithospheric roots. Most of the region was reworked technically during the Pan-African Orogeny. The mafic lithosphere-derived melts have low eNdtvalues -20 to -3, consistent with derivation from pockets of Archaean to Proterozoic lithospheric mantle. The Gondwana margin domain includes West Antarctica, Victoria Land and the Falkland Islands and is characterized by Middle Proterozoic to Palaeozoic exposed crustal ages, although Proterozoic basement may be more common in this domain than is evident from maximum measured formation ages of the regional crust. Seismic data indicate that there are no deep lithospheric roots beneath the domain. The mafic lithosphere-derived melts have relatively high eNdt values of-3.6 to +3.5, consistent with derivation from late Proterozoic to Palaeozoic lithospheric mantle. The cratonic and circumcratonic domain is dominated by EMI compositions, similar to those associated with other Archaean to Early Proterozoic areas, such as Hebridean Scotland and the Wyoming craton, USA. EMI compositions occur on both sides of the possible Pan-African suture. The Gondwana margin domain is dominated by EMU, consistent with modification of this mantle region by Proterozoic to Mesozoic subduction beneath Gondwana.

The authors are grateful to M. A. Menzies, D. G. Pearson and R. J. Pankhurst for helpful comments on the paper.

References ANDRONIKOV, A.V. & FOLEY, S.F. 2001. Trace element and Nd-Sr isotopic composition of ultramafic lamprophyres from the East Antarctic Beaver Lake area. Chemical Geology, 175, 291-305. BARTON, J.M., KLEMD, R., ALLSOPP, H.L., AURET, S.H. & COPPERTHWAITE, Y.E. 1987. The geology and geochronology of the Annandagstoppane granite, western Dronning Maud Land, Antarctica. Contributions to Mineralogy and Petrology, 97, 488-496. BAUER, W., JACOBS, X, FANNING, M. & SCHMIDT, R. 2003. Late Mesoproterozoic arc and back-arc volcanism in the Heimfrontfjella (East Antarctica) and implications for the palaeogeography at the southeastern margin of the KaapvaalGrunehogna craton. Gondwana Research, 6, 449-465. BECKINSALE, R.D., TARNEY, I, DARBYSHIRE, D.P.K & HUMM, H.J. 1976. Re-Sr and K-Ar age determinations on samples of the Falkland Plateau basement at Site 330, DSDP. In: BARKER, P.F, DALZIEL, I.W.D. ETAL. Initial Reports of the Deep Sea Drilling Project, 36, 923-927. BEHRENDT, J.C, LEMASURIER, W.E., COOPER, A.K., TESSENSOHN, F, TREHU, A. & DAMASKE, D. 1991. Geophysical studies of the West Antarctic rift system. Tectonics, 10, 1257-1273. BERGMAN, S.C 1987. Lamproites and other potassium-rich igneous rocks: a review of their occurrence, mineralogy and geochemistry. In: FITTON, J.G. & UPTON, B.GJ. (eds) Alkaline Igneous Rocks. Geological Society, London, Special Publications, 30, 103-190. BLACK, L.P., KINNY, P.D., SHERATON, J.W. & DELOR, C.P. 1991. Rapid production and evolution of late Archaean felsic crust in the Vestfold block of East Antarctica. Precambrian Research, 50, 283-310. BOGER, S.D., WILSON, C.J.L. & FANNING, CM. 2001. Early Paleozoic tectonism within the East Antarctic craton: the final suture between east and west Gondwana? Geology, 9, 463-466. BOGER, S.D., CARSON, C.J., FANNING, C.M., HERGT, J.M., WILSON, C.LJ. & WOODHEAD, J.D. 2002. Pan-African intraplate deformation in the northern Prince Charles Mountains, East Antarctica. Earth and Planetary Science Letters, 195, 195-210. BOYD, FR. & GURNEY, JJ. 1986. Diamonds and the African lithosphere. Science, 232, 472-477. BREWER, T.S., REX, D., GUISE, P.O. & HAWKESWORTH, C.J. 1996. Geochronology of Mesozoic tholeiitic magmatism in Antarctica: implications for the development of the failed Weddell Sea rift system. In: STOREY, B.C., KING, E.C. & LIVERMORE, R.A. (eds) Weddell Sea Tectonics and Gondwana Break-up. Geological Society, London, Special Publications, 108, 45-61. BROMMER, A., MILLAR, I.L. & ZEH, A. 1999. Geochronology, structural geology and petrology of the northwestern La Grange Nunataks, Shackleton Range, Antarctica. Terra Antartica, 6, 269-278.

ANTARCTIC LITHOSPHERIC DOMAINS CANNING, J.C., HENNEY, P.J., MORRISON, M.A. & GASGARTH, J.W. 1996. Geochemistry of late Caledonian minettes from northern Britain: implications for the Caledonian sub-continental lithospheric mantle. Minemlogical Magazine, 60, 221-236. CLARKSON, P.D. 1981. Geology of the Shackleton Range: IV. The dolerite dykes. British Antarctic Survey Bulletin, 53, 201-212. COLLERSON, K.D. & McCuLLOCH, M.T 1983. Nd and Sr isotope geochemistry of leucite-bearing lavas from Gaussberg, East Antarctica. In: OLIVER, R.L., JAMES, PR. & JAGO, J.B. 1983. Antarctic Earth Science. Australian Academy of Science, Canberra, 676-680. CURTIS, M.L. & HYAM, D.M. 1998. Late Palaeozoic to Mesozoic structural evolution of the Falkland Islands: a displaced segment of the Cape Fold Belt. Journal of the Geological Society, London, 155, 115-129. CURTIS, M.L. & STOREY, B.C. 1996. A review of geological constraints on the pre-break-up position of the Ellsworth Mountains within Gondwana: implications for Weddell Sea evolution. In: STOREY, B.C., KING, E.C. & LIVERMORE, R.A. (eds) Weddell Sea Tectonics and Gondwana Break-up. Geological Society, London, Special Publications, 108,123-132. CURTIS, M.L., LEAT, P.T., RILEY, T.R., STOREY, B.C., MILLAR, I.L. & RANDALL, D.E. 1999. Middle Cambrian rift-related volcanism in the Ellsworth Mountains, Antarctica: tectonic implications for the palace-Pacific margin of Gondwana. Tectonophysics, 304, 275-299. DALZIEL, I.W.D. & ELLIOT, D.H. 1982. West Antarctica: problem child of Gondwanaland. Tectonics, 1, 3-19. DAYAL, A.M. & HUSSAIN, S.M. 1997. Rb-Sr ages of lamprophyre dykes from Schirmacher Oasis, Queen Maud Land, East Antarctica. Journal of the Geological Society of India, 50, 457-460. DELOR, C.P. & ROCK, N.M.S. 1991. Alkalineultramafic lamprophyre dykes from the Vestfold Hills, Princess Elizabeth Land (East Antarctica): primitive magmas of deep mantle origin. Antarctic Science, 3, 419-432. DICKIN, A.P. 1995. Radiogenic Isotope Geology. Cambridge University Press, Cambridge. DIVENERE, V., KENT, D.V. & DALZIEL, I.W.D. 1995. Early Cretaceous paleomagnetic results from Marie Byrd Land, West Antarctica: implications for the Weddellia collage of crustal blocks. Journal of Geophysical Research, 100, 8133-8151. DOIN, M.P. & FLEITOUT, L. 1996. Thermal evolution of the oceanic lithosphere: an alternative view. Earth and Planetary Science Letters, 142,121-136. D'SouzA, MJ. & CHAKRABORTY, S.K. 2000. Petrology of Early Palaeozoic lamprophyres from Schirmacher Oasis, East Antarctica. Journal of the Geological Society of India, 56, 593-604 DUEBENDORFER, E.M & REES, M.N. 1998. Evidence for Cambrian deformation in the EllsworthWhitmore Mountains terrane, Antarctica: stratigraphic and tectonic implications. Geology, 26, 55-58.

375

DUNCAN, R.A., HOOPER, PR., REHACEK, J., MARSH, J.S. & DUNCAN, A.R. 1997. The timing and duration of the Karoo igneous event, southern Gondwana. Journal of Geophysical Research, 102, 18 127-18 138. ELLIOT, D.H. & FLEMING, T.H. 2000. Weddell triple junction: the principal focus of Ferrar and Karoo magmatism during initial breakup of Gondwana. Geology, 28, 539-542. ELLIOT, D.H., FLEMING, T.H., KYLE, P.R. & POLAND, K.A. 1999. Long-distance transport of magmas in the Jurassic Ferrar large igneous province, Antarctica. Earth and Planetary Science Letters, 167, 89-104. FAZEL, A., KELLEY, S.P., LEAT, P.T. & HAWKESWORTH, C. 2003. Dykes as tracers of continental break-up: argon geochronology of Mesozoic flood basalts of western Dronning Maud Land, Antarctica. Geophysical Research Abstracts, 5, EAE03-A-00047. FITZSIMONS, I.C.W. 2000. A review of tectonic events in the East Antarctic shield and their implications for Gondwana and earlier supercontinents. Journal of African Earth Sciences, 31, 3-23. FOLEY, S.F., VENTURELLI, G, GREEN, D.H. & TOSCANI, L. 1987. The ultrapotassic rocks: characteristics, classification and petrogenetic models. Earth Science Reviews, 24, 81-134. FRASER, K.J., HAWKESWORTH, C.J., ERLANK, A.J., MITCHELL, R.H. & SCOTT-SMITH, B.H. 1985. Sr, Nd and Pb isotope and minor element geochemistry of lamproites and kimberlites. Earth and Planetary Science Letters, 76, 57-70. FREY, FA., CLAGUE, D., MAHONEY, JJ. & SINTON, J.M. 2000. Volcanism at the edge of the Hawaiian plume: petrogenesis of submarine alkalic basalts from the North Arch volcanic field. Journal of Petrology, 41, 667-691. GIBSON, S.A., THOMSON, R.N., LEAT, P.T, MORRISON, M.A., HENDRY, G.L., DICKIN, A.P. & MITCHELL, J.G 1993. Ultrapotassic magmas along the flanks of the Oligo-Miocene Rio Grande rift, USA: monitors of the zone of lithospheric mantle extension and thinning beneath a continental rift. Journal of Petrology, 34, 187-228. GREENWAY, M.E. 1972. The Geology of the Falkland Islands. British Antarctic Survey Reports, 76. GREW, E.S. & MANTON, WI. 1983. Geochronologic studies in East Antarctica: reconnaissance uranium/thorium/lead data from rocks in the Schirmacher Hills and Mount Stinear. Antarctic Journal of the United States, 18 (5), 6-8. GROENEWALD, P.B., GRANTHAM, G.H. & WATKEYS, M.K. 1991. Geological evidence for a Proterozoic to Mesozoic link between southeastern Africa and Dronning Maud Land, Antarctica. Journal of the Geological Society, London, 148, 1115-1123. GRUNOW, A.M., KENT, D.V. & DALZIEL, I.W.D. 1991. New paleomagnetic data from Thurston Island: implications for the tectonics of West Antarctica and Weddell Sea opening. Journal of Geophysical Research, 96,17935-17954. HANDLER, M.R., WYSOCZANSKI, R.J. & GAMBLE, J.A. 2003. Proterozoic lithosphere in Marie Byrd Land, West Antarctica: Re-Os systematics of

376

P. T. LEATETAL.

spinel peridotite xenoliths. Chemical Geology, 196, 131-145. HARLEY, S.L., SNAPE, I. & BLACK, L.R 1998. The evolution of a layered metaigneous complex in the Rauer Group, East Antarctica: evidence for a distinct Archaean terrane. Precambrian Research, 89,175-205. HART, S.R., BLUSZTAJN, J. & CRADDOCK, C. 1995. Cenozoic volcanism in Antarctica: Jones Mountains and Peter I Island. Geochimica et Cosmochimica Ada, 59, 3379-3388. HART, S.R., BLUSZTAJN, J., LEMASURIER, W.E. & REX, D.C. 1997. Hobbs Coast Cenozoic volcanism: implications for the West Antarctic rift system. Chemical Geology, 13, 223-248. HAWKESWORTH, C., KELLEY, S., TURNER, S., LE ROEX, A. & STOREY, B.C. 1999. Mantle processes during Gondwana break-up and dispersal. Journal of African Earth Sciences, 28, 239-261. HERGT, J.M. & BRAUNS, C.M. 2001. On the origin of Tasmanian dolerites. Australian Journal of Earth Sciences, 48, 543-549. HJELLE, A., OHATA, Y. & WINSNES, T.S. 1982. Geology and petrology of the southern Heritage Range, Ellsworth Mountains. In: CRADDOCK, C. (ed.) Antarctic Geoscience. University of Wisconsin Press, Madison, 599-608. HOCH, M. 1999. Geochemistry and petrology of ultramafic lamprophyres from Schirmacher Oasis, East Antarctica. Mineralogy and Petrology, 65, 51-67. HOCH, M. & TOBSCHALL, HJ. 1998. Minettes from Schirmacher Oasis, East Antarctica - indicators of an enriched mantle source. Antarctic Science, 10, 476-486. HOCH, M., REHKAMPER, M. & TOBSCALL, HJ. 2001. Sr, Nd, Pb and O isotopes of minettes from Schirmacher Oasis, East Antarctica: a case of mantle metasomatism involving subducted continental material. Journal of Petrology, 42,1387-1400. HOFMANN, J., KAISER, G, KLEMM, W. & PAECH, H.-J. 1980. K/Ar-alter von doleriten und metamorphiten der Shackleton Range und der Whichaway Nunataks, ost- und siidostumrandung des Filchner-eisschelfs (Antarktis). Zeitschrift fur Geologische Wissenschaften, 8, 1227-1232. HOKADA,T, MlSAWA, K., YOKOYAMA, K., SHIRAISHI, K.

& YAMAGUCHI, A. 2004. SHRIMP and electron microprobe chronology of UHT metamorphism in the Napier Complex, East Antarctica: implications for zircon growth at >1000°C. Contributions to Mineralogy and Petrology, 147, 1-20. HOLE, MJ. & LEMASURIER, W.E. 1994. Tectonic controls on the geochemical composition of Cenozoic, mafic alkaline volcanic rocks from West Antarctica. Contributions to Mineralogy and Petrology, 117, 187-202. HOLE, M.J., KEMPTON, P.D. & MILLAR, I.L. 1993. Trace-element and isotopic characteristics of small-degree melts of the asthenosphere: evidence from the alkali basalts of the Antarctic Peninsula. Chemical Geology, 109, 51-68. HOLE, M J., STOREY, B.C. & LEMASURIER, W.E. 1994. Tectonic setting and geochemistry of Miocene

alkalic basalts from the Jones Mountains, West Antarctica. Antarctic Science, 6, 85-92. HOTTEN, R. 1993. Die mafischen gange der Shackleton Range/Antarktika: petrography, geochemie, isotopen geochemie und palaomagnetik. Berichte zur Polarforschung, 118. HUNTER, M.A. & LOMAS, S.A. 2003. Reconstructing the Siluro-Devonian coastline of Gondwana: insights from the sedimentology of the Port Stephens Formation, Falkland Islands. Journal of the Geological Society, London, 160, 459^76. JACOBS, J., FANNING, C.M., HENJES-KUNST, K, OLESCH, M. & PAECH, H.-J. 1998. Continuation of the Mozambique belt into East Antarctica: Grenville-age metamorphism and polyphase Pan-African high-grade events in central Dronning Maud Land. Journal of Geology, 106, 385-406. JACOBS, I,THOMAS, R.J., ARMSTRONG, R.A. & HENJESKUNST, F. 1999. Age and thermal evolution of the Mesoproterozoic Cape Meredith Complex, West Falkland. Journal of the Geological Society, London, 156, 917-928. JACOBS, I, FANNING, C.M. & BAUER, W. 2003. Timing of Grenville-age vs. Pan-African medium- to high grade metamorphism in western Dronning Maud Land (East Antarctica) and significance for correlations in Rodina and Gondwana. Precambrian Research, 125, 1-20. JORDAN, T.H. 1988. Structure and formation of the continental tectosphere. In: MENZIES, M.A. & Cox, K.G. (eds) Oceanic and Continental Lithosphere: Similarities and Differences. Journal of Petrology Special Lithosphere Issue, 11-37. KYLE, PR. 1980. Development of heterogeneities in the subcontinental mantle: evidence from the Ferrar Group, Antarctica. Contributions to Mineralogy and Petrology, 73, 89-104. KYLE, PR., PANKHURST, RJ. & BOWMAN, J.R. 1983. Isotopic and chemical variations in Kirkpatrick Basalt Group rocks from Southern Victoria Land. In: OLIVER, R.L., JAMES, PR. & JAGO, J.B. (eds) Antarctic Earth Science. Australian Academy of Science, Canberra, 234-237. LEAT, P.T., THOMPSON, R.N., MORRISON, M.A., HENDRY, G.L. & DICKIN, A.P. 1988. Compositionallydiverse, Miocene-Recent rift-related magmatism in NW Colorado: partial melting and mixing of mafic magmas from 3 different asthenospheric and lithospheric mantle sources. In: MENZIES, M.A. & Cox, K.G. (eds) Oceanic and Continental Lithosphere: Similarities and Differences. Journal of Petrology Special Lithosphere Issue, 351-377. LEAT, P.T. STOREY, B.C. & PANKHURST, RJ. 1993. Geochemistry of Palaeozoic-Mesozoic Pacific rim erogenic magmatism, Thurston Island area, West Antarctica. Antarctic Science, 5, 281-296. LEAT, P.T, SCARROW, J.H. & MILLAR, I.L. 1995. On the Antarctic Peninsula batholith. Geological Magazine, 132, 399-412. LEAT, P.T, RILEY, T.R., STOREY, B.C., KELLEY, S.P. & MILLAR, I.L. 2000. Middle Jurassic ultramafic lamprophyre dyke within the Ferrar magmatic

ANTARCTIC LITHOSPHERIC DOMAINS province, Pensacola Mountains, Antarctica. Mineralogical Magazine, 64, 95-111. LEAT, P.T., RILEY,T.R., WAREHAM, CD., MILLAR, I.L., KELLEY, S.P. & STOREY, B.C. 2002. Tectonic setting of primitive magmas in volcanic arcs: an example from the Antarctic Peninsula. Journal of the Geological Society, London, 159, 31-44. LEEMAN, W.P 1982. Tectonic and magmatic significance of strontium isotopic variation in Cenozoic volcanic rocks from the western United States. Geological Society of America Bulletin, 93, 487-503. LUTTINEN, A.V. & FURNES, H. 2000. Flood basalts of Vestfjella: Jurassic magmatism across an Archaean-Proterozoic lithospheric boundary in Dronning Maud Land, Antarctica. Journal of Petrology, 41, 1271-1305. LUTTINEN, A.V., ZHANG, X. & POLAND, K.A. 2002. 159 Ma Kjakebeinet lamproites (Dronning Maud Land, Antarctica) and their implications for Gondwana break-up processes. Geological Magazine, 139, 525-539. MARSHALL, J.E.A. 1994. The Falkland Islands: a key element in Gondwana paleogeography. Tectonics, 13, 499-514. MCKENZIE, D. 1989. Some remarks on the movement of small melt fractions in the mantle. Earth and Planetary Science Letters, 95, 53-72. MENZIES, M.A. 1989. Cratonic, circumcratonic and oceanic mantle domains beneath the western United States. Journal of Geophysical Research, 94, 7899-7915. MENZIES, M.A. 1992. The lower lithosphere as a major source for continetal flood basalts: a re-appraisal. In: STOREY, B.C., ALABASTER, T. & PANKHURST, R.J. (eds) Magmatism and the Causes of Continental Break-up. Geological Society, London, Special Publications, 68, 31-39. MENZIES, M.A. & HAWKESWORTH, CJ. (eds) 1987. Mantle Metasomatism. Academic Press, London. MENZIES, M. A., HALLIDAY, A.N., PALACZ, Z., HUNTER, R.H., UPTON, B.G.J., ASPEN, P. & HAWKESWORTH, CJ. 1987. Evidence from mantle xenoliths for an enriched lithospheric keel under the Outer Hebrides. Nature, 325, 44-47. MENZIES, M.A., FAN, W. & ZHANG, M. 1993. Palaeozoic and Cenozoic lithoprobes and the loss of >120 km of Archaean lithosphere, Sino-Korean craton, China. In: PRICHARD, H.M., ALABASTER, T, HARRIS, N.B.W. & NEARY, C.R. (eds) Magmatic Processes and Plate Tectonics. Geological Society, London, Special Publications, 76, 71-81. MIKHALSKY, E.V. & SHERATON, J.W. 1993. Association of dolerite and lamprophyre dykes, Jetty Peninsula (Prince Charles Mountains, East Antarctica). Antarctic Science, 5, 297-307. MIKHALSKY, E.V., LAIBA, A.A., BELYATSKY, B.V., SOSEDKO, T.A. & ANDRONIKOV, A.V. 1994. Lamproites from the Rubin Massif, Prince Charles Mountains, East Antarctica. Petrology, 2, 258-264. MIKHALSKY, E.V., SHERATON, J.W., LAIBA, A.A., TINGEY, R.J., THOST, D.E., KAMENEV, E.N. &

377

FEDEROV, L.V. 2001. Geology of the Prince Charles Mountains, Antarctica. AGSO Geoscience Australia Bulletins 247. MILLAR, I.L. & PANKHURST, R.J. 1987. Rb-Sr geochronology of the region between the Antarctic Peninsula and the Transantarctic Mountains: Haag Nunataks and Mesozoic granitoids. In: MCKENZIE, G.D. (ed.) Gondwana Six: Structure, Tectonics and Geophysics. American Geophysical Union, Monographs, 40, 151-160. MILLAR, I.L., WILLAN, R.C.R., WAREHAM, CD. & BOYCE, A.I 2001. The role of crustal and mantle sources in the genesis of granitoids of the Antarctic Peninsula and adjacent crustal blocks. Journal of the Geological Society, London, 158, 855-867. MILLAR, I.L., PANKHURST, R.J. & FANNING, C.M. 2002. Basement chronology of the Antarctic Peninsula: recurrent magmatism and anatexis in the Palaeozoic Gondwana margin. Journal of the Geological Society, London, 159,145-157. MITCHELL, R.H., SMITH, C.B. & VLADYKIN, N.V. 1994. Isotopic composition of strontium and neodymium in potassic rocks of the Little Murun complex, Aldan Shield, Siberia, Lithos, 32, 243-248. MOYES, A.B., BARTON, J.M. & GROENEWALD, P.B. 1993. Late Proterozoic to Early Palaeozoic tectonism in Dronning Maud Land, Antarctica: supercontinent fragmentation and amalgamation. Journal of the Geological Society, London, 150, 833-842. MUKASA, S.B. & DALZIEL, I.W.D. 2000. Marie Byrd Land, West Antarctica: evolution of Gondwana's Pacific margin constrained by zircon U-Pb geochronology and feldspar common-Pb isotopic compositions. Geological Society of America Bulletin, 112, 611-627. MURPHY, D.T., COLLERSON, K.D. & KAMBER, B.S. 2002. Lamproites from Gaussberg, Antarctica: possible transition zone melts of Archaean sediments. Journal of Petrology, 43, 981-1001. NELSON, D.R., MCCULLOCH, M.T. & SUN, S.-S. 1986. The origins of ultrapotassic rocks as inferred from Sr, Nd and Pb isotopes. Geochimica et Cosmochimica Acta, 50, 231-245. PANKHURST, R.J., MILLAR, I.L., GRUNOW, A.M. & STOREY, B.C. 1993. The pre-Cenozoic magmatic history of the Thurston Island crustal block, West Antarctica. Journal of Geophysical Research, 98, 11835-11849. PANKHURST, R.J., WEAVER, S.D., BRADSHAW, J.D., STOREY, B.C. & IRELAND, T.R. 1998. Geochronology and geochemistry of pre-Jurassic superterranes in Marie Byrd Land, Antarctica. Journal of Geophysical Research, 103, 2529-2547. PANTER, K.S., KYLE, PR. & SMELLIE, J.L. 1997. Petrogenesis of a phonolite-trachyte succession at Mount Sidley, Marie Byrd Land, Antarctica. Journal of Petrology, 38, 1225-1253. PANTER, K.S., HART, S.R., KYLE, P., BLUSZTANJN, J. & WILCH, T. 2000. Geochemistry of Late Cenozoic basalts from the Crary Mountains: characterization of mantle sources in Marie Byrd Land, Antarctica. Chemical Geology, 165, 215-141.

378

P.T. LEATETAL.

PARSONS, B. & SCLATER, J.G. 1977. An analysis of the variation of ocean floor bathymetry and heat flow with age. Journal of Geophysical Research, 82, 803-827. PEARSON, D.G. & NOWELL, G.M. 2002. The continental lithospheric mantle: characteristics and significance as a mantle reservoir. Philosophical Transactions of the Royal Society of London, Series A, 360, 2383-2410. PEARSON, D.G., CANIL, D.C. & SHIREY, S.B. 2003. Mantle samples included in volcanic rocks: xenoliths and diamonds. In: CARLSON, R.W. (ed.) Treatise on Geochemistry, vol. 2, Chapter 5. Elsevier, Amsterdam, 171-275. POLLACK, H.N., HURTER, S.J. & JOHNSON, J.R. 1993. Heat flow from the Earth's interior: analysis of the global data set. Reviews of Geophysics, 31, 267-280. RANDALL, D.E. & MACNIOCAILL, C. 2004. Cambrian palaeomagnetic data confirm a Natal embayment location for the Ellsworth-Whitmore Mountains, Antarctica, in Gondwana reconstructions. Geophysical Journal International, 157, 105-116. RANDALL, D.E., CURTIS, M.L. & MILLAR, I.L. 2000. A new late Middle Cambrian palaeomagnetic pole for the Ellsworth Mountains, Antarctica. Journal of Geology, 108, 403-425. RENNE, PR., SWISHER, C.C., DEINO, A.L., KARNER, D.B., OWENS, T.L. & DEPAOLO, DJ. 1998. Intercalibration of standards, absolute ages and uncertainties in Ar-Ar dating. Chemical Geology, 145, 117-152. REX, D.C. 1972. K-Ar age determinations on volcanic and associated rocks from the Antarctic Peninsula and Dronning Maud Land. In: ADIE, R.J. (ed.) Antarctic Geology and Geophysics. Universitetsforlaget, Oslo, 133-136. RILEY, T.R. & KNIGHT, K.B. 2001. Age of pre-breakup Gondwana magmatism. Antarctic Science, 13, 99-110. RILEY, T.R., LEAT, P.T., STOREY, B.C., PARKINSON, I.J. & MILLAR, I.L. 2003. Ultramafic lamprophyres of the Ferrar large igneous province: evidence for a HIMU mantle component. Lithos, 66, 63-76. RILEY, T.R., LEAT, P.T., CURTIS, M.L., MILLAR, I.L. & FAZEL, A. 2005. Early-Middle Jurassic dolerite dykes from western Dronning Maud Land (Antarctica): identifying mantle sources in the Karoo large igneous province. Journal of Petrology, in press. RITZWOLLER, M.H., SHAPIRO, N.M., LEVSHIN, A.L. & LEAHY, G.M. 2001. Crustal and upper mantle structure beneath Antarctica and surrounding oceans. Journal of Geophysical Research, 106, 30 645-30 670. ROCCHI, S., ARMIENTI, P., D'ORAZIO, M., TONARINI, S., WIJBRANS, J.R. & Di VIENCENZO, G. 2002. Cenozoic magmatism in the western Ross embayment: role of mantle plume versus plate dynamics in the development of the West Antarctic rift system. Journal of Geophysical Research, 107 (B9), ECV5, 1-22, 2195, 10.1029/2001JB000515. ROCHOLL, A., STEIN, M., MOLZAHN, M., HART, S.R. & WORNER, G. 1995. Geochemical evolution of rift

magmas by progressive tapping of a stratified source beneath the Ross Sea rift, Northern Victoria Land, Antarctica. Earth and Planetary Science Letters, 131, 207-224. RUDNICK, R.L. & FOUNTAIN, DM. 1995. Nature and composition of the continental crust: a lower crustal perspective. Reviews of Geophysics, 33, 267-309. SAUNDERS, A.D., STOREY, M., KENT, R.W. & NORRY, M.J. 1992. Consequences of plume-lithosphere interactions. In: STOREY, B.C., ALABASTER, T. & PANKHURST, R.J. (eds) Magmatism and the Causes of Continental Break-up. Geological Society, London, Special Publications, 68, 41-60. SCARROW, J.H., LEAT, P.T., WAREHAM, CD. & MILLAR, I.L. 1998. Geochemistry of mafic dykes in the Antarctic Peninsula continental-margin batholith: a record of arc evolution. Contributions to Mineralogy and Petrology, 131, 289-305. SHAPIRO, N.M. & RITZWOLLER, M.H. 2004. Inferring surface heat flux distributions guided by a global seismic model: particular application to Antarctica. Earth and Planetary Science Letters, 223, 213-224. SHERATON, J.W. & CUNDARI, A. 1980. Leucitites from Gaussberg, Antarctica. Contributions to Mineralogy and Petrology, 71, 417-427. SHERATON, J.W. & ENGLAND, R.N. 1980. Highly potassic mafic rocks from Antarctica. Journal of the Geological Society of Australia, 27, 129-135. SHERATON, J.W, TINGEY, R.J., BLACK, L.P. & OLIVER, R.L. 1993. Geology of the Bunger Hills area, Antarctica: implications for Gondwana correlations. Antarctic Science, 5, 85-102. SPAETH, G, HOTTEN, R., PETERS, M. & TECHMER, K. 1993. Mafic dykes in the Shackleton Range, Antarctica. Polarforschung, 63, 101-121. STOREY, B.C. 1991. The crustal blocks of West Antarctica within Gondwana: reconstruction and breakup model. In: THOMSON, M.R.A., CRAME, J.A. & THOMSON, J.W. (eds) Geological Evolution of Antarctica. Cambridge University Press, Cambridge, 587-592. STOREY, B.C. & GARRETT, S.W 1985. Crustal growth of the Antarctic Peninsula by accretion, magmatism and extension. Geological Magazine, 122, 5-14. STOREY, B.C., DALZIEL, I.W.D., GARRETT, S.W, GRUNOW, A.M., PANKHURST, R.J. & VENNUM, W.R. 1988. West Antarctica in Gondwanaland: crustal blocks, reconstruction and break-up processes. Tectonophysics, 155, 381-390. STOREY, B.C., ALABASTER, T., HOLE, M. J., PANKHURST, R.J. & WEVER, H.E. 1992. Role of subductionplate boundary forces during the initial stages of Gondwana break-up: evidence from the protoPacific margin of Antarctica. In: STOREY, B.C., ALABASTER, T. & PANKHURST, R.J. (eds) Magmatism and the Causes of Continental Break-up. Geological Society, London, Special Publications, 68, 149-163. STOREY, B.C., CURTIS, M.L., FERRIS, J.K., HUNTER, M.A. & LIVERMORE, R.A. 19990. Reconstruction

ANTARCTIC LITHOSPHERIC DOMAINS and break-out model for the Falkland Islands within Gondwana. Journal of African Earth Sciences, 29, 153-163. STOREY, B.C., LEAT, P.T., WEAVER, S.D., PANKHURST, R.J., BRADSHAW, J.D. & KELLEY, S. 19996. Mantle plumes and Antarctica-New Zealand rifting: evidence from mid-Cretaceous mafic dykes. Journal of the Geological Society, London, 156, 659-671. STOREY, B.C., LEAT, P.T. & FERRIS, J.K. 2001. The location of mantle plume centres during the initial stages of Gondwana break-up. In: ERNST, R.E. & BUCHAN, K.L. (eds) Mantle Plumes: Their Identification Through Time. Geological Society of America Special Papers, 352, 71-80. STOREY, M., SAUNDERS, A.D., TARNEY, J. ET AL. 1988. Geochemical evidence for plume-mantle interactions beneath Kerguelen and Heard islands, Indian Ocean. Nature, 336, 371-374. STUMP, E. 1995. The Ross Orogen of the Transantarctic Mountains. Cambridge University Press, Cambridge. SWEENEY, R.J., DUNCAN, A.R. & ERLANK, A.J. 1994. Geochemistry and petrogenesis of central Lebombo basalts of the Karoo igneous province. Journal of Petrology, 35, 95-125. TALARICO, F, KLEINSCHMIDT, G. & HENJES-KUNST, F 1999. An ophiolitic complex in the northern Shackleton Range, Antarctica. Terra Antartica, 6, 293-315. TECHMER, K.S., PETERS, M., SPAETH, G, WEBER, K. & LEAT, P.T. 1995. Mafic Dykes. In: CLARKSON, P.D., TESSENSOHN, F. & THOMSON, J.W. and 24 others Geological Map of Shackleton Range, Antarctica. BAS GEOMAP Series, Sheet 4, supplementary text. British Antarctic Survey, Cambridge, 48-52 (with appendix, 73-76). TESSENSOHN, F. & HENJES-KUNST, F. 2005. Northern Victoria Land terranes, Antarctica: far-travelled or local products? In: VAUGHAN, A.P.M., LEAT, P.T. & PANKHURST, RJ. (eds) Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 275-291. TESSENSOHN, F. & WORNER, G. 1991. The Ross Sea rift system, Antarctica: structure, evolution and analogues. In: THOMSON, M.R.A., CRAME, J.A. & THOMSON, J.W (eds) Geological Evolution of Antarctica. Cambridge University Press, Cambridge, 273-277. TESSENSOHN, F, KLEINSCHMIDT, G, TALARICO, F. ET AL. 1999. Ross-age amalgamation of East and West Gondwana: evidence from the Shackleton Range, Antarctica. Terra Antartica, 6, 317-325. THISTLEWOOD, L., LEAT, P.T, MILLAR, I.L., STOREY, B.C. & VAUGHAN, A.PM. 1997. Basement geology and Palaeozoic-Mesozoic mafic dykes from the Cape Meredith complex, Falkland Islands: a record of repeated intracontinental extension. Geological Magazine, 134, 335-367. THOMAS, R.J., HENJES-KUNST, F & JACOBS, J. 1998. Pre-lamprophyre mafic dykes of the Cape Meredith complex, West Falkland. Geological Magazine, 135, 495-500.

379

THOMPSON, R.N., LEAT, P.T, DICKIN, A.P., MORRISON, M.A., HENDRY, G.L. & GIBSON, S.A. 1989. Strongly potassic mafic magmas from lithospheric mantle sources during continental extension and heating: evidence from Miocene minettes of northwest Colorado, U.S.A. Earth and Planetary Science Letters, 98, 139-153. TlNGEY, RJ., MCDOUGALL, I. & GLEADOW, A.J.W

1983. The age and mode of formation of Gaussberg, Antarctica. Journal of the Geological Society of Australia, 30, 241-246. VAUGHAN, A.PM. & STOREY, B.C. 2000. The eastern Palmer Land shear zone: a new terrane accretion model for the Mesozoic development of the Antarctic Peninsula. Journal of the Geological Society, London, 157, 1243-1256. VAUGHAN, A.PM., KELLEY, S.P. & STOREY, B.C. 20020. Mid-Cretaceous ductile deformation on the eastern Palmer Land shear zone, Antarctica, and implications for timing of Mesozoic terrane collision. Geological Magazine, 139, 465-471. VAUGHAN, A.P.M., PANKHURST, RJ. & FANNING, C.M. 2000i>. A mid-Cretaceous age for the Palmer Land event, Antarctic Peninsula: implications for terrane accretion timing and Gondwana palaeolatitudes. Journal of the Geological Society, London, 159, 113-116. VENNUM, W.R., GIZYCKI, P., SAMSONOV, V.V., MARKOVICH, A.G. & PANKHURST, RJ. 1992. Igneous petrology and geochemistry of the southern Heritage Range, Ellsworth Mountains, West Antarctica. In: WEBERS, G.F., CRADDOCK, C. & SPLETTSTOESSER, J.F. (eds) Geology and Paleontology of the Ellsworth Mountains, West Antarctica. Geological Society of America Memoirs, 170, 295-324. WEAVER, S.D., STOREY, B.C., PANKHURST, RJ., MUKASA, S.B., DlVENERE, VJ. & BRADSHAW, J.D.

1994. Antarctica-New Zealand rifting and Marie Byrd Land lithospheric magmatism linked to ridge subduction and mantle plume activity. Geology, 22, 811-814. WHITE, C.M. & CRADDOCK, C. 1987. Compositions of igneous rocks in the Thurston Island area, Antarctica: evidence for a Late PaleozoicMiddle Mesozoic andinotype continental margin. Journal of Geology, 95, 699-709. WILLIAMS, R.W, COLLERSON, K.D., GILL, IB. & DENIEL, C. 1992. High Th/U ratios in subcontinental lithospheric mantle: mass spectrometric measurement of Th isotopes in Gaussberg lamproites. Earth and Planetary Science Letters, 111, 257-268. WILSON, M. & DOWNES, H. 1991. Tertiary-Quaternary extension-related alkaline magmatism in western and central Europe. Journal of Petrology, 32, 811-849. WINBERRY, IP. & ANANDAKRISHNAN, S. 2004. Crustal structure of the West Antarctic rift system and Marie Byrd Land hotspot. Geology, 32, 977-980. WORNER, G. 1999. Lithospheric dynamics and mantle sources of alkaline magmatism of the Cenozoic West Antarctica rift system. Global and Planetary Change, 23, 61-77.

380

P.T.LEATETAL.

Wu, B. & BERG, J.H. 1992. Early Paleozoic lamprophyre dikes of Southern Victoria Land: geology, petrology and geochemistry. In: YOSHIDA, Y., KAMINUMA, K. & SHIRAISHI, K. (eds) Recent Progress in Antarctic Earth Science. Terrapub, Tokyo, 257-264. ZHANG, H.-R, SUN, M., ZHOU, M.-R, PAN, W.-M., ZHOU, X.-H. & ZHAI, M.-G. 2004. Highly heterogeneous late Mesozoic lithospheric mantle beneath the North China craton: evidence from Sr-Nd-Pb isotopic systematics of mafic igneous rocks. Geological Magazine, 141, 55-62.

ZHANG, X., LUTTINEN, A.V., ELLIOT, D.H., LARSSON, K. & ROLAND, K.A. 2003. Early stages of Gondwana breakup: the 40Ar/39Ar 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 (B9), ECV7 1-19, 2449, 1029/2001JB001070. ZINDLER, A. & HART, S. 1986. Chemical geodynamics. Annual Reviews of Earth and Planetary Science, 14, 493-571.

Provenance studies of very low- to low-grade metasedimentary rocks of the Puncoviscana complex, northwest Argentina UDO ZIMMERMANN Department of Geology, University of Johannesburg, PO Box 524, Auckland Park 2006, Johannesburg, South Africa (e-mail: [email protected]) Abstract: A provenance study of Neoproterozoic to Lower Cambrian rocks for the entire Puncoviscana Basin was conducted, using 119 samples from 15 different outcrops. Petrographic data (Qt60-80, F15-35, L5-20, P/F 0.2-0.4, LV/L = 0) show a composition comparable to foreland-basin successions. Lithoclasts are of metamorphic and metasedimentary origin. Volcanic debris is detected only in the form of sanidine, and volcanic lithoclasts were probably decomposed to form pseudo-matrix. Framework clasts are sub-angular to sub-rounded, and the rocks are poorly sorted. Major element geochemistry shows a moderate to high Chemical Index of Alteration (56-77) and failed to provide coherent provenance and rock classification. Trace element geochemistry suggests a rhyodacitic composition overall. Rare earth element patterns are comparable to those of model upper continental crust (UCC), as are concentrations of Nb, Ta, Ti, Th-Sc and Eu/Eu* (0.45-0.87; 95% between 0.4 and 0.7); reworking signatures are not detected. The uniform mineralogical and geochemical composition reflects supra-crustal source(s) for the entire basin, including significant metamorphic rock debris. The Puncoviscana complex is interpreted as a peripheral Pampean foreland basin, fed mainly from an eastern fold-thrust belt, but includes relicts of pre- and syn-collisional magmatic activity as well. A source area of UCC composition to the west is represented by the Arequipa block.

For the last 20 years, the palaeotectonic evolution of the western border of Gondwana during the Neoproterozoic and Palaeozoic has been the object of controversy regarding the basic question, 'is crustal growth related to terrane accretion or to recycling of existing crustal material?' (e.g. Ramos 1986; Bahlburg 1990; Rapela et al 1992; Astini et al 1995; Bahlburg & Herve 1997; Pankhurst & Rapela 1998; Bock et al 2000; Lucassen et al 2000; Zimmermann & Bahlburg 2003). One key element to understanding crustal evolution in this area is the Upper Vendian to Lower Cambrian Puncoviscana Formation. Turner (1960) described pre-Ordovician rock successions from northwestern Argentina (Fig. 1) comprising metagreywackes and psammites, with few intercalations of metacarbonates, but dominated by pelites, which he named the Puncoviscana Formation. Since then, other very low- to low-grade metasedimentary rocks of Late Vendian to Early Cambrian age in the mentioned area have been included in the classification as Puncoviscana Formation, or Puncoviscana Formation s.l (Salfity et al 1975; Toselli & Acenolaza 1978; Acenolaza & Toselli 1981; Omarini 1983; Baldis & Omarini 1984; Acenolaza et al 1988; 1990). Medium- to highgrade metamorphic rocks (Fig. 1), mainly interpreted as metasedimentary, are associated with

the low-grade metamorphosed siliciclastic rocks (e.g. Caminos 1979). This wider Puncoviscana Formation will be referred to here as the 'Puncoviscana complex' (following F. G. Acenolaza, pers. comm. 2004), constituting different, mappable, lithologies that comprise the basin infill, such as conglomerates, sandstones, siltstones, shales and carbonates, all of which were overprinted by very lowto low-grade metamorphism. Classification is based on the nomenclature of the British Geological Survey for metamorphic and sedimentary rocks (Hallsworth & Knox 1999; Robertson 1999). Low-grade metamorphism does not necessarily mask the general provenance information obtained from combined petrographical and geochemical information (Floyd & Leveridge 1987; McLennan et al 1990; 1993; Floyd et al 1991). Detailed provenance studies have been carried out within rock successions at comparable metamorphic grades and in the same regional setting (e.g. Toselli & Weber 1982; Bahlburg 1990; 1998; Zimmermann 1999; Bock et al 2000; Zimmermann et al 2002; Zimmermann & Bahlburg 2003) or elsewhere (e.g. Bock et al 1996; Floyd et al 1989; Thoulkeridis et al 1999). This contribution presents the first dataset and integrated provenance study for the entire Puncoviscana complex in northwestern Argentina in order to

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 381-416. 0305-8719/$15.00 © The Geological Society of London 2005.

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develop a regional model for the depositional setting and the composition of the source rocks. New petrographical data are combined with earlier published results and supported by new geochemical data covering the whole basin, and are interpreted according to modern aspects of provenance analysis (McLennan et al 1990; 1993; Fedo et al 1995; Nesbitt et al 1996; Nesbitt & Markovics 1997; Bahlburg & Floyd 1999). Finally, the resulting data are interpreted on the base of existing models for the evolution of the Puncoviscana Basin. As in all former publications, the rocks are interpreted as deposited in the same basin under a similar palaeotectonic regime during Late Vendian to Early Cambrian time. Future aims will be the radiometric dating of detrital zircons, the development of a complete lithostratigraphy and, finally, comparison with rocks interpreted as Puncoviscana equivalents from other areas in Argentina.

Geological setting The proposed Puncoviscana Basin extends from Bolivia to central Argentina, as its western and eastern boundaries are not well defined, but have been assumed to include the present outcrop areas (Figs 1, 2a). The basin can be divided into two main regions based on the occurrence of different metamorphic rocks (Acenolaza et al 2000). In the northern region very low- to low-grade metasedimentary rocks predominate, while the southern area is characterized by medium- to high-grade metamorphic rocks including gneisses and migmatites (Fig. 1). Contacts between the Puncoviscana complex and the associated gneisses, migmatites and medium- to high-grade metamorphic rocks are not exposed (Caminos 1979). Some authors (WiUner & Miller 1986; Acenolaza et al 1988) argue that the higher deformed and metamorphic rocks represent deeper crustal equivalents of the supra-crustal succession (Puncoviscana complex), whereas Mon & Hongn (1991) interpret the variety of metamorphic rocks as related to different tectono-metamorphic events, and imply an older age for the higher-grade rocks. However, both successions are affected by Cambrian to Early Ordovician felsic to intermediate plutonic activity, a strong Early

383

Fig. 2. (a) The probable terrane distribution in central South America during the Early Cambrian. The Puncoviscana Basin is located on the hypothetical composite Pampia-Arequipa block. The terrane extension to the south is unknown (after Keppie & Bahlburg 1999). (b) Sketch of different tectonic units in northwestern Argentina. 1, Precordillera; 1', inferred southern extension; 2, western Sierras Pampeanas; 2', inferred southern extension; 3, Puna; 4, Famatina; 5, Central batholitic belt; 6, Cordillera Oriental; 7, eastern Sierras Pampeanas; 7', inferred southern extension (after Pankhurst & Rapela 1998).

Cambrian polyphase deformation (Bachmann et al 1987; Mon & Hongn 1991; Rapela et al 1992; Hongn et al 1996; Llambias et al 1998; Rapela et al 1998; Fig. 1) and a Cambrian regional metamorphism (Rapela et al 1992; Lucassen et al 2000; 2002). The Early Cambrian deformation distinguishes the Puncoviscana complex from younger magmato-sedimentary successions (e.g. Mon & Hongn 1991). The Puncoviscana complex consists of sandto gravel-sized channel infills, which cut into dominantly silty and clayey semi-pelites and pelites deposited in a mid- to outer-fan environment (Jezek 1990). According to Jezek & Miller (1986) the main sediment transportation direction was towards the west and northwest. Jezek (1990) presented a sedimentation model for the Puncoviscana complex focusing on outcrops in the eastern part of the basin (Cordillera Oriental and Ambato, the region between Jujuy and Tucuman in Fig. 1). Sedimentation started with sand-rich turbidites of middle-fan channel and lobe facies, a second phase was characterized by marginal middle- to

Fig. 1. Geological map of the Puncoviscana Basin showing the sampled outcrops. Circles indicate sampling for geochemical analyses and squares for quantitative petrography (map redrawn after Acenolaza et al. 2000). The sampling areas are identified as follows: Puna: 1, Rio Taique (RT); 2, Quebrada del Volcan (VOL); 3, Quebrada Randolfo (RAN). Cordillera Oriental: 4, Purmamarca (TOR); 4', Quebrada El Toro and Munano (TOR); 5, Corralito (COR); 6, La Pedrera (FED); 7, Seclantes (SEC); 8, Molinos (MOL); 9, Choromoro (CHO). Ambato: 10, Quebrada de Suncho (SUN); 11, Quebrada La Cebila (CEB); 12, Sijan - clayey pelites (SIJ-P) and medium- to coarse sand psammites (SIJ-G); 13, Conception (CON). Famatina: 14, Los Corrales (NP); 15, Quebrada de Paiman (LA)).

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U. ZIMMERMANN

outer-fan regions associated with hemipelagic to pelagic red pelites, and a third stage exhibits slide masses, debris-flow deposits and common water-escape structures indicating a rise in sedimentation rates and tectonic segmentation of the basin. Finds of trace fossils point to a possible age of deposition as latest Vendian to Early Cambrian (e.g. Durand & Acenolaza 1990) and also give rise to speculation about the palaeogeography of the basin (Acenolaza & Acenolaza 2002; Buatois & Mangano 2003). The interpreted age range coincides with results from U-Pb dating of detrital zircon; lower intercepts of 530-560 Ma were obtained for zircons with a rhyolitic source, and upper intercepts of 1.7-1.8 Ga from rounded zircons (Lork et al 1990). However, Do Campo et al. (1999) presented K-Ar data on authigenic mica reflecting a diagenetic age of 670 ± 27 Ma. A minimum age for the Puncoviscana complex is given by Early to Middle Cambrian calcalkaline and peraluminous granites intruding the basin deposits syn- and post-tectonically during the Pampean Orogeny (Rapela et al. 1992; Bachmann et al. 1987). To the southeast of the basin, in the southern eastern Sierras Pampeanas (Fig. 2b), igneous activity at 560-520 Ma points to arc-collisional characteristics (Pampean arc; Rapela et al. 1998; Stuart-Smith et al. 1999; von Gosen et al. 2002). The plutonic rocks in this area intrude low-grade metasedimentary and highgrade metamorphic rocks. Some outcrops of metasedimentary rocks have been the subject of geochemical analysis (Lopez de Luchi et al. 2003) and they have been interpreted as a southeastern extension of the Puncoviscana complex (Prozzi 1990; Sollner et al 2000). They have been studied partly in terms of zircon age dating by ion-microprobe (Schwartz & Gromet 2004). Detailed studies of mineralogical, sedimentological, geochemical and facies aspects are at a preliminary stage and a comprehensive provenance study of these deposits is in preparation. Geochemical analyses of the Puncoviscana complex are sparse, concentrated on outcrops in the eastern basin area and mainly only for major elements (Willner et al 1985). Trace element data are presented only from a few small outcrops areas (Rossi et al 1997; Bock et al 2000; Do Campo & Guevara 2002) and are insufficient for regional interpretations. However, Willner et al. (1985), Rossi et al (1997) and Do Campo & Guevara (2002) interpret the Puncoviscana complex as deposited at the passive margin of Gondwana based on their data. Interestingly, the sparse Nd-isotope data of the Puncoviscana complex (e.g. Bock et al. 2000, n = 6) are homogeneous, with Nd-model (TDM) ages of about 1.3-1.7 Ga. The medium-

to high-grade metasedimentary rocks associated with the Puncoviscana complex in the northwestern region of Argentina (Puna and Cordillera Oriental, Fig. 2b) yield slightly older Nd-model ages of 1.5-2.1 Ga, 80% scattering around 1.6 Ga (Lucassen et al 2000; n = 22). Unconformably overlying the Puncoviscana complex after the complex deformation, the mainly quartz arenites of the Middle to Upper Cambrian Meson Group, are interpreted as deposited in a shallow-marine system on a passive margin (Acenolaza et al 1982; Kumpa & Sanchez 1988). During the Early Ordovician an active continental margin along the western border of Gondwana was developed (Bahlburg 1990; Pankhurst et al 1998; Zimmermann & Bahlburg 2003) without the addition of allochthonous crustal material (Bahlburg 1998; Bock et al 2000; Zimmermann & Bahlburg 2003). Controversies concerning the evolution of the Puncoviscana complex are due mainly to: (i) lack of a complete lithostratigraphy; (ii) unresolved geological relationships of the low-grade metasedimentary and the medium- to highgrade metamorphic rocks, as no contact between the two successions has been found so far; and (iii) incomplete provenance data for the entire basin, leading to contradictory palaeotectonic models based on restricted datasets. Some authors (Acenolaza et al. 1983; Willner et al. 1985; Jezek & Miller 1986; Rossi et al. 1997; Do Campo & Guevara 2002) interpret the Puncoviscana complex as a typical passive continental margin deposit following an older rifting process, probably with Laurentia as a counterpart (e.g. Rapela et al. 1992). Omarini et al. (1999), using geochemical data for volcanic rocks of uncertain stratigraphic position, suggested that the Puncoviscana complex records progression from a rift environment to a back-arc setting. Kraemer et al (1995) and Keppie & Bahlburg (1999) re-interpret the same succession as foreland-basin infill, syntectonic with the collision of the Pampia block (a microcontinent or microplate, Rapela et al 1998) with the western border of the Rio de la Plata craton (including the Chaco block) accompanied by subduction towards the east (Rapela et al. 1998). During the Early to Middle Cambrian, a volcanic arc (the Pampean arc, exposed in the Sierras de Cordoba, Rapela et al 1998; 2001) represented the western border of the Rio de la Plata craton (Ramos et al 1993, Astini et al. 1996; Rapela & Pankhurst 1996; Rapela et al. 1998). A successful approach deciphering the crustal evolution of the Puncoviscana Basin should comprise a modern provenance study based on geochemical and petrographical data of the basin infill.

385

THE PUNCOVISCANA COMPLEX Table 1. Comparative petrographical data from the Puncoviscana complex

Puncoviscana complex* (38 samples) Puncoviscana complex (24 samples) Rifted margin Biscay1 Guinea1 South Korea1 Puna2 Forearc basin Marianes1 Japan1 Continental arc Middle America1 Java1 Puna 2 ' 3 Peru-Chile1 Foreland basin Madre de Dios4 Puna 3 ' 5 Puna3 El Imperial6 Marnoso7 Continent collision Ganges1 Ganges1

Age

Qt

F

L

P-F

LV/L

NeoproterozoicLower Cambrian NeoproterozoicLower Cambrian

64.8 (55-89) 60-80

11.3 (4-38) 15-35

23.9 (8-50) 5-20

0.65

0

0.2-0.4

0

modern modern modern Tremadoc

86.8 73.1 47.7

9.4 26.1 49.8

3.8 0.7 2.5 4

0.44 0.34 0.36

0.06

94.1 77.9

1 0.9

0.99

15.3

0.78 0.47 0.81 0.79

0.99

0.2-0.4 0.22 0.72

0-0.2

89

7

modern modern

0 3.8

18.3

modern modern Arenig modern

10.1 54.8

74.6 35.4

49 33.3

17 49

74 78 67

5 12 8

21 9 25

85.8

3 28

12.2

modern Llanvirn Llandeilo Mid CarboniferousLower Permian Miocene Neogene Neogene

54 54.4

45

5.9

35.6 32.3

9.8 34 17.7

18 10 22.6

<0.5 0.37 0.45

1

0.98

1 0 0.62 0-0.2 0-0.2

0.01

Qt, total quartz; F, feldspar; L = lithoclasts; P, plagioclase; Lv, volcanic lithoclasts. * Published data (38 samples from the eastern part of the basin) from Jezek (1990). t New data this publication (24 samples from SIJ-G, CON, SUN, CHO, SEC, MOL, TOR and RT; see Fig. 1). Data for comparison from; l McLennan et al (1990);2 Zimmermann & Bahlburg (2003);3 Bahlburg (1990); 4 DeCelles & Hertel (1989);5 Zimmermann et al. (2002); 6 Espejo & Lopez-Gamundi (1994); 7 Zuffa et al. (1980).

Analytical methods Framework mineral composition was quantified using the point-counting method of Gazzi & Dickinson, as described by Ingersoll etal (1984), with a Swift Model F counter, taking 400 to 550 counts in traverses. The framework mineral proportions obtained by point-counting are presented in Table 1, in which comparative data from different tectonic settings are also listed. X-ray diffraction (XRD) was carried out at SPEKTRAU (University of Johannesburg) on 150 samples to determine the mineralogical composition. A Phillips PW 1710 X-ray diffractometer (controlled by Phillips ADP 3.6) was operated at 40 mA and 40 kV using a Cu anode. Scan mode was continuous on a spinning sample and patterns were measured from 3° to 80° 26 (step size 0.017°, scan step-time 3.81 s). Scanning electron microscope (SEM), backscattered electron microscope (BSE), energy dispersive system (EDS) and cathodo-

luminescence (CL) studies were carried out on a JEOL JSM-5600 at SPECTRAU, on polished thin sections covered with carbon using a BIORAD from Poleron Division to analyse textures. The main purpose of these analyses was to control probable recrystallization and new growth of mineral phases, and to determine feldspar compositions. Microprobe (MP) data were processed on the same polished and covered thin-sections on a CAMEBAX, to determine feldspar and mica. The samples for geochemical analyses were cleaned, and weathered and veined surfaces were cut off. The rocks were crushed and milled to a very fine powder and fused with LiBO2. Major and trace element analyses were made by ICP-MS at ACME Laboratories (Vancouver, Canada). Summary data, averaged for the different sampling areas, are presented in Table 2 while discrimination parameters for arc provenance are presented in Table 3. Full data are given in Table 4.

U. ZIMMERMANN

386

Table 2. Average values and standard deviation of selected elements and element ratios for different sampling areas of the Puncoviscana complex PUNA UCC

SiO2 % Ti02 % A1203 % Fe2O3T % MnO % MgO % CaO % Na2O % K20 % P205 % LOI % CIA Ba ppm Rb ppm Sr ppm V ppm Cr ppm Ni ppm Co ppm Cu ppm Y ppm Zr ppm Cs ppm Hf ppm Nb ppm Ta ppm Pb ppm Sc Ppm Th ppm U Ppm La ppm Ce ppm Pr Ppm Nd ppm Sm ppm Eu ppm Gd ppm Tb ppm Dy Ppm Ho ppm Er Ppm Tm ppm Yb ppm Lu ppm Rb/Sr K/Rb Zr/Sc Th/Sc Ti/Zr Nb/Y La/Sc La/Th Th/U LaN/SmN GdN/YbN Sm/Nd LaN/YbN Eu/Eu* Ce/Ce* ZREE

VOL M

66.0 78.51 0.5 0.55 15.2 10.22 4.5 3.29 0.06 0.06 2.2 0.45 0.61 4.2 3.9 0.05 3.4 2.82 0.17 0.15 2.52 71 550 349 112 123 350 32 107 53 41 83 44 13 17 6 25 16 22 24 190 202 4.6 8.9 5.8 5.7 12 12.5 1 1.2 17 6.3 13.6 7.1 10.7 10.8 2.8 1.5 33.4 30 64 70.1 7.1 7.8 31.4 26 4.5 5.8 0.88 1.1 3.8 5.0 0.64 0.7 4.4 3.5 0.8 0.8 2.3 2.5 0.33 0.4 2.4 2.2 0.32 0.4 0.50 4.00 250 190 14.0 27.9 0.80 1.52 12.94 18.9 0.55 0.53 2.2 4.73 3.12 2.8 3.8 6.95 4.2 3.61 1.4 1.70 0.17 0.19 9.3 10.32 0.63 0.58 1.07 1.03 146 166

s.d.

RAN M

1.79 0.04 0.51 0.40 0.03 0.12 0.46 0.04 0.33 0.01 0.43 5 52 12 8 3 10 5 2 7 4 93 0.6 2.5 0.9 0.2 0.9 0.9 1.2 0.1 2.0 3.4 0.4 1.8 0.4 0.1 0.4 0.1 0.5 0.1 0.3 0.0 0.3 0.0 1.02 8 10.3 0.07 7.3 0.08 0.27 0.19 0.63 0.14 0.08 0.01 0.91 0.05 0.04 20

65.71 0.79 15.58 5.65 0.08 2.31 0.66 1.76 3.80 0.18 3.04 65 417 162 86 99 169 31 14 17 35 212 15.3 6.6 16.6 1.2 10.6 14.2 14.9 3.3 39.9 91.3 9.2 36.6 7.3 1.4 6.3 1.0 6.1 1.1 3.3 0.5 3.4 0.5 2.04 197 16.6 1.07 23.7 0.48 2.96 2.73 4.94 3.42 1.50 0.20 8.74 0.62 1.14 208

CORDILLERA ORIENTAL

s.d.

RT M

4.12 0.08 1.86 0.99 0.02 0.33 0.13 0.41 0.88 0.02 0.63 3 116 41 31 11 63 5 2 13 6 65 6.3 1.7 1.8 0.1 11.1 2.2 1.9 1.3 8.0 12.5 1.4 5.4 1.0 0.3 1.2 0.1 0.9 0.2 0.5 0.1 0.4 0.1 0.72 22 12.4 0.21 5.3 0.08 1.21 0.69 1.26 0.25 0.20 0.01 1.53 0.05 0.17 25

68.77 0.66 13.76 5.61 0.09 2.34 0.37 1.88 3.25 0.17 2.71 65 298 146 55 84 109 28 15 15 27 183 10.6 5.7 13.9 1.0 18.5 12.0 12.7 2.7 18.7 41.1 4.6 20.4 4.3 0.8 4.1 0.7 4.3 0.9 2.6 0.4 2.7 0.4 2.63 186 18.2 1.11 22.0 0.52 1.74 1.54 4.65 2.67 1.24 0.22 5.14 0.57 1.05 106

s.d.

TOR M

7.40 0.17 3.65 1.15 0.01 0.39 0.08 0.14 1.45 0.01 0.71 2 125 67 2 30 49 6 2 4 8 12 3.8 0.4 4.4 0.3 12.1 4.7 3.9 0.5 8.1 19.3 2.0 8.5 1.5 0.3 1.2 0.2 1.3 0.3 0.7 0.1 0.7 0.1 1.16 7 9.1 0.21 6.6 0.06 0.87 0.60 0.67 0.31 0.26 0.02 1.82 0.05 0.32 12

75.58 0.58 11.07 3.63 0.07 1.08 0.55 2.25 2.27 0.17 2.21 61 279 104 55 65 230 18 10 7 28 317 5.1 9.2 12.0 0.9 5.0 8.1 11.3 2.7 35.7 75.6 8.3 34.2 6.8 1.2 5.6 0.9 4.8 0.9 2.7 0.4 2.9 0.4 2.14 178 39.8 1.40 11.9 0.42 4.43 3.21 4.30 3.29 1.57 0.20 9.08 0.58 1.03 181

s.d.

COR M

s.d.

FED M

s.d.

SEC M

s.d.

MOL M

s.d.

CHO M

s.d.

1.53 0.06 0.80 0.51 0.08 0.18 0.24 0.37 0.61 0.01 0.38 1 126 19 13 6 52 4 4 6 2 102 1.1 3.0 1.2 0.1 1.7 0.9 1.5 0.3 4.4 8.9 0.9 3.8 0.5 0.1 0.6 0.1 0.5 0.1 0.2 0.0 0.3 0.0 1.21 17 15.7 0.19 3.8 0.01 0.74 0.59 0.67 0.23 0.14 0.01 0.92 0.03 0.03 22

63.43 0.74 15.76 6.07 0.13 2.30 1.00 1.86 4.01 0.16 3.90 62 512 177 46 97 138 30 16 25 35 178 9.3 5.3 15.9 1.2 13.8 14.6 15.5 3.3 38.1 79.4 8.3 36.9 7.6 1.3 6.3 1.0 5.7 1.1 3.2 0.5 3.3 0.5 4.39 188 14.6 1.14 24.9 0.46 2.91 2.52 4.70 3.15 1.57 0.21 8.65 0.57 1.00 193

7.84 0.23 4.26 1.69 0.08 0.85 1.23 0.35 1.54 0.02 0.93 9 128 65 15 34 44 11 5 12 7 18 3.5 0.6 4.5 0.3 10.9 5.5 4.0 0.6 7.1 16.6 1.5 6.3 1.4 0.2 1.0 0.2 1.0 0.2 0.6 0.1 0.8 0.1 2.24 5 7.7 0.24 7.3 0.09 0.95 0.29 0.51 0.18 0.23 0.01 1.30 0.02 0.05 23

65.00 0.76 15.64 6.05 0.05 2.04 0.35 1.88 3.78 0.19 3.45 67 352 170 58 93 112 32 15 19 35 201 8.8 6.2 15.8 1.2 4.4 14.3 15.4 2.9 38.6 82.7 8.9 37.8 7.6 1.3 6.5 1.1 5.8 1.1 3.3 0.5 3.4 0.5 3.13 185 14.0 1.07 22.9 0.45 2.69 2.51 5.31 3.17 1.54 0.20 8.39 0.54 1.03 199

2.48 0.07 1.18 0.39 0.02 0.59 0.04 0.15 0.50 0.02 0.27 1 37 21 19 9 16 5 1 4 2 26 1.1 0.6 1.4 0.1 1.4 1.4 1.1 0.4 3.6 6.8 0.8 3.0 0.6 0.1 0.3 0.1 0.4 0.1 0.2 0.0 0.2 0.0 0.77 6 1.1 0.05 1.2 0.04 0.07 0.11 0.48 0.09 0.03 0.01 0.48 0.01 0.03 23

62.13 0.80 16.93 6.62 0.06 2.69 0.47 1.59 4.25 0.19 4.04 68 413 190 52 102 114 33 15 83 38 177 14.9 5.7 17.4 1.3 3.7 16.0 16.5 3.9 41.9 95.0 10.4 42.7 7.8 1.3 6.8 1.2 6.5 1.3 3.9 0.6 3.7 0.5 3.70 186 11.2 1.04 27.3 0.46 2.64 2.57 4.36 3.35 1.52 0.18 8.49 0.53 1.08 224

2.91 0.08 1.38 0.38 0.01 0.14 0.04 0.27 0.54 0.02 0.95 2 41 26 6 11 57 2 1 74 5 16 2.5 0.5 1.3 0.1 1.5 1.8 2.4 0.9 3.0 7.3 0.8 3.2 0.5 0.1 0.6 0.1 0.7 0.2 0.5 0.1 0.5 0.1 0.67 8 1.9 0.16 3.8 0.05 0.26 0.22 0.80 0.09 0.14 0.01 0.60 0.02 0.02 27

67.46 0.73 14.91 5.33 0.08 2.20 0.79 1.07 4.79 0.19 2.12 64 655 186 59 85 140 27 12 6 32 237 11.5 7.1 15.7 1.2 3.5 13.2 14.6 2.7 28.1 57.4 7.4 30.2 5.7 1.0 5.2 0.9 5.1 1.1 3.3 0.5 3.2 0.5 3.33 214 18.2 1.10 18.7 0.49 2.13 1.94 5.36 3.11 1.30 0.19 6.45 0.52 0.96 149

2.66 0.07 1.23 0.48 0.01 0.21 0.15 0.29 0.51 0.01 0.19 1 105 19 13 6 25 3 1 3 2 28 1.3 0.9 1.8 0.1 0.6 1.1 1.4 0.3 1.5 5.5 0.2 1.9 0.3 0.0 0.3 0.0 0.4 0.1 0.2 0.0 0.2 0.0 1.10 5 3.3 0.06 3.8 0.04 0.17 0.16 0.63 0.18 0.04 0.01 0.43 0.02 0.14 17

66.27 0.73 14.89 6.31 0.07 2.37 0.42 1.92 3.70 0.14 2.96 65 395 156 50 83 75 29 15 13 29 224 9.7 6.7 15.4 1.2 30.1 14.3 13.8 2.7 37.6 83.8 9.1 37.8 6.5 1.1 5.4 0.9 5.0 1.0 3.0 0.5 2.9 0.4 3.26 196 22.6 1.04 25.0 0.52 3.10 2.82 5.18 3.64 1.74 0.17 9.82 0.57 1.06 195

7.95 0.12 3.46 1.93 0.02 0.70 0.15 0.57 1.59 0.04 0.88 4 124 65 7 26 28 8 4 10 3 131 4.1 3.4 3.1 0.2 35.9 4.9 2.8 0.2 5.6 16.5 1.2 4.7 0.8 0.2 0.7 0.1 0.6 0.1 0.4 0.1 0.4 0.1 1.55 13 24.5 0.26 10.8 0.07 1.67 0.76 1.23 0.29 0.18 0.01 2.10 0.04 0.10 24

M, mean; s.d., standard deviation; for outcrop abbreviations see Fig. 1.

Table 3. Comparison of trace element concentrations and ratios considered to be indicators for continental volcanic arc provenance for different sampling areas of Puncoviscana complex normalized to continental arc or UCC PUNA ARC

(Eu/Eu*)NA (Th/Sc)NC (Tippm) NC (Nbppm) NC (Tappm)NC

VOL UCC M

CORDILLERA ORIENTAL

RAN RT s.d. M s.d. M

1.00 0.6-0.7 0.58 0.05 0.62 0.05 <0.8 1.00 1.52 0.07 1.07 0.21 <4100 4100 0.80 0.06 1.15 0.12 «12 12.00 1.04 0.06 1.38 0.15 «1 1.00 1.25 0.20 1.22 0.13

0.57 1.11 0.97 1.16 1.04

TOR s.d. M

COR s.d. M s.d.

0.05 0.21 0.26 0.36 0.31

0.03 0.19 0.09 0.10 0.07

0.58 1.40 0.84 1.00 0.91

0.57 1.14 1.08 1.33 1.23

FED SEC M s.d. M

s.d.

0.02 0.54 0.01 0.53 0.02 0.24 1.07 0.05 1.04 0.16 0.33 1.12 0.11 1.17 0.12 0.37 1.32 0.12 1.45 0.11 0.34 1.18 0.10 1.34 0.14

MOL CHO M s.d. M s.d.

0.52 1.10 1.06 1.31 1.18

0.02 0.06 0.10 0.15 0.15

0.57 1.04 1.83 1.28 1.16

0.04 0.26 0.17 0.26 0.24

M, mean; s.d., standard deviation; for outcrop abbreviations see Fig. 1. Normalization factors: ARC, typical values for a volcanic continental arc provenance; UCC, upper continental crust. NA, arc-normalized; Nc, crust-normalized (from Taylor & McLennan (1985), McLennan (2001) and McLennan et al. (1990; 1993)).

THE PUNCOVISCANA COMPLEX Table 2. (continued) FAMATINA

AMBATO

SUN M 66.27

0.80 15.14 5.91 0.07 2.39 0.70 1.90 3.84 0.20 2.63 64 462 185 56 90 124 27 14 19 36 230 14.3 7.6 16.8 1.3 6.7 14.1 17.2 3.7 37.4 84.6 9.1 36.9 7.3 1.2 6.1 1.1 6.2 1.2 3.5 0.5 3.5 0.5 3.68 173 16.4 1.22 21.2 0.47 2.68 2.19 4.63 3.15 1.46 0.20 7.82 0.54 1.14 199

s.d. 1.49 0.04 0.78 0.32 0.02 0.10 0.32 0.22 0.39 0.01 0.56 2 71 23 18 7 22 3 1 17 4 37 3.6 0.9 1.2 0.1 7.2 0.9 1.7 0.3 11.2 14.7 2.4 10.0 1.7 0.3 1.2 0.1 0.7 0.1 0.3 0.0 0.2 0.0 1.51 16 3.2 0.12 2.8 0.05 0.88 0.65 0.27 0.35 0.24 0.01 2.20 0.03 0.22 24

CEB M

s.d.

71.19 6.75 0.56 0.18 13.95 3.86 4.22 1.20 0.03 0.01 1.44 0.49 1.49 1.26 2.44 1.34 2.37 1.33 0.22 0.22 1.53 0.99 7 60 349 254 60 132 71 166 29 64 55 195 16 5 3 8 12 6 47 40 238 50 9.4 3.1 7.6 1.3 14.9 4.0 1.2 0.3 4.4 3.1 11.1 3.5 14.8 4.7 4.2 2.2 42.3 15.3 89.1 33.9 9.7 3.7 40.4 15.8 4.2 8.8 1.6 0.6 8.0 5.0 0.8 1.3 7.2 5.0 1.4 1.0 4.1 2.9 0.7 0.5 4.3 2.9 0.6 0.4 1.14 1.24 142 22 23.5 9.4 1.36 0.27 14.2 4.2 0.40 0.13 4.11 1.71 2.98 1.03 3.72 0.90 3.11 0.56 1.55 0.25 0.21 0.02 8.28 3.12 0.58 0.09 1.02 0.04 220 25

M 71.69 0.55 12.91 4.48 0.08 1.69 0.34 1.67 3.26 0.13 2.39 64 379 141 56 76 68 21 10 25 27 108 13.9 3.3 11.7 1.0 14.0 10.1 10.7 3.0 29.6 60.6 6.9 28.2 5.6 1.0 5.0 0.8 4.8 0.9 2.6 0.4 2.5 0.4 2.74 204 12.6 1.17 29.9 0.45 3.33 2.82 3.91 3.31 1.65 0.20 8.88 0.55 0.99 149

s.d.

M

7.17 86.57

0.17 3.25 1.81 0.04 0.82 0.22 0.64 1.37 0.02 0.90 5 142 70 9 25 44 9 4 20 10 19 10.0 0.7 3.7 0.2 8.2 4.7 2.9 1.6 7.0 15.2 1.5 6.2 1.4 0.2 1.4 0.3 1.5 0.2 0.7 0.1 0.7 0.1 1.63 31 4.7 0.29 5.1 0.09 1.01 0.30 0.95 0.20 0.28 0.01 0.94 0.06 0.05 17

0.46 6.16 3.19 0.00 0.32 0.03 0.03 1.61 0.03 1.25 77 196 75 40 61 268 11 14 3 33 165 2.8 5.0 8.2 0.7 3.4 6.3 8.0 2.1 38.0 92.9 10.0 39.5 8.1 1.3 6.1 1.0 5.8 1.0 2.8 0.4 2.6 0.4 2.07 177 25.8 1.27 17.2 0.25 6.08 4.91 3.90 2.94 1 .93 0.21 11.08 0.55 1.16 210

NP M

LA M

s.d.

CON M

1.51 0.12 0.94 0.28 0.00 0.09 0.01 0.01 0.24 0.01 0.14 0 36 10 9 6 36 3 4 1 6 58 0.5 1.7 2.0 0.1 0.4 1.0 1.5 0.4 4.8 9.1 1.0 4.7 1.0 0.1 0.7 0.1 1.0 0.2 0.6 0.1 0.5 0.1 1.01 5 6.6 0.23 1.8 0.04 0.93 1.16 0.54 0.11 0.18 0.00 1.90 0.03 0.08 26

66.71 6.72 0.81 0.12 14.49 2.76 5.79 1.57 0.10 0.02 2.69 0.85 1.72 0.55 2.06 0.61 3.35 1.55 0.23 0.05 1.80 0.81 59 5 600 407 153 57 145 37 116 37 103 45 11 31 5 16 20 10 37 6 236 110 60.2 17.4 6.9 3.1 15.2 2.1 0.2 1.1 3.6 1.5 3.4 12.9 3.4 12.9 2.3 0.5 32.1 9.4 73.0 18.2 2.2 7.6 9.4 32.6 1.6 6.5 1.3 0.3 1.3 5.8 0.2 1.0 5.9 1.0 1.2 0.2 0.6 3.6 0.1 0.6 0.7 3.5 0.1 0.5 1.20 0.73 176 19 8.7 19.3 1.02 0.17 9.2 23.6 0.42 0.05 2.57 0.67 2.50 0.37 5.57 0.34 3.10 0.33 1.35 0.14 0.20 0.01 6.74 0.96 0.63 0.03 1.12 0.31 175 20

s.d.

M

s.d.

LA M

s.d.

0.03 0.17 0.80 0.17 0.16

0.52 1.01 1.42 1.51 1.39

0.06 0.13 0.60 0.51 0.44

0.62 1.19 0.94 1.13 1.10

0.17 0.46 0.27 0.30 0.23

SIJ-G

SIJ-P

s.d.

s.d.

65.95 11.03

0.97 0.41 16.02 5.55 5.79 1.81 2.01 0.64 0.10 0.02 0.58 0.36 2.11 0.66 3.22 2.00 0.14 0.06 2.92 0.84 3 67 619 342 154 88 81 43 95 47 19 71 10 33 4 16 7 7 9 41 49 248 2.1 5.0 1.4 6.5 18.1 6.1 0.4 1.4 1.1 1.9 15.8 7.3 15.3 5.5 3.5 1.0 51.5 20.7 105.4 39.7 4.1 11.3 43.0 15.0 2.9 8.8 0.5 1.5 2.4 8.0 1.2 0.3 6.9 1.6 1.3 0.3 4.2 1.0 0.2 0.6 1.2 4.2 0.2 0.6 2.00 0.81 19 171 18.4 7.8 1.01 0.13 8.7 23.6 0.43 0.06 3.35 0.29 3.34 0.25 4.40 0.68 3.62 0.34 1.55 0.10 0.21 0.01 8.88 1.11 0.52 0.06 1.04 0.03 248 30

s.d.

67.17 8.82 0.64 0.18 14.61 4.73 6.85 1.24 0.09 0.02 2.66 0.64 0.42 0.24 1.71 0.71 2.48 1.49 0.10 0.02 3.12 0.79 3 70 432 237 111 65 29 68 39 101 58 135 11 30 11 5 37 35 5 28 146 35 1.6 2.7 4.4 1.0 3.6 13.5 1.1 0.2 15.9 30.0 12.4 4.7 2.9 13.0 4.2 1.6 31.1 11.7 61.8 21.4 2.3 8.0 9.5 33.7 6.4 1.4 0.5 1.2 5.2 1.0 0.9 0.2 4.7 0.8 0.1 0.9 2.8 0.5 0.4 0.1 0.3 2.7 0.4 0.1 1.92 1.67 182 6 17.4 18.6 1.19 0.46 28.7 11.7 0.49 0.19 2.76 1.01 2.35 0.49 3.29 0.81 3.02 0.75 1.27 0.34 0.19 0.02 8.34 2.26 0.62 0.17 0.93 0.12 160 18

Table 3. (continued) AMBATO SUN M

s.d.

CEB M

0.54 1.22 1 1.40 1.29

0.03 0.12 0.05 0.10 0.11

0.58 1.36 0.82 1.24 1.19

FAMATINA

s.d.

M

s.d.

M

s.d.

CON M

0.09 0.27 0.26 0.33 0.31

0.55 1.17 0.81 0.97 0.95

0.06 0.29 0.26 0.31 0.24

0.55 1.27 0.67 0.68 0.65

0.03 0.23 0.17 0.17 0.14

0.63 1.02 1.14 1.27 1.13

SIJ-P

SIJ-G

NP

387

U. ZIMMERMANN

388

TabJe 4. Major and trace element data for the Puncoviscana complex Outcrop Sample Grain-size (phi) Si02 TiO2 A1203 Fe2O3T MnO MgO CaO Na2O K2O P205 LOI TOTAL CIA Ba Rb Sr V Cr Ni Co Cu Y Zr Cs Hf Nb Ta Pb Sc Th U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Cr/V Y/Ni Zr/Sc Th/Sc Ti/Zr Nb/Y La/Sc La/Th Th/U LaN/SmN Sm/Nd LaN/YbN GdN/YbN Eu/Eu* Ce/Ce* 2REE

% % % % % % % % % % % °/ ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm

VOL VOL VOL VOL VOL VOL VR28 VR29 VR33 VR34 VR35 VR36 5 5 8 5 5 5

77.49 76.95 0.52 0.60 10.05 10.82 3.04 3.87 0.08 0.05 0.70 0.49 1.37 0.57 0.12 0.07 2.73 3.08 0.16 0.14 3.27 2.60 99.78 98.98 64 71 296 450 117 139 41 32 50 57 32 62 17 7 8 8 22 8 30 19 127 352 9 10 4 10 12 14 1.1 1.2 5.1 6.3 6.6 9.0 10 13 1.5 1.7 32 38 68 75 7.6 8.4 30 35 6.5 5.7 1.0 1.2 4.5 5.5 0.6 0.9 3.7 4.9 1.0 0.7 2.0 2.8 0.4 0.3 2.0 2.7 0.4 0.3 0.65 1.08 1.07 4.40 19.1 39.1 1.56 1.49 10.2 24.8 0.63 0.46 4.79 4.20 3.07 2.82 7.08 7.88 3.64 3.51 0.19 0.18 11.6 10.5 1.82 1.66 0.56 0.60 1.05 0.97 158 183

79.45 0.56 9.92 3.40 0.01 0.39 0.42 0.01 2.66 0.16 2.11 99.10 73 346 116 28 52 36 12 3 18 19 115 8 4 12 1.1 7.4 6.5 10 1.6 32 69 7.8 31 5.8 1.0 4.6 0.7 3.8 0.7 2.1 0.3 2.1 0.3 0.69 1.58 17.7 1.55 29.6 0.63 4.99 3.22 6.27 3.51 0.19 11.3 1.79 0.56 1.04 162

79.06 0.55 10.32 2.77 0.07 0.35 0.25 0.01 3.12 0.16 2.12 98.80 73 324 131 29 53 36 14 7 18 23 162 8 5 13 1.4 6.1 6.7 10 1.4 33 72 8.1 31 5.9 1.0 5.3 0.7 4.5 0.9 2.5 0.4 2.4 0.4 0.68 1.66 24.1 1.55 20.4 0.55 4.93 3.17 7.47 3.55 0.19 10.0 1.75 0.55 1.05 168

78.84 76.23 0.57 0.54 10.42 10.69 3.61 3.42 0.08 0.04 0.48 0.38 1.14 0.25 0.07 0.02 2.65 3.19 0.17 0.16 2.86 2.26 98.90 99.17 76 65 318 381 134 116 41 37 54 56 37 35 18 16 7 6 21 21 25 26 168 178 10 9 5 5 13 13 1.4 1.5 7.3 7.0 7.2 6.7 11 11 1.7 1.5 33 33 71 71 8.0 7.9 31 31 5.8 5.8 1.1 1.1 5.2 5.3 0.8 0.8 4.7 4.9 0.9 1.0 2.7 2.6 0.4 0.4 2.7 2.6 0.4 0.4 0.64 0.66 1.51 1.47 23.4 26.6 1.52 1.58 18.2 20.3 0.54 0.50 4.61 4.89 3.03 3.09 6.32 7.23 3.56 3.56 0.19 0.19 9.5 9.0 1.65 1.60 0.57 0.57 1.06 1.05 168 167

VOL RAN RAN RAN RAN RAN VR37 RLA-2 RLA-8 RLA-9 RLA-10 RLJ-2 8 8 8 8 8 8

81.54 0.48 9.32 2.90 0.07 0.38 0.26 0.06 2.27 0.13 2.40 99.81 75 326 107 18 48 48 7 5 5 26 315 8 9 11 1.0 5.2 7.0 10 1.5 33 64 7.1 31 5.3 1.2 4.8 0.8 4.0 0.8 2.5 0.4 2.3 0.4 1.00 3.53 45.0 1.37 9.1 0.42 4.73 3.45 6.40 3.91 0.17 10.4 1.67 0.68 0.95 157

65.26 0.78 15.89 5.60 0.07 2.46 0.48 1.66 4.18 0.17 3.00 99.55 66 433 179 66 107 116 32 16 16 41 196 13 6 17 1.2 4.6 15.0 14 2.3 36 74 8.1 32 7.0 1.3 6.8 1.0 6.8 1.3 3.6 0.6 3.7 0.5 1.09 1.29 13.1 0.95 23.8 0.42 2.37 2.51 6.17 3.19 0.22 7.1 1.49 0.57 1.03 182

79.02 0.57 9.92 2.91 0.06 1.20 0.67 2.72 1.31 0.18 1.60 100.16 58 177 58 83 68 260 17 8 20 36 414 3 12 12 0.9 22.0 7.0 12 2.5 47 93 10.0 40 8.0 1.6 7.1 1.1 6.6 1.2 3.1 0.5 2.9 0.5 3.82 2.14 59.2 1.69 8.2 0.34 6.64 3.94 4.72 3.64 0.20 11.7 1.95 0.64 1.00 221

64.84 0.79 16.15 5.73 0.08 2.54 0.53 1.64 4.26 0.17 2.90 99.63 66 423 186 66 102 103 31 16 20 46 204 13 7 18 1.3 3.1 15.0 16 2.5 41 91 9.3 37 8.2 1.5 7.9 1.2 7.8 1.4 4.1 0.6 4.2 0.6 1.01 1.47 13.6 1.08 23.2 0.38 2.75 2.54 6.48 3.15 0.22 7.3 1.54 0.53 1.10 216

Measured by ICP-MS at ACME, Laboratories Vancouver (Canada). See Fig. 1 for locality abbreviations.

63.93 0.82 16.21 6.44 0.09 2.44 0.50 1.77 3.96 0.17 3.10 99.43 67 441 154 61 99 96 30 15 8 28 189 10 6 17 1.3 2.4 16.0 17 2.4 32 81 7.7 31 5.9 1.1 5.1 0.8 5.1 1.0 2.7 0.4 3.1 0.4 0.97 0.94 11.8 1.04 26.0 0.61 2.02 1.93 6.96 3.43 0.19 7.7 1.34 0.60 1.22 178

65.56 0.82 15.64 5.94 0.10 2.42 0.53 1.85 3.27 0.20 3.50 99.83 67 480 118 62 103 198 33 16 9 30 204 10 6 17 1.2 35.6 14.0 14 2.3 27 77 6.9 27 5.6 1.1 5.1 0.8 5.0 1.0 2.9 0.5 3.1 0.5 1.93 0.92 14.5 1.03 24.2 0.55 1.95 1.90 6.26 3.06 0.21 6.5 1.35 0.58 1.35 163

389

THE PUNCOVISCANA COMPLEX Table 4. (continued) RAN RAN RAN RAN RLJ-3 RLJ-4 RLJ-5 RO-1 8 8 8 8

65.73 0.85 15.70 5.87 0.09 2.36 0.77 2.00 3.28 0.20 2.90 99.75 66 580 159 79 97 109 31 15 3 39 244 18 8 18 1.4 4.2 15.0 16 3.1 48 94 10.6 41 7.8 1.4 7.1 1.1 6.5 1.3 3.6 0.5 3.8 0.5 1.13 1.27 16.2 1.05 20.9 0.47 3.19 3.03 5.10 3.85 0.19 9.4 1.54 0.57 0.98 227

67.25 0.73 14.81 4.49 0.04 2.34 0.74 2.42 4.44 0.18 2.20 99.64 59 504 226 73 88 103 25 13 1 44 176 17 6 16 1.2 1.6 14.0 14 3.2 58 111 12.0 47 9.3 2.0 8.8 1.2 7.6 1.4 4.2 0.6 3.8 0.6 1.17 1.75 12.6 1.00 24.8 0.37 4.11 4.11 4.38 3.88 0.20 11.2 1.87 0.66 0.99 267

66.64 0.83 15.30 5.43 0.10 2.32 0.55 1.91 3.11 0.21 3.20 99.60 67 491 117 66 95 123 32 13 11 34 247 10 8 17 1.2 29.8 14.0 15 2.6 39 90 9.2 38 7.4 1.4 6.3 1.0 6.0 1.1 3.3 0.5 3.3 0.5 1.30 1.09 17.7 1.04 20.1 0.49 2.76 2.64 5.62 3.27 0.20 8.5 1.53 0.59 1.12 206

62.41 0.86 17.65 5.58 0.06 2.27 0.77 1.26 4.74 0.16 3.90 99.66 67 575 184 164 105 164 31 14 21 39 189 16 6 19 1.3 17.8 14.0 17 3.8 49 111 11.3 46 8.5 1.7 7.0 1.1 6.7 1.2 3.8 0.6 3.7 0.5 1.56 1.25 13.5 1.19 27.3 0.49 3.51 2.95 4.39 3.63 0.19 9.7 1.52 0.65 1.10 252

RAN RO-3 8

RT RAN RAN RAN RAN RO-4 RO-10 RO-11 RO-12 02-90 8 8 8 8 8

63.11 0.79 16.77 6.12 0.05 2.50 0.64 1.33 4.49 0.16 3.10 99.06 67 459 185 80 96 164 32 15 2 33 152 16 5 16 1.2 3.8 15.0 14 3.5 39 88 9.3 37 7.3 1.5 5.9 1.0 5.7 1.0 3.2 0.5 3.2 0.4 1.71 1.02 10.1 0.96 31.2 0.47 2.60 2.71 4.11 3.35 0.20 9.0 1.50 0.65 1.09 203

63.25 0.86 16.75 5.83 0.14 2.34 0.80 1.62 3.96 0.17 3.80 99.52 67 357 154 142 111 239 34 16 27 34 229 30 7 19 1.4 7.6 16.0 18 7.5 34 95 8.1 33 6.5 1.2 5.2 0.9 5.7 1.1 3.5 0.6 3.5 0.5 2.16 0.99 14.3 1.14 22.5 0.55 2.12 1.86 2.43 3.27 0.20 7.2 1.23 0.60 1.34 198

62.17 0.90 17.45 5.55 0.07 2.29 0.91 1.40 4.61 0.20 3.80 99.35 66 379 180 66 98 185 30 14 25 30 207 16 6 18 1.3 4.4 14.0 16 3.3 34 112 8.4 34 6.5 1.4 4.8 0.8 5.2 1.0 3.0 0.5 3.0 0.5 1.89 0.97 14.8 1.16 26.1 0.60 2.46 2.11 4.94 3.32 0.19 8.3 1.28 0.76 1.55 215

65.59 0.72 15.01 6.88 0.08 2.44 0.70 1.53 3.76 0.15 2.90 99.76 66 273 188 99 109 253 35 15 38 32 161 22 5 16 1.1 5.5 15.0 12 3.3 38 83 9.0 35 6.7 1.2 5.4 0.9 5.5 1.0 3.0 0.5 3.0 0.5 2.32 0.93 10.7 0.80 26.9 0.48 2.50 3.13 3.64 3.51 0.19 9.1 1.43 0.61 1.08 193

RT 02-91 8

65.14 74.11 76.67 0.46 0.52 0.72 9.42 14.88 11.04 4.72 4.73 6.76 0.08 0.07 0.08 1.98 2.03 2.47 0.34 0.30 0.71 2.00 1.96 1.55 1.45 2.20 3.86 0.17 0.19 0.16 2.00 2.10 2.70 99.03 99.29 99.25 64 64 65 141 272 231 62 96 181 52 55 95 104 47 58 198 151 253 24 25 37 14 14 14 42 19 16 21 18 31 198 183 160 7 8 20 6 6 5 9 10 15 0.7 0.8 1.1 38.0 15.4 6.3 6.0 8.0 15.0 13 9 8 2.1 2.3 3.3 21 13 37 28 25 80 4.9 3.1 8.6 22 15 35 3.0 4.6 7.0 0.8 0.5 1.3 4.1 5.2 2.8 0.7 0.5 0.9 5.4 3.4 2.8 0.7 0.6 1.0 2.0 1.8 2.9 0.3 0.3 0.4 2.0 1.9 2.8 0.3 0.4 0.3 4.22 2.60 2.43 0.71 0.84 0.86 33.1 22.9 10.7 1.45 0.98 0.86 17.1 27.0 13.9 0.41 0.56 0.47 3.43 1.58 2.48 1.62 2.88 2.37 4.14 3.39 3.91 2.81 2.63 3.33 0.21 0.20 0.20 7.5 4.9 9.6 1.64 1.17 1.48 0.57 0.56 0.65 0.63 0.91 1.05 95 70 188

RT 02-92 3

RT 02-93 8

RT 02-94 8

RT 02-95 5

61.90 0.82 17.12 6.85 0.09 2.69 0.43 1.65 4.75 0.18 3.40 99.88 66 437 215 56 108 103 35 18 12 35 181 15 6 19 1.5 9.4 16.0 17 3.3 17 29 3.7 17 3.7 0.7 4.4 0.8 5.3 1.1 3.4 0.5 3.5 0.5 0.95 1.00 11.3 1.06 27.1 0.54 1.04 0.98 5.12 2.81 0.22 3.5 1.03 0.54 0.82 90

64.14 0.79 16.04 6.09 0.10 2.60 0.47 1.84 4.25 0.16 3.30 99.78 65 357 181 53 97 82 30 17 12 33 192 11 6 16 1.2 15.1 15.0 16 3.2 29 71 7.0 32 6.4 1.2 5.7 1.0 5.4 1.1 3.0 0.5 3.3 0.5 0.85 1.10 12.8 1.07 24.6 0.48 1.91 1.79 5.00 2.80 0.20 6.4 1.40 0.61 1.14 167

63.68 0.80 16.58 5.96 0.09 2.49 0.38 1.86 4.32 0.15 3.20 99.51 66 395 203 59 115 68 32 17 11 30 178 15 6 18 1.2 6.9 16.0 15 2.9 11 40 3.1 13 3.3 0.7 3.4 0.6 4.5 1.0 2.9 0.5 3.2 0.5 0.59 0.94 11.1 0.96 27.0 0.60 0.68 0.71 5.31 2.07 0.25 2.5 0.85 0.59 1.64 88

78.72 0.45 9.31 3.98 0.07 1.79 0.25 2.06 1.54 0.16 1.80 100.13 63 139 72 55 54 103 19 12 14 17 185 6 5 10 0.7 33.2 7.0 10 2.0 11 29 2.9 12 2.8 0.5 2.9 0.4 2.9 0.6 1.7 0.3 1.9 0.3 1.90 0.91 26.4 1.36 14.6 0.55 1.61 1.19 4.75 2.53 0.24 4.4 1.24 0.48 1.21 69

U. ZIMMERMANN

390

Table 4. (continued) Outcrop Sample Grain-size (phi) SiO2 TiO2 A12O3 Fe2O3T

MnO MgO CaO Na2O

K2O P205

LOI TOTAL

CIA Ba Rb Sr V Cr Ni Co Cu Y Zr Cs Hf Nb Ta Pb Sc Th U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Cr/V Y/Ni Zr/Sc Th/Sc Ti/Zr Nb/Y La/Sc La/Th Th/U La^/SniN Sm/Nd LaN/YbN GdN/YbN Eu/Eu* Ce/Ce* ZREE

°/ % % % % % % % % % % % ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm

RT 02-96 3

62.15 0.79 16.84 6.94 0.10 2.78 0.43 1.77 4.27 0.16 3.20 99.43

67 383 194 55 110 62 33 17 22 34 160 14 5 16 1.2 11.6 16.0

15 3.0 30 66 7.7 31 6.3 1.3 5.8 0.9 5.9 1.1 3.1 0.5 3.3 0.5 0.56 1.02 10.0 0.91 29.5 0.48 1.90 2.10 4.83 3.02 0.20

6.8 1.41 0.64 1.03

164

TOR 72 3

73.74 0.64 11.82 3.92 0.05 1.06 0.73 1.99 2.70 0.16 2.50 99.31

61 278 116 51 70 260 21 11 9 29 329 5 9 13 1.0 7.6 8.0 11 3.0 41 85 9.3 38 6.9 1.3 6.4 0.9 5.0 0.9 2.7 0.4 2.9 0.4

TOR 73 5

TOR 75 5

75.80 75.07 0.59 0.59 11.41 11.97 3.87 2.74 0.05 0.03 1.14 0.80 1.01 0.40 2.05 1.64 2.49 3.23 0.17 0.18 2.20 2.20 99.58 100.05

63 236 137 29 73 308 14 6 4 29 212 5 6 12 0.9 2.7 9.0 11 2.1 32 70 7.7 32 6.5 1.2 5.6 0.9 4.9 0.9 2.8 0.4 2.8 0.4

59 300 107 52 63 239 14 9 3 29 314 4 9 12 0.9 4.4 9.0 12 2.9 35 77 8.2 33 6.5 1.2 5.6 0.9 4.8 0.9 2.6 0.5 2.9 0.4

TOR 67-1 2

77.84 0.53 9.84 3.46 0.03 1.03 0.41 2.38 1.65 0.17 2.00 99.34

60 175 86 62 60 253 17 8 5 27 482 5 14 11 0.9 4.3 7.0 9 2.8 40 82 9.2 37 7.3 1.2 5.2 0.8 4.8 0.9 2.6 0.4 2.9 0.4 4.22 1.62 43.1 1.56 10.4 0.43 4.43 2.85 4.83 3.29 0.20

6.6

9.7

0.40 5.70 4.24 3.36 3.43 0.20 10.2 1.81 0.57 1.01

0.42 4.23 2.45 4.84 3.09 0.21

3.80 2.08 22.3 1.24 17.2 0.43 4.15 3.35 4.13 3.06 0.20

8.0

9.3

8.4

1.63 0.56 1.00

1.56 0.64 1.01

1.46 0.56 1.02

3.71 1.41 68.8 1.34

201

4.22 2.05 44.8 1.73

167

179

195

TOR 69 8

77.39 0.51 10.36 3.21 0.03 1.01 0.37 2.62 1.72 0.15 2.20 99.57

60 183 91 64 55 151 16 7 5 25 314 6 9 11 0.9 4.8 7.0 12 2.5 30 61 6.9 29 6.0 1.0 4.8 0.8 4.0 0.8 2.4 0.4 2.7 0.4 2.74 1.57 41.2 1.41 11.6 0.44 5.09 3.60 3.77 3.70 0.18 10.4 1.42 0.57 1.02

150

TOR 70 2

74.61 0.51 10.65 4.10 0.24 1.38 0.43 2.62 1.68 0.17 2.80 99.19

61 545 85 69 63 205 25 17 20 27 178 7 5 12 0.8 7.1 8.0 10 2.4 33 70 7.6 33 6.8 1.3 5.5 0.9 4.8 0.9 2.5 0.4 2.6 0.3

TOR 71 5

74.60 0.67 11.44 4.11 0.03 1.15 0.49 2.43 2.42 0.18 1.60 99.12

60 235 109 58 70 192 18 9 4 33 388 5 12 14 1.0 4.0 9.0 14 2.9 40 84 9.2 39 7.6 1.3 6.3 0.9 5.6 1.1 3.1 0.5 3.5 0.5

COR 02-24 5

63.71 0.74 16.21 6.13 0.12 2.48 0.38 2.15 3.78 0.15 3.20 99.05

66 396 175 44 97 96 32 15 34 35 165 10 5 17 1.3 2.4

67 495 201 39 113 116 35 18 34 37 169 10 5 17 1.4 14.2 17.0

17 3.3 37 80 8.1 34 7.1 1.2 6.3 1.0 5.8 1.1 3.2 0.5 3.2 0.5

16 3.5 42 87 9.0 39 8.0 1.4 6.3 1.0 6.0 1.1 3.3 0.6 3.5 0.5

2.74 1.82 34.9 1.28 11.3 0.43 3.87 3.03 3.97 3.36 0.20

0.99 1.10 11.0 1.13 26.9 0.48 2.45 2.18 5.12 3.25 0.21

8.5

8.8

8.4

1.69 0.58 1.05

1.45 0.58 1.07

1.57 0.54 1.06

202

60.89 0.83 17.28 6.69 0.15 2.63 0.30 1.80 4.56 0.16 3.70 98.99

15.0

3.26 1.08 23.6 1.21 16.7 0.41 3.56 2.94 5.19 3.09 0.21

170

COR 02-25 8

189

COR 02-26 8

53.83 1.00 20.94 8.22 0.06 3.26 0.36 1.57 6.01 0.19 4.40 99.84

68 504 261 36 138 103 42 22 28 40 198 14 6 22 1.6 3.7 21.0

21 3.9 49 104 10.7

47 9.8 1.6 7.2 1.2 6.6 1.3 4.0 0.6 4.3 0.6

COR 02-27 8

58.88 0.87 17.81 6.98 0.15 2.79 0.97 1.37 4.93 0.18 4.50 99.43

66 693 209 38 113 178 36 18 33 35 171 11 5 18 1.4 29.4 18.0

17 3.4 39 84 8.4 39 8.0 1.5 6.5 1.1 6.1 1.1 3.4 0.5 3.8 0.5 1.57 0.98

1.03 1.04

0.74 0.94

9.9

9.4

9.5

0.95 29.4 0.48 2.46 2.58 4.63 3.27 0.21

1.00 30.2 0.55 2.33 2.32 5.41 3.13 0.21

0.93 30.6 0.50 2.18 2.33 4.94 3.07 0.21

8.8

8.4

7.7

1.47 0.58 1.02

1.35 0.56 1.03

1.41 0.59 1.02

209

Measured by ICP-MS at ACME, Laboratories Vancouver (Canada). See Fig. 1 for locality abbreviations.

247

203

391

THE PUNCOVISCANA COMPLEX Table 4. (continued)

COR

COR

COR

FED

FED

SEC SEC SEC SEC FED FED FED SEC SEC 02-22 02-22-1 02-23 02-161 02-163 02-164 02-165 02-166 02-167 2 5 3 8 5 5 5 5 8

02-28 5

02-29 8

02-30 2

02-19 2

FED 02-20 8

02-21 8

59.86 0.87 18.24 6.73 0.06 2.70 0.33 1.68 4.84 0.17 4.10 99.58 68 428 213 36 117 103 35 18 29 34 183 11 6 18 1.4 3.5 17.0 17 3.7 41 86 8.9 40 8.2 1.4 6.4 1.0 5.6 1.1 3.3 0.6 3.6 0.5 0.88 0.97 10.7 1.01 28.6 0.53 2.42 2.40 4.65 3.15 0.20 8.4 1.43 0.58 1.00 208

69.10 0.47 10.55 4.63 0.30 1.28 3.71 2.15 2.25 0.14 5.10 99.68 45 680 101 78 57 164 16 10 11 41 204 6 6 11 0.8 20.3 8.0 11 2.7 31 60 6.8 32 7.0 1.4 7.2 1.2 6.2 1.2 3.3 0.5 3.1 0.4 2.88 2.54 25.5 1.35 13.8 0.26 3.90 2.89 4.00 2.79 0.22 7.5 1.91 0.58 0.92 161

77.73 0.39 9.29 3.09 0.08 0.96 0.97 2.32 1.72 0.12 2.30 98.97 56 385 77 50 44 205 14 8 6 21 156 4 4 9 0.7 23.0 6.0 10 2.3 28 56 6.3 28 5.2 0.9 4.1 0.6 3.4 0.6 1.9 0.3 1.8 0.3 4.67 1.49 26.1 1.58 14.9 0.44 4.63 2.93 4.13 3.35 0.19 11.4 1.87 0.57 0.96 136

68.04 0.63 13.66 5.78 0.08 2.35 0.29 1.95 3.00 0.16 3.10 99.04 66 320 136 40 79 116 32 16 20 35 154 8 5 14 1.0 4.8 12.0 14 2.3 32 71 7.7 33 6.6 1.2 6.3 1.0 5.3 1.0 3.2 0.5 3.2 0.4 1.47 1.09 12.8 1.13 24.6 0.39 2.69 2.39 5.87 3.07 0.20 7.5 1.60 0.53 1.05 173

67.32 0.78 15.38 5.55 0.05 0.92 0.36 2.00 3.55 0.20 3.30 99.41 66 318 154 95 86 137 23 14 12 31 219 9 7 15 1.2 3.2 14.0 15 3.0 38 85 9.1 37 7.5 1.2 6.1 1.0 5.5 1.0 3.1 0.5 3.2 0.5 1.59 1.34 15.6 1.08 21.4 0.49 2.74 2.54 5.03 3.21 0.20 8.8 1.53 0.53 1.07 198

61.17 64.11 0.78 0.85 17.30 15.99 6.12 6.69 0.03 0.07 2.16 2.63 0.36 0.40 1.90 1.60 3.87 4.50 0.18 0.21 3.50 3.90 99.32 99.00 67 68 351 420 194 179 57 50 98 105 89 109 35 35 15 18 20 24 38 36 197 230 9 11 6 7 16 18 1.2 1.3 3.3 6.1 15.0 16.0 16 17 3.3 3.1 39 44 84 92 9.0 10.0 40 42 7.5 8.3 1.4 1.3 6.5 7.1 1.0 1.2 5.8 6.5 1.1 1.2 3.4 3.6 0.5 0.6 3.5 3.7 0.4 0.5 0.91 1.04 1.01 1.09 14.4 13.1 1.04 1.07 23.7 22.1 0.45 0.46 2.58 2.73 2.40 2.61 4.88 5.39 3.23 3.29 0.19 0.20 8.2 8.7 1.50 1.55 0.56 0.55 1.03 1.02 221 203

64.18 0.78 15.97 6.16 0.03 2.17 0.35 1.88 4.01 0.18 3.50 99.21 66 349 181 54 96 116 35 14 20 34 201 8 6 16 1.2 3.2 15.0 15 3.1 40 85 9.0 38 7.9 1.4 6.6 1.1 5.9 1.1 3.3 0.5 3.4 0.5 1.21 0.98 13.4 0.99 23.3 0.48 2.63 2.65 4.81 3.13 0.21 8.6 1.57 0.55 1.05 203

65.16 0.76 15.55 5.97 0.04 2.00 0.33 1.97 3.77 0.18 3.40 99.13 66 354 177 52 94 103 30 15 17 35 204 9 6 16 1.2 5.9 14.0 16 2.7 39 79 8.6 38 7.9 1.3 6.3 1.0 5.9 1.0 3.3 0.5 3.3 0.5 1.09 1.19 14.6 1.14 22.3 0.45 2.78 2.45 5.89 3.09 0.21 8.6 1.55 0.53 0.98 195

66.50 0.73 14.58 6.03 0.06 2.49 0.51 2.17 3.35 0.21 3.10 99.73 64 358 151 63 91 233 32 14 19 34 207 11 6 16 1.2 1.9 14.0 13 3.1 38 87 9.5 38 7.3 1.2 6.8 1.2 5.9 1.2 3.5 0.5 3.1 0.5 2.56 1.07 14.8 0.91 21.1 0.47 2.69 2.97 4.10 3.24 0.19 8.9 1.75 0.51 1.09 203

64.73 0.81 16.20 6.51 0.07 2.58 0.42 1.53 4.34 0.19 2.40 99.78 67 458 178 46 103 144 34 15 1 37 185 13 6 17 1.4 1.8 15.0 15 2.5 42 94 10.2 45 7.6 1.3 6.9 1.1 6.2 1.3 3.6 0.6 3.5 0.5 1.39 1.07 12.3 1.02 26.2 0.45 2.77 2.71 6.12 3.42 0.17 8.8 1.62 0.52 1.06 223

60.79 0.84 17.47 6.82 0.07 2.78 0.49 1.55 4.21 0.18 4.60 99.80 69 396 193 56 104 82 35 16 62 36 171 16 6 18 1.4 4.1 17.0 16 4.0 41 93 10.2 42 7.9 1.3 6.6 1.2 6.3 1.3 3.7 0.6 3.6 0.6 0.79 1.01 10.1 0.94 29.4 0.49 2.39 2.54 4.00 3.22 0.19 8.3 1.48 0.54 1.08 219

57.67 0.92 19.04 7.16 0.05 2.91 0.52 1.31 5.20 0.19 5.10 100.07 69 474 238 51 123 75 35 16 163 41 183 19 6 20 1.5 5.1 19.0 19 4.5 47 109 11.7 47 8.7 1.5 7.4 1.3

6.9 1.5 4.3 0.7 4.1 0.6 0.61 1.15 9.6 1.02 30.2 0.49 2.48 2.45 4.29 3.40 0.18 8.5 1.47 0.54 1.10 252

61.57 0.84 17.30 6.54 0.05 2.66 0.48 1.61 4.30 0.20 4.20 99.75 68 411 196 51 102 89 34 14 169 37 174 15 5 18 1.2 2.7 17.0 17 4.1 42 94 10.2 42 7.7 1.3 7.1 1.2 6.3 1.3 3.6 0.6 3.6 0.5 0.87 1.09 10.2 0.98 29.0 0.48 2.47 2.53 4.05 3.42 0.18 8.6 1.60 0.52 1.06 221

60.66 0.82 17.41 6.92 0.06 2.74 0.48 1.50 4.33 0.18 4.50 99.60 68 410 194 50 100 82 34 14 21 34 155 16 5 17 1.2 5.4 16.0 16 3.8 41 90 9.9 39 7.4 1.3 5.7 1.1 5.8 1.1 3.6 0.5 3.3 0.5 0.82 1.01 9.7 0.99 31.7 0.49 2.54 2.55 4.18 3.44 0.19 9.1 1.39 0.57 1.06 210

392

U. ZIMMERMANN

Table 4. (continued) Outcrop Sample Grain-size (phi) SiO2 Ti02 A1203 Fe2O3T MnO MgO CaO Na2O K2O P205 LOI TOTAL CIA Ba Rb Sr V Cr Ni Co Cu Y Zr Cs Hf Nb Ta Pb Sc Th U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Cr/V Y/Ni Zr/Sc Th/Sc Ti/Zr Nb/Y La/Sc La/Th Th/U LaN/SmN Sm/Nd LaN/YbN GdN/YbN Eu/Eu* Ce/Ce* 2REE

% % % % % % % % % %
ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm

SEC 02-168 8

MOL MOL MOL MOL MOL 02-200 02-201 02-202 02-203 02-204 2 5 8 3 8

62.99 0.66 16.48 6.36 0.05 2.65 0.42 1.45 4.04 0.16 4.40 99.66 69 382 179 47 91 96 30 14 145 49 167 14 6 17 1.5 4.9 14.0 20 5.2 44 99 11.2 45 8.3 1.4 7.5 1.4 7.8 1.6 4.8 0.7 4.5 0.7 1.05 1.65 11.9 1.39 23.7 0.35 3.14 2.26 3.75 3.32 0.18 7.3 1.36 0.52 1.06 238

67.47 0.73 14.65 5.55 0.07 2.30 0.56 0.60 5.21 0.19 2.10 99.43 66 553 198 38 82 109 26 12 7 31 214 12 6 15 1.2 2.6 13.0 14 2.4 27 60 7.1 28 5.2 0.9 5.2 0.9 5.0 1.0 3.3 0.5 3.2 0.4 1.34 1.20 16.4 1.06 20.5 0.49 2.05 1.93 5.75 3.22 0.19 6.2 1.33 0.53 1.05 147

70.94 0.63 13.28 4.62 0.07 1.95 0.86 1.12 4.14 0.18 2.00 99.79 63 581 161 63 80 178 25 11 5 29 266 10 9 14 1.0 3.2 12.0 14 2.4 28 55 7.5 32 5.4 0.9 4.6 0.9 4.6 1.0 3.0 0.4 2.9 0.4 2.22 1.15 22.2 1.20 14.2 0.49 2.33 1.94 6.00 3.25 0.17 7.1 1.29 0.55 0.90 146

63.50 0.82 16.72 5.93 0.09 2.49 0.83 1.09 5.40 0.20 2.40 99.47 65 816 211 62 96 144 31 14 4 34 206 13 7 19 1.4 4.1 15.0 17 3.0 29 51 7.3 31 5.8 1.0 5.4 0.9 5.6 1.2 3.7 0.5 3.4 0.5 1.50 1.09 13.7 1.13 23.9 0.56 1.96 1.73 5.67 3.18 0.18 6.4 1.30 0.51 0.81 147

68.14 0.72 14.78 5.21 0.07 2.09 0.77 1.12 4.52 0.19 2.20 99.81 64 632 182 60 84 137 25 11 3 33 264 11 7 15 1.2 3.9 13.0 14 2.9 26 66 7.2 28 5.9 0.9 5.1 0.9 5.3 1.1 3.4 0.5 3.3 0.5 1.63 1.29 20.3 1.07 16.4 0.47 2.03 1.90 4.79 2.80 0.21 5.9 1.25 0.51 1.17 154

67.27 0.74 15.10 5.34 0.08 2.19 0.95 1.40 4.68 0.20 1.90 99.85 62 695 178 72 82 130 29 13 12 33 237 12 7 15 1.1 3.9 13.0 14 3.0 30 56 7.7 32 6.0 1.0 5.5 0.9 5.2 1.1 3.2 0.5 3.3 0.5 1.59 1.13 18.2 1.05 18.7 0.47 2.29 2.18 4.57 3.11 0.19 6.6 1.34 0.51 0.88 153

CHO 02-7 5

CHO 02-2 2

CHO 02-4 3

CHO 02-5 5

58.87 0.84 17.80 8.23 0.08 3.00 0.27 1.51 5.21 0.17 3.70 99.68 67 528 213 52 111 55 36 18 6 32 162 13 5 18 1.4 14.0 18.0 16 2.7 33 61 7.9 34 6.0 1.0 4.9 0.9 5.2 1.1 3.4 0.5 3.0 0.5 0.49 0.88 9.0 0.89 31.2 0.57 1.83 2.05 5.96 3.45 0.18 8.2 1.33 0.52 0.87 162

78.26 0.54 9.77 2.91 0.04 1.10 0.70 2.81 1.51 0.19 1.90 99.73 56 226 64 63 44 130 15 7 7 25 381 4 10 11 0.8 3.3 7.0 10 2.9 39 78 8.6 36 6.0 1.1 4.8 0.8 4.3 0.8 2.4 0.4 2.4 0.4 2.95 1.67 54.4 1.36 8.5 0.43 5.57 4.11 3.28 4.07 0.17 12.2 1.65 0.59 0.98 185

77.88 69.63 0.70 0.57 9.62 13.65 5.44 4.13 0.07 0.06 2.26 1.51 0.32 0.48 2.47 2.47 2.86 1.38 0.17 0.17 2.30 1.60 99.87 99.87 64 60 381 207 118 60 51 53 74 44 62 109 33 20 15 10 7 8 28 26 153 484 7 4 14 5 14 11 1.1 0.8 14.9 12.8 12.0 7.0 12 11 3.1 2.5 27 42 59 90 7.0 9.8 29 40 5.0 7.1 0.9 1.1 4.6 4.9 0.8 0.9 4.7 4.6 1.0 0.9 3.2 2.5 0.5 0.4 3.2 2.6 0.4 0.5 0.83 2.49 0.85 1.32 12.7 69.1 0.96 1.54 27.5 7.1 0.50 0.43 2.21 5.97 2.30 3.87 3.71 4.32 3.32 3.69 0.17 0.18 6.1 12.0 1.18 1.54 0.52 0.57 1.03 1.04 145 207

Measured by ICP-MS at ACME, Laboratories Vancouver (Canada). See Fig. 1 for locality abbreviations.

CHO 02-8 5

CHO 02-9 5

60.06 0.81 17.73 7.72 0.08 2.95 0.21 1.38 5.46 0.11 3.30 99.81 67 531 205 42 100 55 37 19 6 26 145 13 4 16 1.3 13.6 19.0 16 2.4 36 89 9.1 38 6.3 1.1 5.4 0.9 4.6 0.9 2.7 0.4 2.7 0.4 0.55 0.71 7.6 0.83 33.6 0.62 1.88 2.27 6.54 3.55 0.16 9.7 1.61 0.54 1.16 197

61.68 0.78 17.30 7.05 0.10 2.88 0.42 1.73 4.26 0.15 3.60 99.95 68 378 193 52 102 62 38 19 14 35 159 12 6 17 1.3 6.3 17.0 16 2.7 41 95 10.4 42 7.8 1.3 6.7 1.1 6.0 1.2 3.6 0.5 3.5 0.5 0.60 0.92 9.4 0.92 29.4 0.48 2.39 2.59 5.81 3.27 0.19 8.5 1.53 0.54 1.10 220

393

THE PUNCOVISCANA COMPLEX Table 4. (continued) CHO 02-12 5

CEB CEB CHO SUN SUN SUN SUN SUN SUN SUN CEB CEB CEB CEB 02-13 02-217 02-218 02-219 02-220 02-222 02-223 02-224 02-262 02-263 02-264 02-265 02-267 02-268 5 5 1 5 3 2 3 5 2 5 5 5 5 8

61.86 0.81 16.64 7.45 0.07 2.61 0.46 1.50 4.42 0.08 3.70 99.60 67 456 198 43 91 62 28 16 30 31 155 12 5 18 1.3 91.7 17.0 17 2.7 42 100 10.2 42 6.7 1.4 5.9 1.0 5.6 1.0 3.1 0.5 2.8 0.4 0.68 1.09 9.1 0.97 31.3 0.58 2.45 2.52 6.11 3.89 0.16 10.8 1.68 0.64 1.15 222

61.95 0.81 16.61 7.52 0.07 2.62 0.47 1.47 4.47 0.08 3.60 99.67 67 454 199 46 94 68 27 16 29 31 155 13 5 18 1.3 83.8 17.0 15 2.6 42 100 10.2 42 6.9 1.3 5.9 1.0 5.4 1.1 3.2 0.5 2.8 0.4 0.73 1.16 9.1 0.87 31.3 0.58 2.49 2.86 5.69 3.85 0.17 11.1 1.70 0.59 1.13 223

69.25 0.76 13.59 5.47 0.07 2.19 0.72 2.37 3.05 0.18 2.20 99.85 61 332 144 69 78 137 24 13 55 36 253 10 8 15 1.2 8.3 13.0 16 3.3 40 82 9.5 38 7.4 1.3 6.4 1.1 6.4 1.2 3.4 0.5 3.4 0.5 1.75 1.46 19.5 1.21 18.0 0.43 3.08 2.55 4.76 3.39 0.20 8.7 1.55 0.54 1.00 201

66.21 0.80 15.63 5.86 0.06 2.41 1.22 1.78 4.09 0.20 1.70 99.96 62 560 172 82 87 123 25 12 10 35 207 13 7 16 1.2 6.8 15.0 17 3.6 41 83 9.4 38 7.4 1.3 6.3 1.1 6.1 1.2 3.3 0.5 3.4 0.5 1.42 1.36 13.8 1.10 23.2 0.47 2.72 2.47 4.58 3.46 0.19 8.9 1.51 0.58 0.99 202

66.21 0.76 14.76 5.97 0.08 2.36 1.07 1.95 3.70 0.20 2.70 99.76 62 455 188 72 87 82 23 14 19 39 228 9 7 17 1.3 22.0 13.0 16 3.6 43 93 10.5 43 8.2 1.4 6.5 1.2 6.8 1.3 3.7 0.6 3.8 0.5 0.94 1.71 17.6 1.19 20.0 0.45 3.32 2.79 4.31 3.30 0.19 8.4 1.39 0.57 1.03 223

65.41 0.87 15.51 6.42 0.08 2.44 0.53 1.87 3.97 0.22 2.60 99.92 65 514 186 47 92 123 28 13 20 42 302 15 9 19 1.5 4.5 14.0 20 4.1 51 106 12.0 50 9.6 1.5 7.5 1.3 6.8 1.3 3.9 0.6 3.8 0.6 1.34 1.47 21.5 1.42 17.3 0.46 3.61 2.54 4.85 3.30 0.19 9.8 1.59 0.53 1.01 255

66.82 0.79 15.00 5.59 0.05 2.45 0.45 1.81 3.79 0.21 2.70 99.66 66 435 199 40 93 123 29 16 8 37 210 17 7 16 1.3 1.8 14.0 18 3.5 44 93 10.7 42 8.4 1.3 6.9 1.2 6.7 1.2 3.5 0.5 3.4 0.5 1.32 1.26 15.0 1.26 22.6 0.43 3.13 2.47 5.06 3.27 0.20 9.5 1.66 0.51 1.02 223

64.82 0.81 15.76 6.12 0.08 2.44 0.48 1.83 4.22 0.21 3.40 100.17 65 458 188 48 98 123 30 14 5 30 196 17 7 16 1.2 2.0 15.0 16 3.7 24 77 6.1 25 5.2 0.9 4.5 0.9 5.1 1.0 3.1 0.5 3.3 0.5 1.26 1.00 13.1 1.07 24.7 0.55 1.60 1.49 4.35 2.89 0.21 5.4 1.11 0.55 1.52 157

65.15 0.81 15.70 5.96 0.04 2.46 0.41 1.71 4.06 0.21 3.10 99.61 66 480 218 36 98 157 30 15 18 33 212 19 8 17 1.3 1.7 15.0 19 4.2 20 59 5.5 22 5.0 0.8 4.3 0.9 5.4 1.1 3.4 0.6 3.6 0.5 1.61 1.10 14.1 1.27 22.9 0.52 1.32 1.04 4.52 2.48 0.23 4.1 0.97 0.50 1.38 132

67.35 0.73 16.00 5.24 0.04 1.86 0.89 1.94 4.43 0.16 1.20 99.84 62 756 206 239 101 198 20 12 14 34 265 15 8 17 1.2 1.7 14.0 13 3.6 30 71 8.0 34 7.4 1.4 6.3 1.0 5.8 1.1 3.3 0.6 3.8 0.5 1.96 1.75 18.9 0.96 16.5 0.48 2.14 2.23 3.72 2.53 0.22 5.8 1.35 0.60 1.09 174

63.05 0.85 18.26 6.17 0.02 2.28 0.30 1.09 3.98 0.13 3.70 99.83 73 599 227 58 105 178 23 9 13 36 228 12 7 20 1.5 10.8 17.0 20 3.8 32 71 7.9 33 7.3 1.2 6.2 1.1 5.9 1.3 3.4 0.6 4.1 0.5 1.69 1.52 13.4 1.19 22.3 0.57 1.88 1.58 5.32 2.74 0.22 5.8 1.23 0.52 1.05 175

62.50 0.30 19.47 2.86 0.03 1.41 4.10 5.07 1.69 0.71 1.20 99.34 52 221 101 254 25 116 10 4 1 138 162 7 7 20 1.7 4.4 13.0 23 9.1 72 161 17.8 75 18.1 2.8 19.4 3.2 18.5 3.8 10.7 1.7 10.7 1.4 4.65 14.09 12.4 1.75 11.1 0.14 5.55 3.17 2.51 2.50 0.24 5.0 1.46 0.45 1.05 416

77.73 0.50 10.49 3.14 0.03 0.92 1.88 2.46 1.17 0.13 0.80 99.25 55 103 82 162 48 178 10 5 10 27 284 8 9 11 1.0 2.2 8.0 13 3.1 32 65 7.0 29 5.9 1.3 5.2 0.8 4.4 0.9 2.4 0.4 2.8 0.4 3.71 2.68 35.5 1.61 10.6 0.43 4.04 2.50 4.16 3.43 0.20 8.6 1.51 0.71 1.00 158

73.63 0.48 11.54 4.73 0.03 1.53 1.48 3.21 1.02 0.18 1.30 99.13 56 63 74 207 56 287 21 11 20 33 191 6 6 12 1.0 2.9 9.0 11 2.9 40 80 8.6 36 7.7 1.5 7.0 1.0 5.6 1.0 2.9 0.5 3.0 0.4 5.13 1.55 21.2 1.24 15.1 0.37 4.44 3.57 3.86 3.26 0.22 9.7 1.87 0.61 0.99 195

77.38 0.49 10.59 3.56 0.03 1.01 0.70 1.50 2.14 0.10 1.60 99.10 64 368 116 113 54 171 15 7 14 31 234 9 7 11 0.8 4.9 8.0 11 3.7 53 102 10.7 42 8.2 1.4 6.6 0.9 4.9 1.0 2.8 0.5 2.8 0.4 3.17 2.05 29.2 1.39 12.6 0.36 6.64 4.78 3.00 4.06 0.19 14.0 1.90 0.55 0.99 238

394

U. ZIMMERMANN

Table 4. (continued) Outcrop Sample Grain-size (phi) SiO2 Ti02 A1203 Fe203T

% % %

MnO MgO CaO

% o/

Na2O

%

cy

%

K2O

%

P205

%

LOI

%

TOTAL

%

CIA Ba Rb Sr V Cr Ni Co Cu Y Zr Cs Hf Nb Ta Pb Sc Th U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Cr/V Y/Ni Zr/Sc Th/Sc Ti/Zr Nb/Y La/Sc La/Th Th/U Lajvf/Snijsf Sm/Nd LaN/YbN GdN/YbN Eu/Eu* Ce/Ce* ZREE

ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm

CEB 02-269 8

SIJ-P 1-2 8

76.68 0.57 11.28 3.87 0.03 1.10 1.08 1.83 2.13 0.14 0.90 99.61

64.05 0.66 16.64 6.14 0.13 2.64 0.21 1.36 4.57 0.11 2.90 99.41

61 334 120 127 57 239 14 7 8 30 303 10 9 13 1.1 3.7 9.0 12 3.5 37 74 8.1 34 7.1 1.4 5.8 0.9 5.0 1.0 2.8 0.5 3.0 0.4

SIJ-P 1-5 8 68.48 0.66 14.09 5.94 0.12 2.19 0.20 1.17 4.13 0.10 2.60 99.68

69 488 206 50 92 96 28 14 68 32 114 15 4 16 1.2

68 453 189 52 84 109 28 12 33 39 116 19 4 13 0.9

26.9 14.0

25.1 14.0

13 2.5 35 79 8.4 34 6.1 1.0 5.1 0.8 5.4 1.0 3.2 0.5 3.1 0.5

12 3.9 35 76 7.8 33 7.2 1.1 6.5 1.1 6.4 1.1 3.3 0.4 3.2 0.5

SIJ-P 1-6 8 62.06 0.75 16.95 6.55 0.08 2.56 0.34 1.62 4.25 0.14 4.10 99.40

68 455 195 54 106 103 31 14 15 33 129 35 4 15 1.2 4.5

SIJ-P 1-12 8 68.03 0.65 14.21 5.99 0.13 2.25 0.21 1.13 4.06 0.12 3.00 99.78

SIJ-P 1-13 8 69.91 0.61 13.36 5.75 0.13 2.46 0.18 0.64 3.91 0.13 2.70 99.78

SIJ-P 4-3 8 79.82 0.35 8.99 2.86 0.03 0.89 0.39 2.27 1.53 0.14 1.44 98.72

60 225 54 68 48 20 11 4 12 15 89 4 3 7 0.7 6.4 4.5 7 1.7 22 43 5.5 22 4.2 0.8 3.6 0.5 3.3 0.7 1.8 0.3 1.7 0.3

SIJ-P 4-4 8 76.78 0.41 10.61 2.80 0.06 0.84 0.92 2.28 2.15 0.15 1.65 98.65

70 424 193 39 85 137 28 11 46 38 110 19 3 13 1.0

17.0

21.8 12.0

13.8 11.0

15 6.3 31 66 7.3 30 5.9 1.1 4.9 0.9 5.5 1.1 3.0 0.5 3.4 0.5

13 4.7 39 77 8.7 35 7.4 1.2 6.4 1.2 6.6 1.2 3.4 0.5 3.0 0.5

11 3.4 32 63 7.2 31 6.9 1.1 7.3 1.2 6.4 1.1 3.1 0.5 2.6 0.4 1.61 1.37 10.0 0.99 33.2 0.34 2.94 2.96 3.21 2.93 0.22

0.42 1.36 19.6 1.58 23.8 0.48 4.83 3.06 4.22 3.24 0.19

0.47 1.54 17.8 1.44 22.9 0.48 4.41 3.06 4.84 3.47 0.19

0.45 1.38 16.6 1.51 26.0 0.46 4.32 2.86 4.55 3.39 0.19

58 294 80 65 60 28 12 6 7 18 107 6 3 9 0.8 6.7 6.0 9 1.8 27 52 6.3 26 4.8 0.9 4.2 0.6 3.6 0.7 2.0 0.3 2.1 0.3

8.2

8.3

7.6

0.89 34.6 0.50 2.51 2.82 5.00 3.62 0.18

0.87 34.0 0.33 2.51 2.89 3.13 3.06 0.22

0.89 34.9 0.46 1.83 2.06 2.40 3.30 0.20

1.02 1.38 10.4 1.07 31.2 0.39 3.24 3.04 2.72 3.29 0.21

9.0

8.2

8.1

6.8

9.6

9.0

9.5

9.6

1.57 0.65 0.98

1.31 0.54 1.09

1.63 0.47 1.06

1.18 0.62 1.03

1.73 0.51 0.98

2.25 0.44 0.95

1.73 0.58 0.94

1.67 0.58 0.94

180

183

1.30 1.42

182

80.08 0.35 9.29 2.23 0.07 0.86 0.28 2.47 1.70 0.13 1.33 98.80

68 440 202 52 87 89 28 13 30 38 125 21 3 15 1.1

4.20 2.11 33.7 1.36 11.3 0.42 4.10 3.02 3.49 3.26 0.21

1.04 1.12

SIJ-P 4-5 8

0.97 1.05

161

191

164

110

130

59 214 57 67 45 20 11 5 7 15 82 5 2 7 0.7 8.5 4.9 7 1.6 21 44 5.1 21 3.9 0.7 3.6 0.5 3.0 0.6 1.7 0.3 1.6 0.2

SIJ-P 4-6 8 66.32 0.74 16.17 4.53 0.06 1.53 0.38 1.39 4.84 0.17 2.73 98.84

66 617 189 56 108 56 24 11 22 24 129 11 4 15 1.3 16.6 13.3

13 2.8 35 68 8.5 33 6.4 1.2 5.5 0.8 4.7 0.9 2.6 0.4 2.7 0.4

SIJ-P 4-7 9-Jan

SIJ-G 67a 1

81.35 0.32 8.75 2.06 0.04 0.64 0.32 2.42 1.45 0.12 1.40 98.87

85.98 0.59 6.63 3.21 0.01 0.27 0.04 0.03 1.73 0.03 1.30 99.82

59 184 51 64 43 18 9 5 6 15 81 4 2 7 0.6 9.5 4.5 7 1.5 18 38 4.6 18 3.5 0.6 3.1 0.5 2.9 0.6 1.6 0.2 1.6 0.2

77 228 79 45 64 260 8 6 3 31 237 3 7 10 0.7 4.0 7.0 10 2.5 33 92 9.1 35 7.2 1.2 5.8 0.9 5.2 0.9 2.8 0.4 2.6 0.4

0.98 34.5 0.63 2.62 2.66 4.72 3.43 0.19

0.43 1.62 17.8 1.47 23.9 0.44 4.08 2.77 4.32 3.31 0.19

4.06 3.67 33.8 1.47 14.9 0.31 4.74 3.22 4.12 2.89 0.21

9.7

9.6

8.7

9.4

1.79 0.59 1.00

1.66 0.58 0.95

1.63 0.56 0.99

1.80 0.56 1.30

107

0.52 1.00

9.7

170

Measured by ICP-MS at ACME, Laboratories Vancouver (Canada). See Fig. 1 for locality abbreviations.

94

197

395

THE PUNCOVISCANA COMPLEX Table 4. (continued) SIJ-G SIJ-G 67c 67b 1 0 87.13 0.36 5.56 3.63 0.01 0.20 0.03 0.02 1.51 0.03 1.30 99.78 76 168 70 39 62 315 9 18 4 29 116 2 4 7 0.6 3.1 6.0 6 2.1 36 93 10.1 40 8.0 1.2 5.5 1.0 5.4 0.9 2.4 0.4 2.3 0.3 5.08 3.08 19.4 1.03 18.5 0.23 5.92 5.73 2.95 2.78 0.20 11.5 1.95 0.54 1.20 206

87.58 0.38 5.78 3.04 0.01 0.30 0.02 0.01 1.54 0.02 1.00 99.68 77 183 75 47 60 226 9 15 3 32 124 3 4 7 0.5 2.8 6.0 7 1.8 43 95 10.6 42 8.7 1.4 6.1 1.1 5.4 1.0 2.5 0.4 2.3 0.4 3.76 3.39 20.6 1.18 18.4 0.23 7.08 5.99 3.94 3.06 0.21 13.8 2.17 0.55 1.08 219

SIJ-G SIJ-G 67e 67d 1 1

SIJ-G 67g 2-Jan

CON Ql 8

CON Q4 8

CON Q7 5

CON Q8 1

CON Q9 2

CON Qll 8

NP NP NP CON Q14 NP270 NP271 NP272 8 8 5 8

86.94 0.38 5.82 3.10 0.01 0.34 0.02 0.03 1.50 0.03 1.30 99.47 77 177 69 43 54 239 11 12 3 32 117 3 4 7 0.6 3.1 6.0 7 1.7 42 98 10.5 42 8.6 1.4 6.1 1.1 5.9 1.0 2.9 0.4 2.5 0.3 4.43 2.94 19.6 1.17 19.4 0.21 6.95 5.96 4.12 3.04 0.21 12.5 2.02 0.58 1.12 222

87.98 0.42 5.30 2.82 0.01 0.35 0.03 0.04 1.35 0.02 1.20 99.52 77 166 66 43 55 260 14 14 4 28 156 3 5 8 0.6 3.7 5.0 8 1.8 32 76 8.5 33 6.8 1.1 5.6 0.9 4.9 0.9 2.5 0.4 2.3 0.3 4.73 2.10 31.2 1.64 16.1 0.27 6.48 3.95 4.56 2.99 0.21 10.6 2.02 0.55 1.12 176

70.63 0.66 13.10 4.74 0.07 2.44 1.67 2.34 2.35 0.19 1.60 99.79 58 307 122 169 89 68 24 14 23 33 171 69 5 14 1.0 3.2 9.0 11 2.0 28 53 6.8 30 6.0 1.2 5.4 1.0 5.3 1.1 3.1 0.5 3.0 0.4 0.77 1.40 19.0 1.19 23.1 0.43 3.16 2.65 5.35 2.97 0.20 7.0 1.48 0.63 0.88 145

61.02 0.82 16.52 7.12 0.13 3.24 1.50 1.82 4.69 0.19 2.70 99.75 60 679 209 121 153 144 43 20 8 35 152 57 5 15 1.1 6.9 14.0 12 2.2 21 79 5.1 22 4.9 1.0 4.5 0.9 5.4 1.1 3.3 0.5 3.1 0.5 0.94 0.82 10.9 0.83 32.3 0.44 1.46 1.77 5.27 2.62 0.22 4.9 1.17 0.62 1.80 151

55.02 0.94 19.31 8.55 0.12 4.18 1.16 1.16 5.89 0.31 3.00 99.64 65 1432 239 96 179 103 49 24 25 38 153 60 5 19 1.4 2.9 19.0 17 2.7 37 80 8.5 35 6.4 1.3 5.7 1.1 6.2 1.3 4.1 0.7 4.1 0.6 0.57 0.78 8.0 0.87 37.0 0.49 1.95 2.25 6.11 3.63 0.19 6.7 1.14 0.62 1.05 191

74.16 0.73 11.33 4.02 0.08 1.67 2.52 2.77 1.50 0.21 0.80 99.79 51 155 83 191 77 55 18 9 28 38 338 38 10 13 1.0 2.9 11.0 12 2.2 30 65 7.3 31 6.7 1.3 6.1 1.0 5.9 1.2 3.6 0.6 3.5 0.6 0.71 2.09 30.8 1.07 12.9 0.34 2.75 2.56 5.36 2.82 0.21 6.4 1.44 0.59 1.02 164

70.88 0.69 13.01 4.90 0.07 2.26 1.26 1.60 3.17 0.21 1.70 99.75 61 569 140 123 99 55 25 14 8 31 217 60 6 14 1.0 2.8 10.0 10 1.8 28 61 6.7 28 5.4 1.2 5.0 0.9 5.3 1.0 2.9 0.5 2.9 0.4 0.55 1.23 21.7 1.03 19.1 0.44 2.84 2.76 5.72 3.30 0.19 7.3 1.41 0.68 1.03 150

65.18 0.83 15.45 6.03 0.10 2.99 1.50 1.92 3.73 0.22 1.80 99.75 61 531 175 131 117 130 35 18 34 35 183 92 5 15 1.1 3.5 12.0 11 2.0 30 65 7.0 31 6.1 1.3 5.3 0.9 5.3 1.2 3.4 0.5 3.2 0.5 1.11 0.99 15.3 0.88 27.1 0.44 2.48 2.84 5.25 3.06 0.20 6.9 1.36 0.67 1.03 160

70.06 0.98 12.70 5.19 0.11 2.04 2.45 2.79 2.09 0.29 1.00 99.70 53 526 106 187 96 164 25 12 14 48 438 45 13 17 1.3 2.7 15.0 19 3.2 51 108 12.1 52 9.7 1.9 8.6 1.5 8.1 1.7 4.8 0.7 4.6 0.7 1.71 1.97 29.2 1.26 13.4 0.35 3.37 2.67 5.91 3.26 0.19 8.0 1.50 0.60 1.01 265

83.81 0.63 7.85 3.35 0.01 0.48 0.04 0.04 2.04 0.02 1.40 99.67 77 252 93 23 71 308 16 18 4 44 239 4 7 12 0.9 3.4 8.0 9 2.5 43 104 11.1 46 9.3 1.4 7.3 1.3 7.7 1.3 3.9 0.6 3.6 0.5 4.34 2.86 29.9 1.15 15.8 0.26 5.31 4.62 3.68 2.86 0.20 8.8 1.67 0.50 1.14 240

65.88 0.88 15.05 6.81 2.46 0.12 0.34 1.42 3.18 0.13 3.40 99.67 70 777 152 67 86 55 36 17 10 39 265 9 7 19 1.4 3.0 13.0 15 2.9 44 91 10.2 40 8.2 1.2 7.4 1.1 6.9 1.3 4.2 0.6 4.3 0.6 0.64 1.08 20.4 1.12 19.9 0.49 3.39 3.02 5.03 3.37 0.21 7.6 1.42 0.45 1.02 221

73.96 79.74 0.74 0.59 12.35 10.20 2.61 4.23 1.53 0.90 0.06 0.09 0.35 0.41 2.25 1.90 1.44 2.12 0.12 0.08 2.40 1.70 99.85 99.92 67 63 608 289 84 103 41 66 47 66 62 62 18 25 8 12 8 7 36 39 239 327 3 5 6 9 12 16 1.0 1.3 1.1 1.7 9.0 11.0 11 13 2.5 4.0 33 40 72 85 7.6 9.1 30 35 6.2 7.4 1.0 1.2 6.4 6.9 1.1 1.0 6.3 5.8 1.2 1.1 3.4 3.9 0.5 0.6 3.2 3.8 0.6 0.5 1.31 0.93 2.05 1.57 26.5 29.7 1.17 1.14 14.8 13.6 0.32 0.41 3.68 3.63 3.15 3.19 4.20 3.13 3.35 3.38 0.21 0.21 7.6 7.8 1.62 1.49 0.47 0.49 1.07 1.05 172 202

U. ZIMMERMANN

396

Table 4. (continued) Outcrop Sample Grain-size (phi) Si02 TiO2 A1203 Fe203T MnO MgO CaO Na2O K2O P205 LOI TOTAL CIA Ba Rb Sr V Cr Ni Co Cu Y Zr Cs Hf Nb Ta Pb Sc Th U La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Cr/V Y/Ni Zr/Sc Th/Sc Ti/Zr Nb/Y La/Sc La/Th Th/U LaN/Snijvi Sm/Nd LaN/YbN GdN/YbN Eu/Eu* Ce/Ce* ZREE

% % % % % % % % (V

%

% %

ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm ppm

NP NP273 8

NP NP NP NP NP NP274 NP276 NP278 NP280 NP281 5 5 8 5 8

59.09 77.53 0.58 1.05 10.25 18.80 4.35 8.09 1.42 2.62 0.09 0.11 0.35 0.53 1.35 2.25 1.73 3.62 0.07 0.13 2.20 3.50 99.79 99.92 68 69 308 737 80 166 44 101 46 123 48 82 24 42 21 13 22 3 32 44 224 195 3 6 6 5 12 20 1.0 1.4 1.1 1.7 9.0 20.0 10 17 2.2 4.0 33 61 69 121 7.3 13.1 28 52 6.3 9.4 1.7 0.9 5.6 8.5 0.9 1.3 5.0 7.3 1.1 1.5 3.2 4.7 0.4 0.7 2.7 4.9 0.4 0.7 1.04 0.67 1.37 1.06 24.9 9.7 1.06 0.83 15.5 32.4 0.36 0.45 3.63 3.06 3.44 3.71 4.32 4.13 3.25 4.08 0.22 0.18 9.0 9.3 1.42 1.71 0.45 0.56 1.06 1.00 287 163

61.13 74.18 0.64 0.96 17.08 11.83 5.07 7.43 1.67 2.69 0.06 0.10 0.43 1.47 2.04 3.63 1.68 2.19 0.12 0.14 2.20 2.90 99.72 99.92 67 61 306 307 72 133 49 178 60 93 62 68 42 28 17 18 13 1 32 39 210 186 2 5 6 5 14 16 1.0 1.2 1.2 4.4 10.0 16.0 11 16 2.7 3.2 36 47 78 93 8.3 10.2 32 37 6.7 7.5 1.2 1.5 6.1 7.0 0.9 1.1 5.7 6.4 1.1 1.3 3.3 4.0 0.5 0.6 3.2 3.6 0.5 0.6 1.03 0.74 1.16 0.92 21.0 11.6 1.05 0.98 18.2 30.9 0.42 0.41 3.59 2.95 3.42 3.01 3.89 4.91 3.94 3.36 0.21 0.20 8.3 9.6 1.57 1.57 0.55 0.63 1.02 1.09 183 222

50.97 1.64 24.34 6.76 2.38 0.12 0.67 2.08 6.56 0.24 4.00 99.76 67 1122 296 92 166 96 42 19 1 55 292 6 8 28 2.2 1.6 27.0 23 5.0 83 169 17.7 67 13.8 2.3 11.6 1.7 9.3 1.8 5.6 0.8 6.1 0.9 0.58 1.29 10.8 0.85 33.7 0.51 3.08 3.63 4.58 3.77 0.21 10.1 1.56 0.53 1.04 391

LA PI 3

51.05 82.52 0.33 1.63 6.31 24.32 5.05 6.78 0.06 2.38 0.12 1.67 0.24 0.67 1.71 2.07 0.08 6.49 0.07 0.23 1.90 4.00 99.74 99.94 65 68 1114 26 4 302 54 94 35 166 103 239 20 42 9 20 20 1 28 57 294 203 0 6 6 8 7 27 2.0 0.7 3.3 1.5 4.0 27.0 8 25 2.1 4.6 17 86 39 171 5.0 18.0 22 66 5.1 13.3 0.7 2.3 12.2 5.0 0.9 1.7 4.6 9.5 0.9 1.8 2.7 5.7 0.4 0.9 2.3 6.0 0.9 0.3 6.84 0.62 1.40 1.36 50.7 10.9 2.00 0.92 33.2 9.8 0.26 0.48 4.15 3.18 2.08 3.45 5.41 3.81 2.04 4.05 0.23 0.20 5.4 10.6 1.81 1.66 0.41 0.54 1.04 1.02 106 395

LA P2 8

LA P7 5

62.22 0.78 17.05 7.47 0.08 2.78 0.52 1.85 3.12 0.09 3.80 99.76 70 591 138 79 126 116 26 11 89 22 126 4 4 16 1.3 1.8 15.0 14 6.5 25 39 6.1 26 4.7 0.9 3.7 0.6 3.6 0.7 2.2 0.4 2.4 0.3 0.92 0.83 8.4 0.91 37.0 0.76 1.65 1.82 2.09 3.31 0.18 7.5 1.24 0.65 0.75 115

66.08 0.64 15.07 7.57 0.11 3.16 0.33 1.44 2.57 0.10 2.90 99.97 72 461 119 54 94 103 40 14 8 33 154 3 5 14 1.2 2.1 15.0 14 3.6 34 69 8.5 36 6.8 1.2 6.0 1.1 5.4 1.0 3.4 0.5 3.0 0.4 1.09 0.83 10.2 0.93 25.0 0.42 2.27 2.43 3.89 3.13 0.19 8.4 1.62 0.54 0.95 175

LA P8 8

64.26 0.74 17.07 6.10 0.08 2.43 0.22 0.81 4.12 0.11 3.80 99.74 73 612 180 40 123 109 21 4 58 25 122 4 4 15 1.2 69.6 14.0 14 5.0 32 78 10.1 43 7.7 1.3 5.0 0.8 4.2 0.9 2.7 0.4 2.7 0.4 0.89 1.23 8.8 1.10 35.6 0.44 3.44 3.13 3.85 4.03 0.18 11.7 1.51 0.87 0.88 189

Measured by ICP-MS at ACME, Laboratories Vancouver (Canada). See Fig. 1 for locality abbreviations.

LA P12 8

60.76 0.73 17.53 8.04 0.10 3.25 0.80 2.76 2.50 0.11 3.20 99.78 67 470 116 114 126 109 44 16 1 35 123 3 4 15 1.1 2.9 14.0 15 4.0 48 85 10.1 42 7.5 2.0 6.3 1.1 5.6 1.1 3.2 0.5 3.0 0.5 0.87 0.79 8.7 1.01 36.4 0.58 2.29 2.28 2.82 2.61 0.18 8.7 1.67 0.63 1.05 216

THE PUNCOVISCANA COMPLEX

Sampling areas and outcrop description Sample areas for provenance study were distributed over the entire Puncoviscana Basin in northwestern Argentina (Fig. 1). The northwestern and western part are represented by samples from the Puna. The eastern part of the basin is covered by sampling areas between Jujuy and Tucuman, referred to as 'Cordillera Oriental'. The southern part is divided into two areas. The first, 'Ambato', is situated in the Central batholith belt (Fig. 2b; after Pankhurst & Rapela 1998), where mainly medium- to high-grade metamorphic rocks of unknown age and post-depositional plutonic igneous rocks are associated with outcrops of the Puncoviscana complex and includes the Sierra de Ovejerua with the outcrop Sundro (SUN). The second is located in the Sierra de Famatina, and referred to here as Famatina (Fig. 1, Table 5). A compilation of the characteristics of the sampling areas is given in Table 5. The aim of sampling was to collect rocks from welldescribed sequences of the Puncoviscana complex, but several new outcrops were found and sampled. All sampled rocks exhibit the typical pre-Ordovician deformation described by Toselli (1990), Willner (1990) and Mon & Hongn (1991). Regional metamorphism of these rocks does not exceed lower greenschist facies (Toselli & Rossi de Toselli 1982; Do Campo 1999), with temperatures below 350 °C and 3 kbar estimated by Toselli (1990). Only rocks whose original principal textures were not changed dramatically during metamorphism and/or deformation were analysed. No gneisses, schists or migmatites were selected.

Region 1: Puna 1.

2.

3.

Quebrada del Volcan (VOL). The rocks exposed in the Quebrada del Volcan at the western border of the Salar Antofalla (S25°39'23.5"; W67°47'24.4") represent the most westerly exposed Puncoviscana Basin deposits. The outcrop dimensions are 150 X 500 m and there is a tectonic contact with highly deformed metagabbros. Quebrada Randolfo (RAN). These rocks are exposed in several outcrops of 50-200 m along the Quebrada Randolfo (S26°51'll"; W66°44'22.6") and are cut by several felsic to intermediate intrusive rocks of unknown age. Rio Taique (RT). Low-grade metasedimentary rocks are exposed with an unconformable contact with Lower Tremadocian quartz arenites (S23°50'29.3";

397

W66°15'47.5"). The outcrop occupies an area of about 200 X 500 m.

Region 2: Cordillera Oriental 1.

2.

3.

4.

5.

Quebrada El Toro, Cuesta Munano and Purmamarca (TOR). Samples were taken from the 'classic' Puncoviscana Formation outcrops in the Quebrada Purmamarca (S23°42'0.5"; W65°39'21.5") and along the road to the west towards San Antonio de los Cobres. Rio Corralito (COR). Samples were taken along the river (S24°59'06.5"; W65°41'42.0"), below a conglomeratic layer (Durand 1990) reinterpreted as a pebble-rich mud flow (van Staden & Zimmermann 2004). La Pedrera (PED). The outcrop lies 5 km south of Salta towards the Cuesta de Quesera (S24°52'22.9"; W65°20'46.2"). The rocks are overlain discordantly by the Meson Group. Seclantes and Molinos (SEC, MOL). Both outcrops (Molinos S25°26'06.8"; W66°17'10.0"; Seclantes 1 km to the north of the town) are located in the Valle Calchaquies south of Cachi, and situated close to intrusive bodies. Choromoro (CHO). The outcrop is to the west of Choromoro (S26°23'25.4"; W65°28'12.6"). The locality has been the subject of intensive sedimentological study (see Acenolaza et al 1990; Acenolaza 2004).

Region 3: Ambato 1.

2.

3.

4.

Quebrada de Suncho (SUN). The rock succession in the Quebrada de Suncho (Sierra de Ovejeria, S27°25'55.9"; W66°38'08.3") comprises more than 300 m of metaconglomerates, embedded in a thick packet (>1 km) of metasedimentary rocks. Quebrada La Cebila (CEB). The rocks in this c. 30 km long valley (sample at S28°49'50.8"; W66°24'9.3") are described in detail by Hockenreiner (1998). Sijan (SIJ-P). In the region between Saujil and Poman on the western flank of the Sierra de Ambato to the east of Sijan (S28°12'-S28°17'; W66°08'-W66°09'), outcrops of the Puncoviscana complex are exposed as windows in a giant modern alluvial fan. Contacts with the country plutonic rocks, migmatites and gneisses are not exposed. Sijan (SU-G). A packet (20 m) of red, hard

Table 5. Summary of the geological characteristics of the different outcrops Outcrop Puna Quebrada del Volcan Quebrada Randolfo Rio Taique Cordillera Oriental El Toro Mufiano Permamarca

Lithology

red slaty semi pelites with silty to clayey grain-sizes black slaty clayey pelites with few intercalations of black to blue psammites brown to green slaty pelites with muddy to very fine sand grain sizes with a few intercalations of slaty psammites brown to green metalitharenitic wackes of mud to fine sand grain sizes intercalated with clayey metapelites

Rio Corralito

brown muddy to fine-sand pelites and semi-pelites (metawackes)

La Pedrera

brown to green muddy to fine-sand pelites and semi-pelites (metawackes)

Seclantes Molinos

grey to brown fine-sand semi-pelites and psammites (metawackes) intercalated with clayey pelites: Scelantes psammites can be described as phyllites brown to green fine-sand semi-pelites and psammites (metawackes) intercalated with clayey pelites

Choromoro

Ambato Quebrada de Suncho Quebrada La Cebila SijanPelites SijanPsammites Conception

Famatina Quebrada de Paiman Los Corrales

brown to green slaty silty to sandy pelites and semi-pelites with intercalated sandy psammites (metawackes) brown to grey sandy psammites (metawackes) with intercalated slaty muddy pelites and phyllitic pelites blue, red, brown and green clayey pelites with rare intercalations of blueish finesand pelites red-brown quartz-rich medium- to coarsesand psammites blue medium-sand semi-pelites and psammites (metawakes) with intercalated clayey pelites

blue to brown muddy pelites to finesandy psammites brown muddy pelites to semi-pelites to psammites

Petrology-Mineralogy

Sedimentology

Tectonics

rounded Qp, sub-rounded undulose Qm, altered kfs and alb (P/F=0,1), sub-rounded mms L, mus, zir, kao (XRD),hem, hbl; M (> 15%): qtz, alb, ill, ser, chl sub-rounded to rounded Qm partly undulose, altered alb (EDS), mus, bio, ilm (EDS), gla; M (5-35%): qtz, alb, kfs, mus poorly sorted; small rounded Qm (< 60 um), angular large Qm (60-100mm); altered kfs, alb, oli (MP), mms L, mm L ('phyllites'); M (< 20%): qtz, alb, kfs, ill, chl, mus

fine laminated, with cross-beddings, small-scale channels of silty semi-pelites incising clayey semi-pelites and pelites normal grading, laminations, but mainly massive rocks without sedimentary structures normal grading, laminations

polyphase deformation, strongly folded with kink-bands and isoclinal small-scale folds polyphase deformation with two cleavages developed polyphase deformation with two cleavages developed, folded and slaty aspect in the field

poorly sorted; sub-rounded to sub-angular mainly undulose Qm, Qp small rounded qtz (< 70 um), angular large qtz (70-300 um), strongly altered F mainly sub-euhedral alk (< 300 um), alb, small (<60 um) sub-anhedral oli (MP), silty mms L, mus, bio, gla (XRD); M (<25%) qtz, alb, kfs, ill, chl, and probable braunite poorly sorted; sub-rounded to sub-angular mainly undulose Qm, Qp small rounded qtz (< 70 m), euhedral Qm (80-120 um) altered F (P/F= 0.2-0.4), mms L, mus, bio; M (<25%): ill, qtz, kfs, alb, chl, apa poorly sorted; sub-rounded to sub-angular undulose Qm, Qp; F is strongly altered, kfs, alb, oli; M (<20%): qtz, hlb, alb, mus, ill, kao, mmt, chl, ser, rut, gla, hem poorly sorted; sub-angular to sub-rounded undulose Qm (80-300 um) sub-rounded Qp; strongly altered F (P/F= 0.2), kfs, alb, sed L, mms L, mm L ('phyllites'), mus, bio; M (5-25%): qtz, kfs, alb, ill, chl, mus (Seclantes) - qtz, alb, bio, P (Molinos) poorly sorted; sub-angular undulose Qm and rarely Qp, strongly altered angular F (P/F< 0.4), kfs, alb, silty to fine-sand sed L, psammitic mms L, mm L ('phyllites'), mus, cha (XRD), hem; M (8-18%): qtz, alb, ill, chl, mus

normal grading, laminations, crossbedding, slumping, ripple marks, load marks, flute-casts, trace fossils

polyphase deformation with two cleavages developed, folded

thin-laminated, cross-bedding, trace fossils

polyphase deformation with two cleavages developed, strongly folded polyphase deformation with two cleavages developed, folded

poorly sorted; sub-angular undulose Qm, Qp, strongly altered angular F (P/F= 0.1-0.3), kfs, alb, mus, bio, hem, tit, zir, apa; M (<18%): qtz, alb, kfs, chl, mus and bio poorly sorted; sub-angular undulose Qm, Qp, strongly altered angular F (P/F= < 0.5), kfs, alb, mus, rounded zir, apa, chl; M (25%): qtz, alb, kfs, chl, mus and bio sub-angular to sub-rounded undulose Qm, rarely Qp, altered subangular F (P/F< 0.2), mic, san, alb, silty sed L, mm L, mus, bio, zir, hem, hbl; M (15-30%): qtz, alb, kfs, ill, mmt, chl, mus moderate sorting; sub-angular to sub-rounded undulose Qm, altered sub-angular large kfs (200-350 urn), smaller alb (150-200 um) rounded muddy to silty sed L, hem; M (5-18%): qtz, alb, kfs, ill, mmt, chl poorly sorting; angular to sub-rounded undulose Qm partly bi-modally distributed, sub-angular large kfs (200-300 um), san, mic, smaller alb (50-100 um) partly altered to calcite, rounded silty qtz-rich sed L, mm L ('phyllites'), mms L; bio, mus, hbl, pyr, ilm, apa, hem, rut, well rounded euhedral zir; M (15-25%): qtz, alb, kfs, ill, mus, chl, epi, tur

laminated, flute marks, trace fossils

poorly sorting; angular to sub-rounded undulose Qm, altered F (P/F 0.1-0.4), kfs, alb, mms L, mmL ('phyllites'),epi, bio, well-rounded zir and rut, gla (XRD); M (<35%): qtz, alb, kfs, ill, mmt, chl, ser, mus, epi, tur (description for both outcrops)

laminated, cross-bedding, normal grading laminated, normal grading, flute marks, slumping, trace fossils laminated, normal grading, flute marks, slumping, trace fossils

laminated, flute marks, cross-bedding, normal grading, slumping, convolute bedding, trace fossils laminated, flute marks, cross-bedding, normal grading, slumping partly laminated, normal grading

polyphase deformation with two cleavages developed, strongly folded; both outcrops are situated close to intrusive complexes polyphase deformation with two cleavages developed, folded

polyphase deformation with two cleavages developed, strongly folded; partly phyllites polyphase deformation with two cleavages developed, strongly folded; phyllites and schists polyphase deformation, strongly folded with kink-bands and isoclinal small-scale folds polyphase deformation with two cleavages developed, folded

thin-laminated, normal grading, convolute bedding, cross-bedding, Bouma sequence Tt,_d, ripples

polyphase deformation with two cleavages developed, penetrated by abundant quartz veins

thin-laminated, normal grading, trace fossils

polyphase deformation with two cleavages developed, strongly folded, penetrated by abundant quartz veins

thin-laminated, normal grading

The localities are shown in Figure 1. Abbreviations: qtz, quartz; Qm, monocrystalline quartz; Qp, polycrystalline quartz; F, feldspar; kfs, potassium feldspar; mic, microcline; oli, oligoclase; san, sanidine; P, plagioclase; alb, albite; L, lithoclasts; mm, metamorphic; mms, metasedimentary; sed, sedimentary; mus, muscovite; bio, biotite; chl, chlorite; ill, illite; kao, kaolinite; mmt, montmorillionite; ser, sericite; epi, epidote; cha, chamosite; hem, hematite; ilm, ilmenite; gla, glauconite; zir, zircon; apa, apatite; hbl, hornblende; tur, tourmaline; rut, rutile; tit, titanite; pyr, pyrite; M, matrix; MP, microprobe; XRD, x-ray diffraction; EDS, energy dispersive system. Matrix content was determined by XRD, EDS and microprobe.

THE PUNCOVISCANA COMPLEX

5.

metawackes is exposed to the east of Colana (S28°22'9.8"; W66°8'20.7")- The rocks concordantly overlie a quartz-rich conglomerate, which, in turn, overlies the rocks described above (SIJ-P). Conception (CON). The outcrop lies 50 km south of San Fernando del Valle de Catamarca, next to El Quemadito (S28°38'59.5"; W66°03'14.4") and is surrounded by migmatites and gneisses.

Region 4: Famatina The Famatina region has been described by several authors (see compilation in Acenolaza et al 1996) as a distinct tectonic unit separated from the Puna region to the north and the Sierras Pampeanas to the east (Fig. 2b). Two pre-Ordovician formations, La Aguadita (LA) and Negro Peinado (NP), crop out on the eastern flank of the Sierra de Famatina, and were sampled in well-described localities at Los Corrales (NP; S28°51'45.3"; W67°35'55") and Quebrada de Paiman (LA).

Provenance Provenance studies on sedimentary rocks aim at deciphering both the composition and geological evolution of their source areas, and the tectonic setting of the depositional basin. Meaningful results depend on the evaluation of indicators that mirror the original composition of source rocks and areas; data need to be tested for the effects of secondary factors that have the potential to obscure this information, such as sorting, weathering, metamorphism and further alteration processes (e.g. Morton & Hallsworth 1999). However, provenance studies that rely only on petrographic studies can lead to misinterpretations, and combination with major and trace element geochemistry is necessary to obtain reliable information (e.g. Condie & Martell 1983; von Eynatten et al. 2003).

Petrography Quantitative petrography based on the GazziDickinson method ideally requires clastic rocks of a certain maturity, grain size and composition (Ingersoll et al. 1984). For about 75% of the rocks of the Puncoviscana complex these preconditions are not met. First, most of the rocks are too fine grained (< 50 um). Secondly, they are mainly poorly sorted, making pointcounting at one step-size impossible. Thirdly, the matrix content lies between 8% and 35% and is, in many samples, too high (> 20%) for

399

point-counting according to several authors (e.g. Ingersoll et al. 1984; Cox & Lowe 1996). Furthermore, framework minerals such as feldspars and lithic components are weathered strongly and sometimes impossible to classify. The metasedimentary rocks of the Puncoviscana complex are dominated by quartz grains, mainly monocrystalline, with a low abundance of plagioclase. Lithoclasts are mainly of (meta) sedimentary origin, but fragments of phyllites and schists occur in some outcrops. Generally, quartz and alkali-feldspar grains are larger than the subhedral to euhedral grains of plagioclase. Alteration affected the feldspars intensely, commonly causing sericitization and sometimes calcification. The grains are mostly angular and seldom rounded. New growth on quartz grains, which could mask the probable rounded nature of clasts, was not observed using BSE and CL analysis. The matrix is rich in phyllosilicates, quartz and feldspar, and mainly consists of weathered and dissolved lithoclasts, the origin of which is difficult to determine. Most rocks are affected by hydrothermal alteration introducing chlorite and epidote (van Staden & Zimmermann 2004). The strong deformation is often marked by orientated new growth, mainly of muscovite, as determined by microprobe and EDS. Compared to the data of Jezek (1990), a lower abundance of lithoclasts and plagioclase, and volcanic grains in the form of sanidine are observed: Jezek (1990) and Jezek & Miller (1986) did not mention volcanic input nor record matrix percentages. The observed high proportion of pseudomatrix (Table 5) is ascribed to the dissolution of the less stable lithoclasts such as very fine-grained sedimentary components and volcanic lithoclasts, as described elsewhere (Dickinson 1970; Condie & Martell 1983; von Eynatten et al. 2003). However, petrographic indicators of volcanic input except sanidine, were not observed. The rocks show neither recrystallized glass fragments, clayey (silicified) volcanic ash clasts, nor quartz or feldspar grains with resorption embay ments produced by rapid cooling, as described by Schneider (1993), Dutta & Wheat (1990) and Zimmermann & Bahlburg (2003). The data would indicate a recycled orogen source in Dickinson diagrams (QtFL, QmFLt, QpLvLs, LvLmLs, QmPK).

Implications of the petrographic data Derivation from a recycled orogenic source, composed mainly of metamorphic and sedimentary rocks contrasts with interpretation of Jezek (1990), who reported an extremely high

400

U. ZIMMERMANN

plagioclase content suggesting a dissected volcanic arc source in the Qm-P-K diagram (after Dickinson 1988). Jezek & Miller (1986) and Jezek (1990) also considered the Puncoviscana complex to comprise polycyclic reworked material on a passive continental margin. However, compared to typical recent or older (Ordovician) continental margin deposits (Table 1), the rocks of the Puncoviscana complex are enriched strongly in lithoclasts and plagioclase (feldspar), and the poor sorting and the mainly subhedral grain shape do not necessarily indicate such a deposit. The closest comparison in terms of mineralogical composition is to foreland-basin infill (young as well as older, Table 1). In such settings, quartz dominates over mainly alkali-feldspar and largely metamorphic and sedimentary lithoclasts, whereas plagioclase is subordinate (Valloni & Zuffa 1984; DeCelles & Hertel 1989; Bahlburg 1990; Espejo & Lopez-Gamundi 1994; Zimmermann et al 2002). However, the amount of volcanic lithoclasts in foreland-basin deposits can vary according to the weathering profile and the proximity of the volcanic arc, and the foreland basin can be situated peripherally on the subducted plate, or behind the arc on the overriding plate (Jordan 1995). The basin morphology, determined by active tectonism, can allow funnelling of volcanic debris into the foreland basin (DeCelles & Hertel 1989; Zuffa et al. 1980; Cibin et al. 2001). The deposited labile volcanic lithoclasts can then be weathered and dissolved to pseudomatrix (e.g. Condie & Martell 1983; von Eynatten et al 2003), as interpreted here. As the pseudomatrix proportion is generally very high, the absence of volcanic lithoclasts cannot be used to exclude a volcanic source terrane. The sedimentary rocks of the Puncoviscana complex were deposited mainly by means of turbidity currents; rapid burial probably took place, as only minor reworking occurred. An important characteristic of the petrography of the Puncoviscana complex is the relatively uniform lithotype distribution throughout the depositional area. Apart from a few conglomerates and carbonates, deposited mainly in the eastern part of the basin, rocks from the different basin areas are petrographically nearly indistinguishable from each other. Compressive tectonics during Palaeozoic and Tertiary times shortened the crust by more than 40% (Kley et al. 1997; Kley 1998; Killer & Oncken 2003). Outcrops in the Western Puna and those in the east, such as the Cordillera Oriental, originally would have been more than 500 km apart. From the example of recent

passive continental margins, such a distance would reach water depths of several thousand metres, as off the Brazilian coastline. Thus, in the passive continental margin model of Jezek (1990), it could be expected that pelagic sediments would have been deposited in the western basin area. They should differ significantly in composition and grain size from those deposited at the eastern margin of the basin, which so far has not been shown to be the case. Recent isotope studies by Lucassen et al. (2000; 2002) suggest that the Arequipa terrane is Gondwana-related and, thus, may have functioned as the western basin boundary, as presumed by Keppie & Bahlburg (1999) and Acenolaza & Acenolaza (2002). Major element geochemistry Geochemistry of sedimentary rocks is a valuable tool for provenance studies of matrixrich sandstones as long as the bulk composition is not affected strongly by diagenesis, metamorphism and/or other alteration processes (McLennan et al. 1993). Abundances and ratios of major elements thus need to be checked for mobility, especially following diagenesis or lowgrade metamorphism (e.g. McLennan et al. 1990; McLennan 2001). Von Eynatten et al. (2003) showed that major element geochemistry is less effective for provenance purposes than trace element geochemistry. However, major element geochemistry can give important information about the weathering profile of rocks (e.g. Boles & Franks 1979; Nesbitt & Young 1984; Nesbitt et al. 1996; Bahlburg 1998; Zimmermann & Bahlburg 2003). An insight into the degree of weathering can be obtained from the Chemical Index of Alteration (CIA = molar [A12O3/A12O3 + CaO* + Na2O +K2O] X 100 where CaO* is CaO in silicates only; Nesbitt & Young 1982). The resultant CIA value is a measure of the proportion of A12O3 versus the mobile oxides in the analysed samples, typically representing the alteration of feldspars and volcanic glass to clay minerals. Overall, CIA values for the Puncoviscana complex scatter between 56 and 77 independently of grain size (Table 4 and see below), but some are elevated (VOL, SIJ-P, SIJ-G, SEC), whereas others are moderate (TOR, MOL, CON). Calcite as a secondary phase is rare, but when present it lowers the CIA significantly (e.g. CEB, Table 2). Excluding disturbed samples and those with high silica concentrations (which mask the alteration, e.g. Nesbitt & Young 1982), 90% of the samples have a mean of 65 (s.d. = 4) for the CIA. Nesbitt et al. (1996) show that in situ alteration

THE PUNCOVISCANA COMPLEX

401

degree. On the other hand, CIA values in the Ordovician foreland basin of the Puna are generally higher, averaging 71 (Bahlburg 1998). In diagrams using major element concentration (Bhatia 1983; Roser & Korsch 1986; 1988), the Puncoviscana complex samples plot in different tectonic settings, even including samples from the same outcrop. This can be explained by the mobility of K and Na during the weathering of feldspar, as shown above. Provenance discrimination using major elements is, therefore, unsuccessful. Bahlburg (1998) and Zimmermann & Bahlburg (2003) reached the same conclusion for overlying Ordovician rocks of the same region, and excluded major element geochemistry for provenance analysis because of their mobility during weathering. Fig. 3. A-CN-K diagram combined with CIA (after Fedo et al 1995) showing the weathering trend of the Puncoviscana complex. Note the relatively homogeneous trend towards illite. The solid arrow is an idealized weathering trend for upper continental crust. Average shale after Taylor & McLennan (1985). kao, kaolinite; ill, illite; ksp, alkali-feldspar; plag, plagioclase; CIA after Nesbitt & Young (1982).

of feldspar could produce a quartz-rich composition, causing difficulties in the interpretation of the CIA. This could explain the signature of the SIJ-G and SIJ-P samples (CIA 76-77 and 59-70, respectively). Also, a few samples (LA PI, CHO 2-4 and 2-5, SIJ-P 4-7, all samples from SIJ-G, and NP NP272) exhibit low concentrations of K2O and CaO, which may reflect albitization (e.g. Milliken 1988). Weathering trends are examined in the triangular diagram of Figure 3. The Puncoviscana complex data are scattered slightly, but deviate clearly from the ideal weathering trend for upper continental crust towards illite composition. This could be a result of metasomatic increase of potassium during diagenesis, caused either by the conversion of aluminous clay minerals to illite or by transformation of plagioclase to K-feldspar (Fedo et al 1995). Different angles of deviation from the ideal trend parallel to the A-CN line may indicate mixing of different sources variously affected by weathering, or secondary gain or loss, especially of Na and K (and potentially Ca) in the silicate fraction, e.g. during albitization (McLennan et al. 1993). Thus, the major element data presented show that the Puncoviscana rocks were affected by weathering to a significant

Trace element geochemistry The high field strength trace elements (HFSE) Th, Sc, Zr and rare earth elements (REE) - are useful for provenance analysis as they are insoluble and normally immobile under surface conditions. Moreover, due to their typical behaviour during fractional crystallization, weathering and recycling, they preserve characteristics of the source rocks in the sedimentary record (Taylor & McLennan 1985; Bhatia & Crook 1986; McLennan 1989; McLennan et al. 1993; Roser et aL 1996). The overall composition of the Puncoviscana complex samples, using immobile element ratios like Zr/Ti and Nb/Y, points to a uniform rhyodacitic source (after Winchester & Floyd 1977; Table 2). General characteristics of the deposits of the Puncoviscana Basin are those of upper continental crust (UCC, McLennan et al 1990). Compared to UCC they are enriched in Cs and depleted in Sr and Ni. LILE (Large-ion lithophile elements) are scattered, due to their more sensitive reaction to alteration and weathering, but are generally slightly enriched, as expected in upper crustal rocks (Floyd & Leveridge 1987; McLennan et al 1990). The HFSE mostly show typical UCC patterns: depletion in Sc, a compatible element, and enrichment in LREE (light rare earth elements) and Th. Ta and Nb also occur in typical UCC concentrations. Silica-enriched samples (e.g. SIJ-G, VOL) show dilution effects, lowering their trace element concentrations. The behaviour of Zr and Hf in the Puncoviscana complex displays a broad scatter, but only in few samples are both elements enriched relative to UCC (Table 2). Mafic source terranes would have high ferromagnesian mineral abundances, resulting in high Cr/V ratios (> 8) and low Y/Ni (< 0.5). Cr/V

402

U. ZIMMERMANN

Fig. 4. Th/Sc vs Zr/Sc diagram after McLennan et al (1990). Data for the Puncoviscana complex show some spread of Zr/Sc ratios even within each outcrop area, but the individual values are mainly lower than 30. Th/Sc ratios are very constant and point to a UCC composition. The symbols indicate the mean of each sampled outcrop.

ratios indicate enrichment of Cr over other ferromagnesian trace elements, where Y/Ni tests the abundance of ferromagnesian minerals (McLennan et al. 1993). The Puncoviscana complex shows low Cr/V ratios and elevated Y/Ni abundances (Tables 2, 4). Sc is a good tracer of mafic source components, particularly when compared with Th, which is incompatible and, thus, strongly enriched in felsic, upper crustal rocks. Both elements are generally immobile under surface conditions, therefore preserving the characteristics of their source. Thus, the Th/Sc ratio is considered a robust provenance indicator (Taylor & McLennan 1985; McLennan et al. 1990); its value in average upper continental crust is 0.79 (McLennan 2001). All samples of the Puncoviscana Basin show an upper crustal composition with Th/Sc ratios between 0.75 and 2. The Zr/Sc ratio is used commonly as a measure of the degree of sediment recycling leading to the enrichment of the ultra stable mineral zircon in the deposits (e.g. McLennan et al. 1990; 1993). Zr/Sc ratios in the Puncoviscana samples range between 7 and 70 (Fig. 4), but 87% of the samples have values below 30. In some outcrops, silica-rich samples show slightly elevated Zr/Sc ratios (e.g. VOL VR37, CHO 02-4 and 02-5, LA PI; Table 4). However, the medium- to coarse-sand SIJ-G psammites, with silica concentrations over 80%, have values of only 19.4-33.8. Zr/Sc ratios appear to be relatively unaffected by grain-size distribution, showing no systematic trend (Fig. 5).

Rare earth elements (REE)

Fig. 5. Zr/Sc ratios versus grain size. The symbols mark the mean of each grain-size class, the thin line represents the standard deviation, and the crosses mark samples with higher Zr/Sc ratios. The light grey arrow shows the trend of elevated Zr/Sc resulting from concentration of zircon by reworking, after McLennan et al. (1990). The grain sizes are marked in phi; n = number of samples per grain-size category. For comparison, two deposits of reworked sedimentary rocks are plotted: 1, Rio Taique (Tremadoc, data from Estelau & Zimmermann 2002); and 2, Meson Group (Middle Cambrian, unpublished data and from Bock et al 2000). Note that Zr/Sc is not correlated positively with grain size and reflects mainly UCC composition. All samples are poorly sorted and comprise a range of grain sizes; the values used (Table 4) are estimated to be the most representative for each sample.

The REE are well-established provenance indicators (McLennan 1989; 2001; McLennan et al. 1990; 1993), although some mobility may occur during weathering and diagenesis (Milodowski & Zalasiewicz 1991; Zhao et al. 1992; Bock et al. 1994; McDaniel etal 1994; Utzmann etal 2002). The REE abundances and patterns of nearly all samples from the Puncoviscana complex are homogeneous, independent of grain size (Table 4 and Figs 6a-e), and comparable to those of PAAS (post-Archaean Australian average shale, Nance & Taylor 1976; Fig. 6a). In general, the rocks exhibit slightly elevated REE concentrations (both LREE and HREE). Expressed as 2REE (mean = 186; Table 2), this is higher than UCC (148; McLennan 1989). The only excep tions are the rocks from Rio Taique (2REE = 106, s.d. = 12). Nearly all samples show a pronounced negative Eu-anomaly, a relatively steep pattern with LaN-YbN mostly > 4.4 and flat HREE (heavy rare earth element) patterns typical for upper crustal rocks (Fig. 6; McLennan

THE PUNCOVISCANA COMPLEX

et al 1990). Some samples are diluted in REE due to their high silica content (e.g. sample COR 02-30, with SiO2 - 77.73 and 2REE = 136; for comparison sample 02-28, from the same outcrop, has SiO2 = 59.86 and 2REE = 207 (Table 4, Fig. 6c). Divergence from the overall REE characteristics are observable in the outcrops of LA (Fig. 6e) and RT (Fig. 6b; discussed later), as well as in one sample from CEB (Fig. 6d). Eu anomalies are calculated as Eu/Eu* - Eu N /((Sm N )-(Gd N )) 1/2 , where the subscript N denotes the chondrite-normalized value and Eu* represents the Eu value expected for a smooth chondrite-normalized REE pattern. Eu/Eu* values are good indicators of source-rock composition, as high values reflect plagioclase enrichment in the source rocks (McLennan 1989). The Eu/Eu* values are relatively constant throughout the Puncoviscana Basin, with average values of individual outcrops between 0.5 and 0.78 (Tables 2 & 4), and 80% of the samples display values below 0.6; only three samples exhibit values higher than 0.7 (Table 4). Ce/Ce* anomalies (calculated as CeN/((LaN)2/3- (NdN)1/3)) are relatively constant, with values, mostly around 1 or negative (Tables 2, 4). On the other hand, RT, RAN, SUN and CON in the northern, western and central area of the basin (Fig. 1) comprise samples with strong positive Ce anomalies of 1.2-1.8 (Table 4; Figs 6b, d). High Ce/Ce* ratios coincide normally with moderate depletion in LREE, indicated by lower LaN/SmN ratios and low ZREE (McDaniel et al 1994); this signature is seen in the RT samples 02-94 and 02-95, which have Ce/Ce* of 1.21-1.64, LaN/SmN ratios of 2.07-2.53, and very low 2REE of 69-87.

Implications of the trace and rare earth element geochemistry Comparison of the studied units with the average composition of UCC (McLennan 2001) shows that the detrital material is enriched in incompatible and partly depleted in compatible elements. The Puncoviscana complex shows the influence of different source rocks, but predominantly of UCC composition, which is recorded in the entire basin. Compared to the nearly uniform PAAS-like REE patterns with high LaN/YbN values, low values are recorded from samples of RT and SUN, which will be discussed below. The LA samples show a great deal of variability, due to widely ranging LREE concentrations, whereas the HREE are fairly constant (Fig. 6e). This outcrop is deformed strongly and penetrated by

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mafic dykes, which could have caused mobility of LREE as a result of intensive fluid flow. One sample of CEB (02-264) exhibits a SiO2 content of 62.5%, with 2REE of 416 (Table 4), but displays the same shaped pattern as other samples from the same outcrop (Fig. 6d). This could be related to an addition of metamorphic heavy mineral phases, as the rocks in the sampling area are partly metamorphosed by contact with a Palaeozoic pluton (Hockenreiner 1998). The low EuN/Eu* ratios are typical for older continental crust rocks (McLennan et al 1993; Tables 2, 4). Only some samples with a high plagioclase content, have slightly elevated Eu/Eu* values (e.g. 0.87 for LA P8). Volcanic arc terranes show significantly higher Eu/Eu* values (0.9-1.2 or higher), which reflect the enrichment of plagioclase in those rocks. The strongly pronounced positive anomalies in the northern central part of the basin can be explained by intense surface weathering caused in a strong oxidizing environment, where most LREE are removed preferentially (McDaniel et al 1994). Ce, when oxidized to Ce4+, is removed less readily from the system as it is then incorporated more readily into insoluble hydroxides and oxides (e.g. Banfield & Eggleton 1989). This could imply a very shallow-water or a subaerial environment. Several authors have suggested a deepening of the basin towards the west away from the proposed passive margin and the detrital source, based on palaeocurrents and sedimentology (Jezek et al 1985; Jezek 1990), geochemistry (Willner & Miller 1986; Willner et al 1990) and trace fossil ichnofacies distribution (Acenolaza 1982; Durand & Acenolaza 1990; Aceolaza & Acenolaza 2002). This should imply mainly deep-marine sediments and anoxic conditions, which are not observed. The positive Ce/Ce* in samples of outcrops in the north, west and central part of the proposed Puncoviscana Basin thus point to a heterogeneous basin morphology. Th-Sc and Zr-Sc ratios are slightly variable and point to UCC composition; no sample shows ratios consistent with a volcanic arc source terrane (Tables 3 & 4). Furthermore, the Zr/Sc ratios are relatively low in comparison with typical passive margin deposits. In Figure 5 the data for the Puncoviscana complex are compared with the UCC, typical values from different modern rifted or passive margin successions (McLennan et al 1990), and overlying reworked quartz and silica-rich rocks of the Cambrian Meson Group (mean Zr/Sc 130; from Bock et al (2000, and unpublished data), and the Tolar Chico Formation (Lower

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THE PUNCOVISCANA COMPLEX

Tremadoc) in northwestern Argentina (mean Zr/Sc 80; Zimmermann & Bahlburg 2003), both interpreted as passive margin deposits. Reworking processes on basin margins resulting in well-sorted quartz arenites are strongly dependent on the climate and morphology (e.g. Suttner et al. 1981), but can sometimes be identified in clastic sedimentary rocks. The associated stronger weathering profile should be recognizable in strongly elevated Th/U, Rb/Sr and lower K/Rb values than UCC (McLennan et al. 1993). As shown, CIA varies only from moderate to high in the entire basin, without being dependent on grain size or basin position. The Th/U ratio correlates almost perfectly with the CIA (Tables 2,4). Exceptions are the medium- to coarse-sand psammites from Sijan and two samples from LA, which show disturbances probably related to their high silica contents (Nesbitt & Young 1982). Rb/Sr ratios are evolved (>0.5; cf. UCC, McLennan et al. 1993; Tables 2, 4), since Rb+ is retained more readily on exchange sites of clays than the smaller Sr2+ ion; K/Rb ratios are relatively low, lower than 250 ppm (UCC after McLennan 2001). Alteration took place but is moderate and constant throughout the entire basin. Polycyclic reworking will concentrate heavy minerals like zircon, monazite, rutile and allanite, even small amounts of which can control REE patterns (McLennan 1989). The enrichment of zircon in a sedimentary rock, for example, will cause higher REE concentrations and especially enriched HREE, leading to lower GdN/YbN ratios. From 119 samples only a few samples show low GdN/YbN values (< 1.2; Table 4) associated with higher absolute Zr concentrations. However, McLennan (1989) argues that the addition of 0.1 % of zircon in sandstones would raise the Zr concentration to 600 ppm and lower the GdN/YbN by c. 10%; none of the Puncoviscana samples exhibits such high Zr concentrations and the higher values are not related to grain size (Fig. 7). Thus, the low GdN/YbN ratios cannot be explained by zircon addition alone. Zircon addition should also be reflected in higher Hf concentrations relative to Th and Co (Basu et al. 1990). This can be observed only in three samples (TOR 72, CHO 02-4 and 02-5). The other 116 samples scatter

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Fig. 7. Element distribution vs grain sizes (in phi). The symbols mark the mean of each element per grain-size class, the thin lines standard deviations (see Tables 2, 4). Note that there is a slight trend from clayey to very-fine sand samples correlating with an increase in element concentration. However, the values are only slightly enriched compared to UCC composition (after McLennan 2001). Interestingly, medium- to coarse-grained rocks show no pronounced reworking, except for SiO2 concentrations. All samples are poorly sorted and comprise a range of grain sizes; the values used (Table 4) are estimated to be the most representative for each sample.

around the UCC composition (Taylor & McLennan 1985; not shown here, but see Table 4). Similarly, monazite addition of about 0.005% in coarse-grained material and 0.02% in shales will lift the GdN/YbN ratios to >2.0. This is only the case in three samples of the

Fig. 6. REE element patterns normalized to chondrites (Taylor & McLennan 1985) for each outcrop of the Puncoviscana complex, (a) Plot of the averages of each outcrop. They show similar patterns to UCC and PAAS. RT (Rio Taique) exhibits depleted REE concentrations and CEB (La Cebila) enriched, but they have the same shape as PAAS (post-Archaean Australian average shale; Taylor & McLennan 1985). MOL (Molinos) is depleted in LREE, and NP (Los Corrales) is strongly enriched in LREE only, (b) REE patterns for outcrops in the Puna, (c) REE patterns for Cordillera Oriental, (d) REE patterns for Ambato. (e) REE patterns for Famatina. The grey areas show the envelope of REE patterns for each outcrop.

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Fig. 8. Provenance discrimination diagram using ratios of Th, Sc and Zr/10 after Bhatia & Crook (1986) and Bahlburg (1998). A, oceanic island arc; B, continental island arc; C, active continental margin; D, rifted margin; TE, trailing edge; CA, continental arc. The shadowed area shows a revision of the field boundaries based on data from modern clastic rocks (Bahlburg 1998).

medium- to coarse-sand psammites from Sijan with slightly evolved LaN/YbN values (Table 4). Allanite addition will result in extremely high LaN/YbN ratios; addition of only 0.02% of this mineral would produce remarkable LaN/YbN values of > 15, which are not observed. Addition of mixtures of these three heavy minerals could mask the effect on REE concentrations, but would result in higher Zr, Th and Hf concentrations. In Figure 7, grain sizes in phi (= -Iog2 of the grain diameter, Wentworth 1922) are plotted versus trace and major elements. Silica contents rise with the coarseness of the material, but only the rocks with phi values between 1 and 0 show this trend clearly. As expected, a higher silica concentration dilutes the rocks in elements not concentrated in stable minerals (the example here is A12O3; Korsch et al 1993). On the other hand, elements that are concentrated in stable heavy minerals can be enriched during reworking, e.g. TiO2 (rutile), and Zr and Hf (both in zircon). A slight trend, or reworking, from clay to very fine sand is observable, but element abundances do not exceed those of UCC by more than 15%. Trace element ratios such as La/Th, La/Sc, Zr/Sc and Th/Sc have been used successfully to discriminate tectonic settings (Taylor & McLennan 1985; Bhatia & Crook 1986; McLennan et al 1990; 1993; Zimmermann &

Bahlburg 2003). However, such an approach has to be used with caution because it has been shown that specific tectonic settings do not necessarily produce sedimentary rocks with unique geochemical signatures (McLennan et al 1990; Bahlburg 1998). Th-Sc-Zr/10 characteristics of the Puncoviscana complex (Fig. 8) scatter from the continental arc to rifted margin fields defined by Bhatia & Crook (1986). However, the average values for each outcrop point to a provenance of a rifted margin and/or active continental margin rather than to a continental arc source. In a diagram of La/Th ratios against Hf (Fig. 9), most of the samples plot with uniform La/Th ratios, close to a UCC composition. Only samples from SIJ-G, SIJ-P and LA show lower Hf values, probably mainly caused by high SiO2 concentrations (Table 4). The generally upper crustal Hf concentrations (Table 2) are not consistent with a strong influence from old crustal material (Fig. 9). Certain trace element ratios and element concentration can be used to detect the compositional influence of a volcanic arc source terrane in the source (Floyd & Leveridge 1987; Hofmann 1988; 1997; McLennan et al 1990; 1993; Bock et al 2000; Zimmermann & Bahlburg 2003). As mentioned above, volcanic debris was deposited in the Puncoviscana complex, but it has to be determined if these volcanic and associated plutonic components were derived from a volcanic arc source. In Table 3, mean Th/Sc and Eu/Eu* ratios and Ta, Nb and Ti concentrations are normalized to typical UCC (Nc) and to values for a volcanic arc signature (NA) (after McLennan et al 1990; 1993). Th/Sc values (0.8-2.0; Tables 3 & 4) are atypical for continental volcanic arc terranes, as they are far too elevated. The Puncoviscana complex shows a trend opposite to that expected in volcanic arc settings (McLennan et al 1993). Arc terranes generally show no Euanomalies 1 (Eu/Ea* ~ 1; McLennan et al 1993), whereas in the Puncoviscana complex Eu/Eu* values are clearly <1 and typical of UCC. However, dissected arc terranes and depositional areas in certain distances from the continental arc can comprise a wider range of Eu/Ea values (0.64-0.94; McLennan et al 1990), which are rarely represented (Table 4). Strong negative Ta, Nb and Ti anomalies are often used for deciphering a continental arc source or component (e.g. Floyd & Leveridge 1987; Hofmann 1988; 1997). For all three elements, the samples show concentrations above UCC values. The samples from VOL, TOR and SIJP/G are relatively depleted in Ta, Nb and Ti, but this is caused by dilution by high SiO2 not

THE PUNCOVISCANA COMPLEX

Fig. 9. La/Th versus Hf diagram after Floyd & Leveridge (1987). Fields indicate the composition for sedimentary rocks related to different tectonic settings. Most samples point to an upper continental crust composition. The rocks from SIJ-P show low Hf concentrations and the metagreywackes from the same area point to an acidic arc source. Note the relatively uniform La/Th ratios.

because of a continental arc provenance. The sample CEB 02-264, with low silica concentration (62%) and low TiO2 abundance (0.3), seems to have been affected by a post-depositional event, such as the emplacement of a pluton (Hockenreiner 1998), as its whole geochemistry is apparently disturbed (Table 4; Fig. 6d). Overall, trace element compositions show that no continental arc component or mafic component contributed to the detrital material of the Puncoviscana complex. Nevertheless, the provenance diagrams using Th-ZrllO-Sc ratios should be considered cautiously. Most of the samples show affinities to the modified fields for rifted margin rocks presented by Bahlburg (1998) based on modern turbidites (Fig. 8). However, passive margins and foreland-basin deposits cannot be distinguished using provenance diagrams such as Figures 8 and 9.

Discussion and interpretation The Puncoviscana complex is a key element in understanding the evolution of the western margin of Gondwana from the Late Neoproterozoic until the Early Cambrian. However,

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interpretation of this large basin is complicated by isolated outcrops, polyphase deformation, metamorphism and massive intrusive activity during Early-Mid-Cambrian and Early Palaeozoic times. Former provenance interpretations of the whole Puncoviscana complex were based mainly on petrography and major element geochemistry. Von Eynatten et al (2003) have shown that the most reliable tool for such an approach is trace element geochemistry, following the approach of McLennan et al (1993), who proposed similar techniques to decipher the origin and composition of source terranes for sedimentary rocks. These techniques cannot, at this early stage, overcome the problem of incomplete lithostratigraphy, but can offer substantial data for the understanding of the basin. Previously, petrography and major element geochemistry led to the interpretation that the Puncoviscana complex represents polycyclic passive continental margin deposits, with the main sediment supply from the east into a c. N-S elongated basin (e.g. Acenolaza et al 1988; Jezek 1990; Willner et al 1990). Deepening to the west was suggested on the basis of sedimentological data, including trace fossil assemblages (Acenolaza & Acenolaza 2002). However, recent approaches to the problem of Upper Vendian life forms have reinterpreted models of ichnofacies substantially, pointing to a more complicated model, but affirming a depositional age of uppermost Vendian to Lower Cambrian (Buatois & Mangano 2003). The new petrographic analysis presented here shows that there are no significant petrographic changes from west to east in the Puncoviscana Basin. Grain sizes, matrix compositions and proportions, as well as framework mineral abundances, crystal forms and the general poorly sorted character of the rocks are relatively constant throughout the entire basin. A detailed study shows that mobility (especially of alkali elements) makes provenance interpretations based on major elements unsuccessful. The most reliable interpretations of tectonic setting for clastic sedimentary rocks are based on trace element and/or isotope geochemistry (e.g. McLennan et al 1990; von Eynatten et al 2003). Combining petrography with new geochemical data, the hypothesis of a passive continental margin for the Puncoviscana complex is seen as unlikely. A foreland-basin setting explains the results more successfully and fits with proposed tectonic models. There is no indication of a dominant process of recycling, recorded in the geochemistry of several passive margin deposits, modern as well as

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Fig. 10. (a) Depositional model for the Puncoviscana complex as a peripheral foreland basin. Material from the east was probably mixed with exhumed basement detritus of UCC composition. These 'arches' could represent heterogeneities in the basin morphology and explain the variations in Ce-anomalies. A further source to the west is proposed (after Ricci Lucchi 1984). (b) The proposed tectonic setting of the Puncoviscana complex is a peripheral foreland basin on the hypothetical composite Pampia-Arequipa block. The deposits in the peripheral foreland basin could have been folded and, in part, cannibalistically recycled within the same depositional area. The petrographical and geochemical dataset presented here supports the model of Keppie & Bahlburg (1999) and Kraemer et al (1995). Sketch modified after Keppie & Bahlburg (1999).

ancient (e.g. McLennan et al. 1990, 1993; Esteban & Zimmermann 2002; Zimmermann & Bahlburg 2003). REE and trace element ratios do not point to a significant accumulation of heavy minerals. Furthermore, strong reworking or recycling would produce clear evidence of weathering, which is not the case. Minimal recycling, together with the relatively moderate and constant weathering profile, would favour the hypothesis that alteration processes affected the sediments mainly during or after deposition. Volcanic debris was found in the rocks in the form of sanidine, and volcanic pebbles and lithoclasts were observed in conglomeratic deposits of Quebrada de Suncho (Durand 1990; van Staden & Zimmermann 2004). Most volcanic components probably were decomposed to form pseudomatrix. Geochemical tests

for continental volcanic arc influence, using element ratios such as Th/Sc and Eu/Eu* and concentrations of Ta, Nb and Ti (e.g. Floyd & Leveridge 1987; Hofmann 1988; 1997; McLennan et al 1990; 1993), failed to decipher such a component in the rocks presented here. Only isotope analysis can finally exclude or trace a volcanic arc component (McLennan et al 1993), but existing preliminary data exclude such a source (Bock et al 2000). Furthermore, the geochemical data show that certain areas were exposed to an environment with a high oxidation potential where strong positive Ce/Ce* anomalies occur. In all cases, except samples from Quebrada Randolfo (Puna) and Conception (Ambato), this signature correlates with depletion of LREE, which can be removed more easily under highly oxidizing conditions (Banfield & Eggleton 1989; McDaniel et al

THE PUNCOVISCANA COMPLEX

1994). This suggests a heterogeneous basin morphology unlikely for a classic passive continental margin, as oxidizing conditions are recorded in outcrops in the central and western parts of the basin but are found near the eastern boundary as well. The Puncoviscana foreland basin could have been related to the Pampean Orogeny (Rapela et al 1998): it was situated on the Pampia block, at that time probably one entity with the Arequipa block (Figs, 2a, 10). The dominant source for the basin came from the east according to palaeocurrents (Jezek 1990) and was composed mainly of sedimentary rocks, exhumed metamorphic material and subordinate volcanic debris, funnelled through the foredeep and propagating fold-thrust belt into the peripheral foreland basin (e.g. Valloni 1984; Schwab 1986; Dutta & Wheat 1990; Fig. lOa). Sources to the east of the proposed Puncoviscana Basin (Fig. 2a), such as the eastern Sierras Pampeanas, including the Sierras de Cordoba and the Sierra de San Luis, could have provided the necessary detrital material to supply the basin infill. Felsic volcanic rocks are reported from the La Lidia Formation in the Sierra Norte de Cordoba and dated by conventional U-Pb on zircons at 584 ± 22 Ma (Llambias et al 2003) and might be a candidate for the volcanic input. However, volcanic debris is not recorded generally in ancient peripheral foreland basins (Lawton 1986; Wuellner et al 1986). It is sometimes seen in the early phases of continent-continent collision, e.g. the Ouachita orogenic belt (Houseknecht 1986). However, detrital supply into foreland basins can change during their evolution, depending on the regional tectonic processes (Zuffa et al 1980; Cibin et al 2001). Basin morphology in peripheral foreland basins is controlled by on-going tectonism, which can deform and break the basin along thrusts and uplift material (e.g. Homewood et al. 1986; Houseknecht 1986). During the propagation of fold-thrust belts in foreland basins, small local arches can be developed (Tankard 1986). These can exhume deeper crustal levels or previously deposited sediments for cannibalistic supply (Ricci Lucchi 1984) without long distance material transport. This type of process can explain best the petrographical and geochemical characteristics of the Puncoviscana complex (Fig. lOa). Bock et al (2000) and Zimmermann & Bahlburg (2003) demonstrated, using geochemistry and Nd-Pb isotope systematics, that the Neoproterozoic to Permian supracrustal rocks in northwest Argentina are composed mainly of older regional recycled crustal material of similar composition.

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A peripheral foreland-basin model makes a second main detrital source in the west probable (Kraemer et al 1995; Figs lOa, b). Lucassen et al (2000; 2002) used Pb isotopes to show that the basement material of the Arequipa block, a probable western source, and the Pampia block are difficult to distinguish. The relatively homogeneous distribution of geochemical and petrographic characteristics of the Puncoviscana complex makes such interpretation possible. After the cessation of sedimentation and deformation of the Puncoviscana complex, an intra-cratonic basin containing strongly reworked quartz-rich arenites of the Meson Group transgressed from the north (e.g. Kumpa & Sanchez 1988). During the Late Cambrian an east dipping subduction zone was initiated at the western border of the Pampia terrane (e.g. Rapela et al 1998), probably caused by a flipping of the subduction zone from east further to the west with the same dip direction. During the Ordovician the Puncoviscana complex was partly an erosional area. The rocks were recycled into quartz-rich arenites of the Tremadocian Tolar Chico and Rio Taique Formation, as both carry debris of the Puncoviscana complex (Esteban & Zimmermann 2002; Zimmermann & Bahlburg 2003), and into the subsequently developed Arenigian and Llanvirnian retro-arc and foreland basins (Bahlburg 1990,1998; Zimmermann et al 2002; Zimmermann & Bahlburg 2003).

Conclusions 1.

2.

Petrographic analyses show that the rocks of the Puncoviscana complex are sorted poorly, the principal components mainly sub-angular and that they exhibit a high amount of matrix (8-35%), very probably produced during the dissolution of labile lithoclasts. Quantitative petrographic studies point to a composition comparable to a recycled orogenic succession dominated by quartz and metamorphic, as well as sedimentary, lithoclasts. Sparse volcanic input can be recorded in the form of sanidine and probable volcanic lithoclasts, which contributed through decomposition to the high matrix content. The composition of the rocks (Qt 60-80 F 15-35 L 5-20; P/F 0.2-0.4; LV/L 0) is more compatible with foreland-basin deposits than a passive continental margin succession. Major element geochemistry shows a moderate to strong alteration (CIA between 56 and 77), independent of grain size. Provenance determination using

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3.

4.

5.

6.

7.

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major element chemistry failed, because of the mobility of the major elements like K, Na and Ca. The effect of weathering is also represented in elevated Rb/Sr and Th/U ratios. More than 90% of the samples show typical UCC composition using trace element geochemistry with robust element ratios such as Ti/Zr and Nb/Y (showing a rhyodacitic composition), LaN/YbN (> 4.4), Th/Sc (> 0.8), Zr/Sc (> 10) and REE patterns (mainly similar to PA AS). More than 80% of the samples show EuN/Eu* values between 0.45 and 0.6. Polycyclic reworking can be excluded on the basis of the petrographic results, as well as trace element geochemistry using REE, Zr/Sc, GdN/YbN, LaN/YbN, Zr, Hf and Ti. Heavy mineral addition of zircon, monazite, allanite or Ti-oxide phases that would point to recycling is not recorded. In the west and in the centre of the basin, high positive Ce/Ce* anomalies between 1.2 and 1.7 indicate a strongly oxidizing near-surface environment. The values coincide with a loss of LREE and consequently low ZREE. This excludes a simple, westerly deepening, basin morphology. Trace element ratios and concentrations of Th/Sc, Eu/Eu*, Nb, Ta and Ti exclude a volcanic arc or mafic source for the detritus. The small amount of observed volcanic debris could have come from a magmatic source other than a volcanic arc. The Puncoviscana Basin developed on the subducting Pampia block in a peripheral foreland-basin position. The Arequipa block most probably represented the western border of the basin. Ongoing tectonic activity could have fragmented the peripheral foreland basin and produced palaeohighs. The exhumed metamorphic and sedimentary successions were of UCC composition, cannibalistically reworked into the same basin, but transported only a short distance. Volcanic debris was probably funnelled from the collisional belt in the east into the foreland basin.

Finally, the data presented here constitute the first complete sampling of currently-known Puncoviscana complex deposits in northwestern Argentina. The lack of a complete lithostratigraphy is the most problematical aspect for understanding the palaeogeographical and tectonic evolution of Western Gondwana during the Neoproterozoic and Early Cambrian. However, these data mark a starting point from

which to compare other outcrop areas of the Puncoviscana complex, equivalent regions and probable source terranes by means of petrographical and geochemical data. A further understanding of the problem will be possible if age determinations of detrital minerals are obtainable on a large scale. This study benefited from the logistic help of Rudolfo Lucero (Institute Superior San Martin at San Fernando del Valle de Catamarca), financial help from Mario Contreras (Universidad Nacional de Catamarca) and funding by the University of Johannesburg and the Faculty of Science (UJ) through the SASOL fund (UJ). The idea and motivation for this study grew during stimulating discussions with Fernando Hongn and Heinrich Bahlburg. The author would like to thank his wife for company and help in the field and during the preparation of the sample material. Thanks go to R. J. Pankhurst who corrected the English. The contribution was improved by the reviews of H. von Eynatten and an anonymous reviewer. This is a contribution to IGCP Projects 436 (Pacific Gondwana Margin) and 478 (Neoproterozoic-Early Palaeozoic Events in SW Gondwana).

References ACENOLAZA, F.G. & TOSELLI, A.J. 1981. Geologia del Noroeste Argentina. Publication Especial de la Facultad de Ciencias Naturales, Universidad Nacional de Tucuman, 1287, 1-212. ACENOLAZA, F.G & ACENOLAZA, G.F. 2002. The Puncoviscana Formation (Northwest Argentina), an approach to the evolution of a basin in the Gondwanan border during the Neoproterozoic/ Early Cambrian. // International Colloquium Vendian-Cambrian of W-Gondwana, Montevideo, Abstract Volume, 7-9. ACENOLAZA, F.G., FERNANDEZ, R.I. & MANCA, N. 1982. Caracteres bioestratigraficos y paleoambientales del Grupo Meson (Cambrico MedioSuperior), centro-oeste de America del Sur. Estudios Geologicos, 38, 385-392. ACENOLAZA, F.G, MILLER, H. & TOSELLI, A.J. (eds) 1983. La geologia de la Sierra de Ancasti. Minister. Forschungshefte zur Geologic und Palaontologie, 59, 1-372. ACENOLAZA, F.G, MILLER, H. & TOSELLI, A.J. 1988. The Puncoviscana Formation (Late PrecambrianEarly Cambrian). Sedimentology, tectonometamorphic history and age of the oldest rocks of NW Argentina. In: BAHLBURG, H., BREITKREUZ, C. & GIESE, P. (eds) The Southern Central Andes: Contributions to Structure and Evolution of an Active Continental Margin. Lecturer Notes on Earth Sciences, 17, Springer Verlag, Berlin, 25-38. ACENOLAZA, F.G, MILLER, H. & TOSELLI, A.J. (eds) 1990. El Ciclo Pampeanas en el Noroeste Argentino. Universidad Nacional de Tucuman, Serie Correlation Geologica, 4,1-227. ACENOLAZA, F.G, MILLER, H. & TOSELLI, A.J. (eds)

THE PUNCOVISCANA COMPLEX 1996. Geologia del Sistema de Famatina. Miinchener Geologische Hefte, 19,1-410. ACENOLAZA, EG., MILLER, H. & TOSELLI, A.J. 2000. The Pampean and Famatinian cycles - superimposed erogenic events in West Gondwana. In: MILLER, H. & HERVE, F. (eds) Geoscientific Cooperation with Latin America. Zeitschrift fiir Angewandte Geologic, Sonderheft, 1, 337-344. ACENOLAZA, G. 2004. Precambrian-Cambrian ichnofossils, an enigmatic 'annelid tube' and microbial activity in the Puncoviscana Formation (La Higuera; Tucuman Province, NW Argentina). Geobios, 37, 127-133. ASTINI, R.A., BENEDETTO, J.L. & VACCARI, N.E. 1995. The early Paleozoic evolution of the Argentine Precordillera as a Laurentian rifted, drifted, and collided terrane: A geodynamic model. Geological Society of America Bulletin, 107, 253-273. ASTINI, R. A., RAMOS, V. A., BENEDETTO, J.L., VACCARI, N.E. & CANAS, EL. 1996. La Precordillera: un terreno exotico a Gondwana. XIII Congreso Geologico Argentina y HI Congreso de Exploration de Hidrocarburos, Buenos Aires, Actas, 5, 293-324. BACHMANN, G, GRAUERT, B., KRAMM, U, LORK, A. & MILLER, H. 1987. El magmatismo del Cambrico Medio/Cambrico superior en el basamento del noroeste Argentino: Investigaciones isotopicas y geocronologicas sobre los granitoides de los complejos intrusivos de Santa Rosa de Tastil y Canani. 10th Congreso Geologico Argentino, Tucuman, Actas, 4, 125-127. BAHLBURG, H. 1990. The Ordovician basin in the Puna of NW Argentina and N-Chile: geodynamic evolution from back-arc to foreland basin. Geotektonische Forschungen, 75, 1-107. BAHLBURG, H. 1998. The geochemistry and provenance of Ordovician turbidites in the Argentine Puna. In: PANKHURST, R.J. & RAPELA, C.W. (eds) The Proto-Andeau Margin of Gondwana. Geological Society, London, Special Publications, 142, 127-142. BAHLBURG, H. & HERVE, F. 1997. Geodynamic evolution and tectonostratigraphic terranes of northern Argentina and northern Chile. Geological Society of America Bulletin, 109, 869-884. BAHLBURG, H. & FLOYD, PA. 1999. Advanced techniques in provenance analysis of sedimentary rocks. Sedimentary Geology, 124, 1-220. BALDIS, B. & OMARINI, R. 1984. El Grupo Lerma (Precambrico-Cambrico) en la comarca central Saltena y su position en el borde Pacifico Americano. 9th Congreso Geologico Argentino, Buenos Aires, 1, 64-78. BANFIELD, J.F. & EGGLETON, R.A. 1989. Apatite replacement and rare earth element mobilization, fractionalisation, and fixing during weathering. Clays and Clay Minerals, 37, 113-127. BASU, A.R., SHARMA, M. & DECELLES, PG. 1990. Nd, Sr-isotopic provenance and trace element geochemistry of Amazonian foreland basin fluvial sands, Bolivia and Peru: implications for

411

ensialic Andean orogeny. Earth and Planetary Science Letters, 100, 1-17. BHATIA, M.R. 1983. Plate tectonics and geochemical composition of sandstones. Journal of Geology, 91, 611-627. BHATIA, M.R. & CROOK, K.A.W. 1986. Trace elements characteristics of graywackes and tectonic setting discrimination of sedimentary basins. Contributions to Mineralogy and Petrolology, 92, 181-193. BOCK, B., MCLENNAN, S.M. & HANSON, GN. 1994. Rare earth element redistribution and its effects on the neodymium isotope system in the Austin Glen Member of the Normanskill Formation, New York, USA. Geochimica et Cosmochimica Acta, 58, 5245-5253. BOCK, B., MCLENNAN, S.M. & HANSON, G.N. 1996. The Taconic orogeny in southern New England: Ndisotope evidence against addition of juvenile components. Canadian Journal of Earth Sciences, 33,1612-1627. BOCK, B., BAHLBURG, H., WORNER, G. & ZIMMERMANN, U. 2000. Tracing Crustal Evolution in the Southern Central Andes from Late Precambrian to Permian with Geochemical and Nd and Pb Isotope Data. The Journal of Geology, 108, 513-535. BOLES, J.R. & FRANKS, S.G. 1979. Clay diagenesis in Wilcox Sandstones of Southwest Texas: Implications of smectite diagenesis on sandstone cementation. Journal of Sedimentary Petrolology, 49, 55-70. BUATOIS, L.A. & MANGANO, M.G. 2003. La icnofauna de la Foramcion Puncoviscana en el noroeste argentine: la colonization de fondos oceanicos y reconstruction de paleoambientes y paleoecosistemas de la transition Precambrica-Cambrica. Ameghiniana, 40, 103-117. CAMINOS, R. 1979. Sierras Pampeanas noroccidentales. Salta, Tucuman, Catamarca, La Rioja y San Juan. Segundo Simposio de la Geologia Regional Argentina. Academia Nacional de Ciencias, Cordola, 225-291. CIBIN, U, SPADAFORA, E., ZUFFA, G.G. & CASTELLARIN, A. 2001. Continental collision history from arenites of episutural basins in the Northern Apennines, Italy. Geological Society of America Bulletin, 113, 4-19. CONDIE, K.C. & MARTELL, C. 1983. Early Proerozoic metasediments from north-central Colorado: metamorphism, provenance, and tectonic setting. Geological Society of America Bulletin, 94, 1215-1224. Cox, R. & LOWE, D.R. 1996. Quantification of the effects of secondary matrix on the analysis of sandstone composition, and a petrographicchemical technique for retrieving original framework grain modes of altered sandstones. Journal of Sedimentary Research, 66, 548-558. DECELLES, G. & HERTEL, B.I. 1989. Petrography and sedimentology of the Peruvian foreland basin in the Central Andes. Geological Society of America Bulletin, 101, 1552- 1562. DICKINSON, W.R. 1970. Interpreting detrital modes of

412

U. ZIMMERMANN

greywacke and arkose. Journal of Sedimentary Petrology, 40, 695-707. DICKINSON, W.R. 1988. Provenance and sediment dispersal in relation to paleotectonics and paleogeography of sedimentary basins. In: KLEINSPEHN, K.L. & PAOLA, C. (eds) New Perspectives in Basin Analysis. Springer Verlag, New York, 3-26. Do CAMPO, M. 1999. Metamorfismo del basamento en la Cordillera Oriental y borde oriental de la Puna. In: GONZALO HONORING, G., OMARINI, R. & VIRAMONTE, J. (eds) Relatorio de la Geologia del Noroeste de Argentina, Salta, 41-51. Do CAMPO, M. & RIBEIRO GUEVARA, S.R. 2002. Geoquimica de las sequencias clasticas de la Formacion Puncoviscana (Neoproterozoico, NO Argentina), proveniencia y marco tectonico. XV Congreso Geologico Argentino, El Calafate (CD). Do CAMPO, M., NIETO, E, OMARINI, R. & OSTERA, H. 1999. Neoproterozoic K-Ar ages for the metamorphism of the Puncoviscana Formation, northwestern Argentina. 20th International Simposio Sudamericano de Geologia Isotopica, Villa Carlos Paz, Abstract volume, 37-42. DURAND, ER. 1990. Los conglomerados del Ciclo Pampeanos en el Noroeste Argentino. In: ACENOLAZA, EG., MILLER, H. & TOSELLI, AJ. (eds) El Ciclo Pampeano en el Noroeste Argentino. Serie Correlacion Geologico, Universidad Nacional de Tucuman, 12, 61-69. DURAND, ER. & ACENOLAZA, EG 1990. Caracteres biofaunisticos, paleoecologicos y paleogeograficos de la Formacion Puncoviscana (Precambrico Superior-Cambrico Inferior) del Noroeste Argentino. In: ACENOLAZA, EG, MILLER, H. & TOSELLI, AJ. (eds) El Ciclo Pampeano en el Noroeste Argentino. Serie Correlacion Geologico, Universidad Nacional de Tucuman, 12, 71-112. DUTTA, P.K. & WHEAT, R.W. 1990. Climatic and Tectonic Control on Sandstone Composition in the Permo-Triassic Sydney Foreland Basin, Eastern Australia. In: JOHNSSON, M.J. & BASU, A. (eds) Processes Controlling the Composition of Clastic Sediments. Geological Society of America Special Publications, 284, 187-202. ESPEJO, I. & LOPEZ-GAMUNDI, O.R. 1994. Source versus depositional controls on sandstone composition in a foreland basin: the El Imperial Formation (Mid-Carboniferous-Lower Permian), San Rafael basin, Western Argentina. Journal of Sedimentary Research, 64, 8-16. ESTEBAN, S.B. & ZIMMERMANN, U. 2002. Tremadocian successions at the western border of Gondwana (northwestern Argentina): Initiation of the evolution of retro-arc basins? 16th International Sedimentological Congress, Rand Afrikaans University, Johannesburg, South Africa, 94-95. FEDO, C.M., NESBITT, H.W. & YOUNG, GM. 1995. Unravelling the effects of potassium metasomatism in sedimentary rocks and paleosoils, with implications for paleoweathering conditions and provenance. Geology, 23, 921-924. FLOYD, PA. & LEVERIDGE, B.E. 1987. Tectonic

environment of the Devonian Gramscatho basin, south Cornwall: Framework mode and geochemical evidence from turbidite sandstones. Journal of the Geological Society, London, 144, 531-542. FLOYD, PA., WINCHESTER, J.A. & PARK, R.G. 1989. Geochemistry and tectonic setting of Lewisian clastic metasediments from the Early Proterozoic Loch Maree Group of Gairloch, NW Scotland. Precambrian Research, 45, 203-214. FLOYD, P.A., SHAIL, R., LEVERIDGE, B.E. & FRANKE, W 1991. Geochemistry and provenance of Rhenohercynian synorogenic sandstones: implications for tectonic environment discrimination. In: MORTON, A.C., TODD, S.P. & HOUGHTON, P.D.W. (eds) Developments in Sedimentary Provenance Studies. Geological Society, London, Special Publications, 57, 173-188. HALLSWORTH, C.R. & KNOX, R.W. O'B. 1999. BGS Rock Classification Scheme, Volume 3, Classification of sediments and sedimentary rocks. British Geological Survey Research Reports, RR 99-03. HOCKENREINER, M. 1998. Die Formation La Cebila und ihr geologisches Umfeld (Sierra de Ambato, NW Argentinien). MSc thesis, Ludwig-Maximilians Universitat Miinchen, Germany. HOFMANN, A. 1988. Chemical differentiation of the Earth: The relationship between mantle, continental crust and oceanic crust. Earth Planetary Science Letters, 90, 297-314. HOFMANN, A. 1997. Mantle geochemistry: the message from oceanic volcanism. Nature, 385, 219-229. HOMEWOOD, P., ALLEN, PA. & WILLIAMS, GD. 1986. Dynamics of the Molasse Basin of western Switzerland. In: ALLEN, PA. & HOMEWOOD, P. (eds) Foreland Basins. International Association of Sedimentologists, Special Publications, 8, 199-217. HONGN, E, MON, R., CUEVAS, J. & TUBIA, M. 1996. Zones de cisaillement caledoniennes a hautes temperature dans la Quebrada Barranquilla (Puna Orientate, Argentine): donnees structurales et cinematiques. Compte Rendu Academic des Sciences, 323, Ha, 809-815. HOUSEKNECHT, D.W. 1986. Evolution from passive margin to foreland basin: the Atoka Foramtion of the Arkoma Basin, south-central U.S.A. In: ALLEN, PA. & HOMEWOOD, P. (eds) Foreland Basins. International Association of Sedimentologists, Special Publications, 8, 327-345. INGERSOLL, R.V., BULLARD, T.E, FORD, R.L., GRIMM,

J.P., PICKLE, J.D. & SARES, S.W 1984. The effect of grain size on detrital modes: a test of the Gazzi-Dickinson point-counting method. Journal of Sedimentary Petrology, 54, 103-116. JEZEK, P. 1990. Analisis sedimentologico de la Formacion Puncoviscana entre Tucuman y Salta. In: ACENOLAZA, EG, MILLER, H. & TOSELLI, AJ. (eds) El Ciclo Pampeano en el Noroeste Argentino. Serie Correlacion Geologico, Universidad Nacional de Tucuman, 12, 9-35. JEZEK, P. & MILLER, H. 1986. Deposition and facies Distribution of turbiditic sediments of the Puncoviscana Formation (Upper Precambrian-Lower

THE PUNCOVISCANA COMPLEX Cambrian) within the basement of the NW Argentine Andes. Zentralblatt fur Geologic und Paldontologie, Teil I, 9/10, 1235-1244. JEZEK, P., WILLNER, A.P., ACENOLAZA, F.G. & MILLER H. 1985. The Puncoviscana trough - a large basin of Late Precambrian to Early Cambrian age on the Pacific edge of the Brazilian shield. Geologische Rundschau, 74, 573-584. JORDAN, T.E. 1995. Retroarc foreland and related basins. In: BUSBY, CJ. & INGERSOLL, R.V. (eds) Tectonics of Sedimentary Basins. Blackwell Science, Oxford, 331-362. KEPPIE, J.D. & BAHLBURG, H. 1999. Puncoviscana Formation of northwestern and central Argentina: Passive margin or foreland basin deposits? In: RAMOS, V.A. & KEPPIE, J.D. (eds) Laurentia-Gondwana Connections before Pangaea. Geological Society of America Special Publications, 336, 139-144. KLEY, J. 1998. Variable foreland shortening along the southern Central Andes (15°-42° S): Correlation with crustal thickness, lithospheric structure and the geometry of the Nazca Plate. X Congreso Latinoamericano de Geologia y 6th Congreso Nacional de Geologia Economica, Buenos Aires, 2, 88-93. KLEY, J., MULLER, J., TAWACKOLI, S., JACOBSHAGEN, V. & MANUTSOGLU, E. 1997. Pre-Andean and Andean-Age Deformation in The Eastern Cordillera of Southern Bolivia. Journal of South American Earth Sciences, 1, 1-19. KORSCH, R.J., ROSER, B.P & KAMPRAD, J.L. 1993. Geochemical, petrographic and grain-size variations within single turbidite beds. Sedimentary Geology, 83, 15-35. KRAEMER, P.E., ESCAYOLA, M.P. & MARTINO, R.D. 1995. Hipotesis sobre la evolution tectonica neoproterozoica de las Sierras Pampeanas de Cordoba (30°40'-32°40'), Argentina. Revista de la Asociacion Geologica Argentina, 50, 47-59. KUMPA, M. & SANCHEZ, T.M. 1988. Geology and sedimentology of the Cambrian Grupo Meson (NW Argentina). In: BAHLBURG, H., BREITKREUZ, C. & GIESE, P. (eds) The southern Central Andes: Contributions to Structure and Evolution of an Active Continental Margin. Lecturer Notes on Earth Sciences, 17, Springer Verlag, Berlin, 39-53. LAWTON, T.F. 1986. Compositional trends within a clastic wedge adjacent to a fold-thrust belt: Indianola Group, central Utah, U.S.A. In: ALLEN, PA. & HOMEWOOD, P. (eds) Foreland Basins. International Association of Sedimentologists, Special Publication, 8, 411-423. LLAMBIAS, E.J., SATO, A.M., ORTIZ SUAREZ, A. & PROXY, C. 1998. The granitoids of the Sierra de San Luis. In: PANKHURST, R.J. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 325-341. LLAMBIAS, E.J., GREGORI, D., BASEI, M.A., VARELA, R. & PROZZI, C. 2003. Ignimbritas rioliticas neoproterozoicas en la Sierra Norte de Cordoba: /,evidencia de un arco magmatico temprano en el

413

ciclo Pampeano? Revista de la Asociacion Geologica Argentina, 58, 572-582. LOPEZ DE LUCHI, M.G., CERREDO, M.E., SIEGESMUND, S., STEENKEN, A. & WEMMER, K. 2003. Provenance and tectonic setting of the protoliths of the metamorphic complexes of Sierra de San Luis. Revista de la Asociacion Geologica Argentina, 58, 525-540. LORK, A., MILLER, H. & JEZEK, P. 1990. Sistematica U-Pb de circones detriticos de la Fm. Puncoviscana y su significado para la edad maxima de sedimentacion en la Sierra de Cachi (Prov. de Salta) Argentina. In: ACENOLAZA, F.G, MILLER, H. & TOSELLI, A.J. (eds) El Ciclo Pampeano en el Noroeste Argentino. Serie Correlation Geologico, Universidad Nacional de Tucuman, 12, 199-208. LUCASSEN, F, BECCHIO, R., WILKE, H.G., FRANZ, G, THIRLWALL, M.F., VIRAMONTE, J. & WEMMER, K. 2000. Proterozoic-Paleozoic development of the basement of the Central Andes (18-26°S) - a mobile belt of the South American craton. Journal of South American Earth Science, 13, 697-715. LUCASSEN, F, HARMON, R., FRANZ, G, ROMER, R.L., BECCHIO, R. & SIEBEL, W. 2002. Lead evolution of the Pre-Mesozoic crust in the Central Andes (18-27°): progressive homogenisation of Pb. Chemical Geology, 186, 183-197. MCDANIEL, D.K., HEMMING, S.R., MCLENNAN, S.M. & HANSON, G.N. 1994. Resetting of neodymium isotopes and redistribution of REEs during sedimentary processes: the Early Proterozoic Chelmsford Formation, Sudbury Basin, Ontario, Canada. Geochimica et Cosmochimica Ada, 58, 931-941. MCLENNAN, S.M. 1989. Rare earth elements in sedimentary rocks: Influence of provenance and sedimentary process. In: LIPIN, B.R. & McKAY, G.A. (eds) Geochemistry and Mineralogy of Rare Earth Elements. Mineralogical Society of America, Reviews in Mineralogy, 21, 169-200. MCLENNAN, S.M. 2001. Relationships between the trace element composition of sedimentary rocks and upper continental crust. Geochemistry, Geophysics, Geosystems, 2, 2000GC000109. MCLENNAN, S.M., TAYLOR, S.R., MCCULLOCH, M.T. & MAYNARD, IB. 1990. Geochemical and Nd-Sr isotopic composition of deep-sea turbidites: crustal evolution and plate tectonic associations. Geochimica et Cosmochimica Acta, 54, 2015-2050. MCLENNAN, S.M., HEMMING, S., MCDANIEL, D.K. & HANSON, G.N. 1993. Geochemical approaches to sedimentation, provenance and tectonics. In: JOHNSSON, MJ. & BASU, A. (eds) Processes Controlling the Composition of Clastic Sediments. Geological Society of America Special Publications, 284, 21-40. MILLIKEN, K.L. 1988. Loss of provenance information through subsurface diagenesis in Plio-Pleistocene sandstones, northern Gulf of Mexico. Journal of Sedimentary Petrology, 58, 992-1002. MILODOWSKI, A.E. & ZALASIEWICZ, J.A. 1991.

414

U. ZIMMERMANN

Redistribution of rare earth elements during diagenesis of turbidite/hemipelagite mudrock sequences of Llandovery age from Central Wales. In: MORTON, A.C., TODD, S.P. & HOUGHTON, P.D. W. (eds) Developments in Sedimentary Provenance Studies. Geological Society, London, Special Publications, 57, 101-124. MON, R. & HONGN, F. 1991. The structure of the Precambrian and Lower Paleozoic basement of the Central Andes between 22° and 32° Lat. Geologische Rundschau, 80, 745-758. MORTON, A.C. & HALLSWORTH, C.R. 1999. Processes controlling the composition of heavy mineral assemblages in sandstones. Sedimentary Geology, 124, 3-29. NANCE, W.B. & TAYLOR, S.R. 1976. Rare earth element patterns and crustal evolution - I. Australian post-Archean sedimentary rocks. Geochimica et Cosmochimica Acta, 40, 1539-1551. NESBITT, H.W. & MARKOVICS, G. 1997. Weathering of granodioritic crust, long-term storage of elements in weathering profiles, and petrogenesis of siliciclastic sediments. Geochimica et Cosmochimica Acta, 61, 1653-1670. NESBITT, H.W. & YOUNG, G.M. 1982. Early Proterozoic climates and plate motions inferred from major element chemistry of lutites. Nature, 299, 715-717. NESBITT, H.W & YOUNG, G.M. 1984. Prediction of some weathering trends of plutonic and volcanic rocks based on thermodynamic and kinetic considerations. Geochimica et Cosmochimica Acta, 48, 1523-1534. NESBITT, H.W, YOUNG, G.M., MCLENNAN, S.M. & KEAYS, R.R. 1996. Effects of chemical weathering and sorting on the petrogenesis of siliciclastic sediments, with implications for provenance studies. Journal of Geology, 104, 525-542. OMARINI, R.H. 1983. Caracterizacion litologia, diferenciacion y genesis de la Formacion Puncoviscana entre el Voile de Lerma y la Faja Eruptiva de la Puna. PhD thesis, Universidad National de Salta, Argentina. OMARINI, R.H., SUREDA, R.J., GOTZE, H.J., SEILACHER, A. & PFLUGER, F. 1999. Puncoviscana folded belt in northwestern Argentina. Testimony of Late Proterozoic Rodinia fragmentation and pre-Gondwana collisional episodes. International Journal of Earth Science, 88, 76-97. PANKHURST, R.J., & RAPELA, C.W. 1998. The protoAndean margin of Gondwana: an introduction. In: PANKHURST, R.J. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 1-10. PANKHURST, R.J., RAPELA, C.W, SAAVEDRA, J., BALDO, E., DAHLQUIST, J., PASCUA, I. & FANNING, C.M. 1998. The Famatinian magmatic arc in the central Sierras Pampeanas: an Early to Mid Ordovician continental arc on the Gondwana margin. In: PANKHURST, RJ. & RAPELA, C.W (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 181-217. PROZZI, C. 1990. Consideraciones acerca del Basa-

mento de San Luis. X7 Congreso Geologico Argentina, San Juan, Actas, 1, 452-455. RAMOS, V.A. 1986. Late Proterozoic-Early Paleozoic of South America - a collisional history. Episodes, 11, 168-173. RAMOS, V.A., VUJOVICH, G, MAHLBURG KAY, S. & MCDONOUGH, M. 1993. La orogenesis de Grenville en las Sierras Pampeanas occidentals: la Sierra Pie de Palo y su integration al supercontinente Proterozoico. 12th Congreso Geologico Argentino y II Congreso de Exploracion de Hidrocarburos, Mendoza, Actas, 3, 343-358. RAPELA, C.W. & PANKHURST, RJ. 1996. The Cambrian plutonism of the Sierra de Cordoba: preFamatinian subduction and crustal melting? Congreso Geologico Argentino, Actas, Buenos Aires, 5, 491-492. RAPELA, C.W, COIRA, B., TOSELLI, A. & SAAVEDRA, J. 1992. El magmatismo del Paleozoico en el sudoeste de Gondwana. In: GUTIERREZ MARCO, J.G., SAAVEDRA, J. & RABANO, I. (eds) Paleozoico Inferior de Ibero-America, Universidad de Extremadura, 21-68. RAPELA, C.W, PANKHURST, R.J., CASQUET, C., BALDO, E., SAAVEDRA, J., GALINDO, C. & FANNING, C.M. 1998. The Pampean Orogeny of the southern proto-Andes: Cambrian continental collision in the Sierras de Cordoba. In: PANKHURST, RJ. & RAPELA, C.W. (eds) The Proto-Andean Margin of Gondwana. Geological Society, London, Special Publications, 142, 181-217. RAPELA, C.W, PANKHURST, R.J., BALDO, E., CASQUET, C., GALINDO, C., FANNING, C.M. & SAAVEDRA, J. 2001. Ordovician metamorphism in the Sierras Pampeanas: New U-Pb SHRIMP ages in central east Valle Fertil and the Velasco batholith. 3rd South American Symposium on Isotope Geology, Pucon, Actas (CD). RICCI LUCCHI, F. 1984. Influence of transport processes and basin geometry on sand composition. In: ZUFFA, G.G (ed.) Provenance ofArenites. NATO ASI Series, 148, 19-45. RILLER, U. & ONCKEN, O. 2003. Growth of the central Andean plateau by tectonic segmentation is controlled by the gradient in crustal shortening. The Journal of Geology, 111, 367-384. ROBERTSON, S. 1999. BGS Rock Classification Scheme, Volume 2, Classification of metamorphic rocks. British Geological Survey Research Reports, RR 99-02. ROSER, B.P. & KORSCH, RJ. 1986. Determination of tectonic setting of sandstone-mudstone suites using SiO2 and K2O/Na2O ratio. The Journal of Geology, 94, 635-650. ROSER, B.P. & KORSCH, R.J. 1988. Provenance signatures of sandstone-mudstone suites determined using discrimination function analysis of majorelement data. Chemical Geology, 67, 119-139. ROSER, B.P, COOPER, R.A., NATHAN, S. & TULLOCH, AJ. 1996. Reconnaissance sandstone geochemistry, provenance, and tectonic setting of the lower Paleozoic terranes of the West Coast and Nelson, New Zealand. New Zealand Journal of Geology and Geophysics, 39, 1-16.

THE PUNCOVISCANA COMPLEX ROSSI DE TOSELLI, IN., DURAND, F.R., TOSELLI, AJ. &

SARDI, F.G. 1997. Aspectos estratigraficos y geoquimicos comparatives del basamento metamorfico de bajo grado del Sistema de Famatina, Argentina. Revista de la Asociacion Geologica Argentina, 52, 469-480. SALFITY, J.A., OMARINI, R.H., BALDIS, B. & GUTIERREZ, W. 1975. Consideraciones sobre la evolution geologica del Precambrico y Paleozoico del norte argentine. 2nd Congreso Iberoamericano de Geologia Economica, Buenos Aires, 4, 341-361. SCHNEIDER, N. 1993. Das lumineszenzaktive Strukturinventar von Quarzphanokristen in Rhyolithen. Gottinger Arbeiten zur Geologie und Palaontologie, 60, 1-81. SCHWAB, F.L. 1986. Sedimentary 'signatures' of foreland basin assemblages: real or counterfeit. In: ALLEN, PA. & HOMEWOOD, P. (eds) Foreland Basins. International Association of Sedimentologists, Special Publications, 8, 395-410. SCHWARTZ, J.J. & GROMET, L.P. 2004. Provenance of a late Proterozoic-early Cambrian basin, Sierras de Cordoba, Argentina. Precambrian Research, 129, 1-21. SOLLNER, F, BRODTKORB, M.K., MILLER, H., PEZZUTTI, N. & FERNANDEZ, R.F 2000. Dataciones U-Pb en circones de rocas metavolcanicas de la Sierra de San Luis. Revista de la Asociacion Geologica Argentina, 55, 15-22. STUART-SMITH, P.G., CAMACHO, A., SIMS, J.P. ET AL. 1999. Uranium-lead dating of felsic magmatic cycles in the southern Sierras Pampeanas, Argentina: implications for the tectonic development of the proto-Andean Gondwana margin. In: RAMOS, V.A. & KEPPIE, ID. (eds) LaurentiaGondwana Connections before Pangaea. Geological Society of America Special Publications, 336, 87-114. SUTTNER, L.I, BASU, A. & MACK, G.H. 1981. Climate and the origin of quartz arenites. Journal of Sedimentary Petrology, 51, 1235-1246. TANKARD, AJ. 1986. On the depositional response to thrusting and lithosphere flexure: examples from the Appalachian and Rocky Mountains basins. In: ALLEN, PA. & HOMEWOOD, P. (eds) Foreland Basins. International Association of Sedimentologists, Special Publications, 8, 369-392. TAYLOR, S.R. & MCLENNAN, S.M. 1985. The Continental Crust: its Composition and Evolution. Blackwell Scientific, Oxford. THOULKERIDIS, T, CLAUER, N., KRONER, A., REIMER, T. & TODT, W. 1999. Characterization, provenance, and tectonic setting of Fig Tree greywackes from the Archaean Barberton Greenstone Belt, South Africa. Sedimentary Geology, 124, 113-129. TOSELLI, A. I 1990. Metamorfismo del ciclo Pampeano. In: ACENOLAZA, F.G, MILLER, H. & TOSELLI, A.I (eds) El Ciclo Pampeano en el Noroeste Argentino. Serie Correlacion Geologico, Universidad Nacional de Tucuman, 12, 181-197. TOSELLI, A. I & ACENOLAZA, F.G. 1978. Geocronologia de las Formaciones Puncoviscana

415

y Suncho, Provincias de Salta y Catamarca. Revista de la Asociacion Geologica Argentina, 33, 76-80. TOSELLI, AJ. & Rossi DE TOSELLI, IN. 1982. Metamorfismo de la Formation Puncoviscana en las provincias de Salta y Tucuman, Argentina. 5th Congreso Latinoamericano de Geologia, Buenos Aires, Actas, 2, 37-52. TOSELLI, A.I & WEBER, K. 1982. Anquimetamorfismo en rocas del Paleozoico Inferior de Argentina. Acta Geologica Lilloana, 16, 187-200. TURNER, J.C.M. 1960. Estratigrafia de la Sierra de Santa Victoria y adyacencias, Boletin de la Academia Nacional de Ciencias, Cordoba, 41,163-196. UTZMANN, A., HANSTEEN, T.H. & SCHMINCKE, H.U. 2002. Trace element mobility during sub-seafloor alteration of basaltic glass from Ocean Drilling Program site 953 (off Gran Canada). International Journal of Earth Science, 91, 661-679. VALLONI, R. 1984. Reading Provenance from modern sand. In: ZUFFA, G.G. (ed.) Provenance of Arenites. NATO ASI Series, 148, 309-332. VALLONI, R. & ZUFFA, G.G. 1984. Provenance change for arenaceous formations of the northern Apennines, Italy. Geological Society of America Bulletin, 95, 1035-1039. VAN STADEN, A. & ZIMMERMANN, U. 2004. Tillites or ordinary conglomerates? Provenance studies on diamictites of the Neoproterozoic Puncoviscana Formation in NW Argentina. Geoscience Africa 2004, Abstract Volume, University of Witwatersrand, Johannesburg, South Africa, 668-669. VON EYNATTEN, H., BARCELO-VIDAL, C. & PAWLOWSKY-GLAHN, V. 2003. Composition and discrimination of sandstones: a stochastic evaluation of different analytical methods. Journal of Sedimentary Research, 73, 47-57. VON GOSEN, W., LOSKE, W. & PROZZi, C. 2002. New isotope dating of intrusive rocks in the Sierra de San Luis (Argentina): implications for the geodynamic history of the Eastern Sierras Pampeanas. Journal of South American Earth Sciences, 15, 237-250. WENTWORTH, C.K. 1922. A scale of grade and class terms for classifying sediments. The Journal of Geology, 30, 377-392. WILLNER, A.P. 1990. Division tectonometamorfica del basamento del Noroeste Argentino. In: ACENOLAZA, F.G, MILLER, H. & TOSELLI, A.I (eds) El Ciclo Pampeano en el Noroeste Argentino. Serie Correlacion Geologico, Universidad Nacional de Tucuman, 12, 113-159. WILLNER, A.P. & MILLER, H. 1986. Structural division and evolution of the lower Paleozoic basement in the NW-Argentine Andes. Zentralblatt Geologie Palaontologie, Teil I, 9/10, 1245-1255. WILLNER, A.P, MILLER, H. & JEZEK, P. 1985. Geochemical features of an Upper PrecambrianLower Cambrian greywacke/pelite sequence (Puncoviscana through) from the basement of the NW-Argentine Andes. Neues Jahrbuch Geologie Palaontologie Mitteilungen, 33,498-512. WILLNER, A.P, MILLER, H. & JEZEK, P. 1990. Composition geoquimico del basamento sedimentario/

416

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metamorfico de los Andes del NW Argentine (Precambrico Superior/Cambrico Inferior). In: ACENOLAZA, KG., MILLER, H. & TOSELLI, A.J. (eds) El Ciclo Pampeano en el Noroeste Argentina. Serie Correlation Geologico, Universidad Nacional de Tucuman, 12,161-179. WINCHESTER, J.A. & FLOYD, PA. 1977. Geochemical discrimination of different magma series and their differentiation products using immobile elements. Chemical Geology, 20, 325-343. WUELLNER, D.E., LEHTONEN, L.R. & JAMES, W.C. 1986. Sedimentary-tectonic development of the Marathon and Val de Verde basins, West Texas, U.S.A.: a Permo-Carboniferous migrating foredeep. In: ALLEN, PA. & HOMEWOOD, P. (eds) Foreland Basins. International Association of Sedimentologists, Special Publication, 8,347-368. ZHAO, J.X., MCCULLOCH, M.T. & BENNETT, V.C. 1992. Sm-Nd and U-Pb zircon isotope constraints on the provenance of sediments from the Amadeus Basin, central Australia: evidence for REE fractionation. Geochimica et Cosmochimica Acta, 56, 921-940.

ZIMMERMANN, U. 1999. Sedimentpetrographische, geochemische und isotopengeochemische Methoden zur Bestimmung der Beziehung von Provenienz und Ablagerungsraum an aktiven Kontinentalrandern: Das ordovizische Back-arcBecken in der Siid-Puna, Hochland im Nordwesten Argentiniens. PhD thesis, University of Heidelberg, Germany. ZIMMERMANN, U. & BAHLBURG, H. 2003. Provenance analysis and paleotectonic setting of the clastic Ordovician deposits in the southern Puna, NW Argentina. Sedimentology, 50, 1079-1104. ZIMMERMANN, U, LUNA TULA, G, MARCHIOLI, A., NARVAEZ, G, OLIMA, H. & RAMIREZ, A. 2002. Analisis de la procedencia de la Formacion Falda Cienaga (Ordovicico Medio, Puna Argentina) por petrografia sedimentaria, elementos trazas e isotopia de Nd. Asociacion Argentina de Sedimentologia Revista, 9, 1-24. ZUFFA, G.G., GAUDIO, W. & ROVITO, S. 1980. Detrital mode evolution of the rifted continental-margin Longobucco Sequence (Jurassic), Calabrian arc, Italy. Journal of Sedimentary Petrology, 50,51-61.

Kinematic history of western Marie Byrd Land, West Antarctica: direct evidence from Cretaceous mafic dykes CHRISTINE S. SIDDOWAY1, LOUIS C. SASS III1 & RICHARD P. ESSER2 1 Department of Geology, The Colorado College, 14 E. Cache la Poudre, Colorado Springs, Colorado 80903, USA (e-mail: [email protected]) 2 New Mexico Geochronological Research Laboratory, New Mexico Bureau of Mines and Mineral Resources, 801 Leroy Place, Socorro, NM 87801, USA Abstract: Intracontinental deformation occurred in West Antarctica during the final stages of plate convergence along the Cretaceous Gondwana margin. In western Marie Byrd Land, 115 Ma to 95 Ma A-type granitoids and mafic dykes record a change in plate kinematics. The magmatism typically is viewed as a record of extension leading to orthogonal break-up between New Zealand and Marie Byrd Land by c. 67 Ma. This paper presents new kinematic and 40Ar/39Ar age data for a mafic dyke array in the Ford Ranges, a region >1000 km2 dominated by plutonic and metamorphic bedrock. The mean dyke trend of N16W corresponds to a maximum finite strain axis orientated N74E, highly oblique to the N58E-trending margin and to on-land crustal structures defined from airborne geophysics. 40 Ar/39Ar emplacement ages for most dykes fall between 114 Ma and 97 Ma, coeval with emplacement of a gneiss dome at 101-96 Ma and with development of mylonitic shear zones at 100-95 Ma in coastal western Marie Byrd Land. The oblique orientation of maximum finite strain with respect to large faults, geophysical lineaments and the rifted margin of western Marie Byrd Land is consistent with transcurrent tectonics along this segment of the Gondwana margin at c. 100 Ma.

Continental extension occurred within the western Marie Byrd Land terrane (Fig. 1) during final stages of plate convergence along the Gondwana margin in West Antarctica-New Zealand. The event is recorded by c. 107 Ma mafic dykes in central Marie Byrd Land (Storey et al 1999) and by c. 115-95 Ma A-type granite plutons (Weaver et al 1992; 1994) overprinted by high strain zones having 101-94 Ma 40 Ar/39Ar mineral cooling ages (Richard et al 1994; Siddoway et al 20040) in the Ford Ranges of western Marie Byrd Land (Fig. 2). In the absence of regional kinematic data, magmatism and mid-crustal flow typically are interpreted in terms of pure shear extension orthogonal to the continental margin (e.g. Richard et al 1994; Storey et al 1999). A persistent problem for kinematic interpretation in the region is the comparative homogeneity of bedrock units, absence of dynamic fabrics in plutonic rocks, lack of exposure of large-scale faults or crustalscale shear zones and the lack of precise age control upon mesoscopic brittle structures (Luyendyk et al 2003; Siddoway et al 20040). Comprehensive structural analysis, including study of anisotropy of magnetic susceptibility (AMS) fabrics, has been conducted previously only on the singular exposures of middle crustal

rocks that form the Fosdick Mountains migmatite dome (Siddoway et al 20046). Mafic dykes present throughout the Ford Ranges offer the means to determine directly the regional finite strain ellipsoid during Cretaceous magmatism in the Ross Province of western Marie Byrd Land. This paper presents a geometrical analysis of the regional dyke array, together with 40Ar/39Ar age data that provide constraints on timing of dyke emplacement. The kinematic record from mafic dykes is considered together with kinematic information from brittle faults (Luyendyk et al 2001; 2003) and ductile structures (Richard et al 1994; Siddoway et al 20046) in the Ford Ranges, in order to develop a transcurrent model for Early Cretaceous tectonism along this segment of the Gondwana margin in Albian time.

Geological background Marie Byrd Land in West Antarctica (Dalziel & Elliot 1982; Storey et al 1988; Bradshaw et al 1997) forms the northern flank of the West Antarctic rift system (Storey et al 1999) and has geological affinities with New Zealand and northern Victoria Land (Ireland et al 1994; Muir et al 1994; Bradshaw et al 1997). Pankhurst et al

From: VAUGHAN, A. P. M., LEAT, P. T. & PANKHURST, R. J. (eds) 2005. Terrane Processes at the Margins of Gondwana. Geological Society, London, Special Publications, 246, 417^38. 0305-8719/$15.00 © The Geological Society of London 2005.

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Fig. 1. Tectonic reconstruction of the Marie Byrd Land-New Zealand sector of the Cretaceous Gondwana margin at c. 100 Ma. The rifted margin corresponds to the c. 1500 m contour (dashed-line pattern), which defines a linear to curvilinear trace for a distance >1500 km bordering the Ross Sea and western Marie Byrd Land. The margin trends N58E along Ruppert Coast, changing to N75E along Hobbs Coast. The dramatic increase in depth along a linear or small circle trace, together with the very close fit with the -1500 m contour for Campbell Plateau, suggest structural controls upon the rifted margin. CT, Campbell Trough; EP, Edward VII Peninsula; HC, Hobbs Coast; IB, Iselin Bank; RC, Ruppert Coast. Based on Sutherland (1999).

(1998) distinguish two provinces within Marie Byrd Land; the eastern Amundsen Province of arc character and the western Ross Province of continental affinity. Mafic dykes pervade both provinces; however, the prevalent orientation in the Ross Province (this paper) differs from that in the Amundsen province (Storey et al. 1999). Storey et al. (1999) determined that the Amundsen Province array exposed along Ruppert Coast (Fig. 1) is orientated N80W, subparallel to the coast except where dykes intrude pre-existing N-S structures. The Palaeozoic Ross Province, the focus of this study, comprises folded, low greenschistgrade metagreywackes of the Swanson Formation (Bradshaw et al 1983), intruded by

Ford Granodiorite (I-type; Weaver et al 1991; Pankhurst et al 1998) (Fig. 2). Swanson Formation represents detritus shed from the early Palaeozoic Ross Orogen (Ireland et al 1994; Pankhurst et al 1998). In middle Cretaceous time, A-type granitoids collectively known as Byrd Coast Granite were emplaced; these have a geochemical character indicative of a continental extensional province (Weaver et al 1992; 1994; Adams et al 1995). Furth evidence of tectonic divergence comes from a discordance in Cretaceous palaeomagnetic poles between East Antarctica and Marie Byrd Land (DiVenere et al 1994; Luyendyk et a 1996); gravity data that delimit a 8-9 km contrast in crustal thickness between the Ross

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Fig. 2. Geological map of the Ford Ranges, western Marie Byrd Land. Location and trend of mafic dykes are shown with orange bars generally, and with brown bars in the Fosdick Mountains. Sample sites for dykes used for 40Ar/39Ar age determination are numbered. The Ford Ranges comprise early Palaeozoic greywacke and argillite of the Swanson Formation intruded by Devonian Ford Granodiorite and Cretaceous Byrd Coast Granite. These formations are, in turn, intruded by mafic dykes; felsic dykes intrude the Byrd Coast Granite only. Inferred faults correspond with contrasts in metamorphic grade between ranges, geophysical lineaments or boundaries, and zones of penetrative brittle deformation.

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Sea and central Marie Byrd Land (Luyendyk et al. 2003); Late Cretaceous mineral cooling and apatite fission track ages throughout the Ross Province (Richard et al 1994; Lisker & Olesch 1998; Pankhurst et al 1998); and, in the northern Ford Ranges, a migmatite-cored gneiss dome in the Fosdick Mountains, emplaced and cooled rapidly between 105 Ma and 94 Ma (Richard et al 1994). Textural and petrological evidence for rapid decompression, together with the thermal history, indicate that diapirism played a role in dome emplacement (Siddoway et al 20046). Offshore, high-angle normal faults have been imaged by marine seismic survey and, along one of them, samples of mylonitized Byrd Coast Granite were recovered by dredge (Luyendyk et al 2001). The mylonitic gneisses formed in a shear zone at c. 98 Ma and underwent rapid cooling as a consequence of tectonic exhumation (Siddoway et al 20040). On land, field investigation and analysis of airborne geophysics data identify a regional fault pattern of dextral and sinistral faults in a conjugate geometry (Fig. 2; Luyendyk et al 2003). In this region dominated by isotropic plutonic rocks, the primary means to obtain kinematic data directly is through analysis of mesoscopic-scale brittle structures including minor faults, shear fractures and mafic dykes that cut Byrd Coast Granite and older units (Luyendyk et al 2003). The paucity of reliable geological markers to quantify fault offsets or provide age control make the mafic dykes of great value for strain studies, using the assumption that dykes emplaced in previously unfractured rock propagate in a plane normal to the maximum principal finite strain (i.e. least compressive stress direction, Tsunakawa 1983; Best 1988). Recent efforts to integrate findings from brittle kinematic studies throughout the Ford Ranges (Luyendyk et al 2003) with results from the Fosdick Mountains dome (Richard et al 1994) found a transcurrent model to be most compatible with regional structural patterns (Siddoway et al 20046). The mafic dykes throughout the Ford Ranges provide a direct means to determine age and kinematic setting for the time of their emplacement and, consequently, to determine whether their kinematics are consistent with transtension (e.g. Luyendyk et al 1992, fig. 6) or orthogonal extension (e.g. Storey et al 1999) along this sector of the Cretaceous Gondwana margin.

The Ford Ranges dyke array: description and structural geometry Present in both the Carboniferous Ford Granodiorite and Cretaceous Byrd Coast Granite, mafic dykes of the Ford Ranges are typically 2-4 m thick and have steeply dipping, planar contacts with host rock (Figs 3a, b; Table 1). They have great lateral extent along-strike; however, they are widely spaced at 15 m to 1500 m. Thicknesses range from 0.5 m to 14 m. The dykes typically are dolerite, with finegrained plagioclase laths within an aphanitic groundmass of anhedral to subhedral hornblende and pyroxene. Several dykes are plagioclase- and pyroxene-porphyritic. One unusual dyke contains large pyroxene phenocrysts up to 3 cm in length; the rock is vesicular and developed columnar joints (Fig. 3c). Petrographic study shows that typical groundmass consists of varying proportions of augite, amphibole, biotite and plagioclase. Hornblende is commonly zoned, with some alteration to actinolite: Epidote, chlorite or calcite may be present, particularly in dykes with sheared margins. Pyrite and non-magnetic opaque grains are abundant accessory minerals and accessory apatite is present in some samples. Ford Ranges dyke orientations, Fosdick Mountains excluded, are summarized in Figure 4. Figures 4a and b provide a comparison of data from the northern versus southern Ford Ranges as a means to assess regional variations. Fosdick Mountains data are considered separately (Fig. 5) because migmatite gneisses representative of middle crust (Siddoway et al 2004Z?) are exposed in that range. All structural attitudes were measured with respect to geographical north at outcrops throughout the Ford Ranges between 142° W and 146° W longitude (Fig. 2). A prevalent mafic dyke array in the southern Ford Ranges is orientated N16W with subvertical dips (Fig. 4a), suggesting a predominant ENE-WSW stretching direction at the time of dyke emplacement. This population is also present in the northern Ford Ranges; however, a conjugate dyke array distinguished by moderate dips to NE and SW (Fig. 4b) is also present. Seven of thirteen of the NE-striking dykes intrude brittle faults with sinistral sense, low-raking striae; thus, it is probable that these dykes intruded upon contemporaneous or preexisting structures. The eastern Phillips Mountains (Fig. 2) host a suite of felsic porphyry dykes cutting Byrd Coast Granite and these steeply dipping felsic dykes strike NNW-SSE (Fig. 4c), consistent with the prevalent mafic dyke orientation for the region. The small

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dataset in Figure 4d summarizes the foliation and lineation fabrics of a rare exposure of mylonitized Ford Granodiorite at Mt Cooper in the Sarnoff Range. The mylonite data are included here because a mylonite gneiss sample investigated using 40Ar/39Ar thermochronology yields similar kinematic and age results as mafic dykes. The 40Ar/39Ar age for dynamically recrystallized biotite in the normal-sense shear zone, reported below, is 97 Ma, concordant within error with the 96 Ma ages of the youngest dykes. Finally, Figure 4e shows the orientations of mafic dykes that intrude Byrd Coast granite plutons and, as such, are sure to be Albian aged or younger. Evident in Figure 5, dyke geometries in the Fosdick Mountains contrast with those elsewhere. The mean orientation of 102/80 SW is nearly orthogonal to the dyke trends in the broader Ford Ranges summarized in Figure 4. 40

Ar/39Ar age investigation

Fig. 3. Outcrop photos illustrating dyke geometry and extent, (a) Dyke of width 2 m is orientated 340/89 NE. The dyke is cut by a NE-striking brittle fault with low-raking striae; the offset is 5 m in a sinistral sense. 40Ar/39Ar sample 9N23-4 was collected from this dyke at Mt Ralph, (b) Large, steep dyke at Andrews Ridge orientated 000/81 E; thickness is 4 m (not analysed), (c) Wide dyke of columnar-jointed, vesicular plagioclase basalt, orientated 355/82 E. 40Ar/39Ar sample L9N17-1 was collected from this site at Mt Little.

Mafic dykes and host rocks from 11 sites distributed widely within the Ford Granodiorite (Fig. 2) were sampled for 40Ar/39Ar age dating. The aphanitic texture of the dyke rock necessitated the use of groundmass concentrate separates as the phase to be dated. Biotite and muscovite mineral separates were obtained from four hostrock samples, including sample 9D7-4 from a narrow mylonitic shear zone at Mt Cooper. Samples were analysed by the resistancefurnace incremental-heating age spectrum method at the New Mexico Geochronology Research Laboratory (NMGRL). Analytical methods are summarized in Table 2. The nine groundmass concentrate samples yield slightly to severely discordant age spectra (Fig. 6; Table 3). Five of the samples (9D9-1, 9N17-2, 9N23-4, 9N25-8 and L9N16-2) yield spectra with initial anomalously young apparent ages at the lowest temperature steps, followed by monotonically increasing ages at the intermediate to highest temperature steps. For these samples, high K-Ca ratios generally coincide with older apparent ages. For most samples, isotope correlation diagrams (also known as inverse isochron diagrams) yield trapped 40 Ar/36Ar ratios equivalent to present-day atmosphere (295.5) and ages are analytically indistinguishable from their respective spectrum-weighted mean ages. For most of the groundmass concentrate samples, rising age spectra and correspondence of initial young ages to the lowest radiogenic yields (Figs 6, 8) are attributed to alteration of finer-grained groundmass, with non-diffusive

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Table 1. Summary of field relationships for dykes used for 40Ar/39Ar study Location, site #, and strike/dip of margin Mt Little 143°50' W 77°00' S Site 2 in Figure 2

Schematic diagram

Sample number and description L9N16-2 Dolerite with acicular plagioclase in fine-grained groundmass.

Age and cross-cutting relationships 142.0 ± 0.7 Ma Cut by NE faults with strike-slip striae.

300/90

Mt Little 143°50'W76°58'S Site 2 in Figure 2 Various

Mt Gilmore 144°35' W76°56'S Site 3 in Figure 2 241/80

Mt Ralph 144°31'W76°58'S Site 4 in Figure 2 340/89

Ranney Nunatak 143°55' W76°53' S Site 5 in Figure 2

L9N17-1 Vesicular basalt intrusion with columnar joints. Pyroxenes up to 10 cm; fine-grained plagioclase laths.

9N20-1 Fine-grained light grey dolerite hornblende and pyroxene; contains resorbed Qz reaching 8 mm and xenoliths with reaction rims.

96.38 ± 3.54 Ma Intrudes E-W and NE-SW fractures. It is not cut by faults.

102.2 ± 1.7 Ma Intrudes brittle fault with shallow fault striae.

9N23-4 Medium-grained dolerite with augite >1 mm. Px replaced by Hbl. Contains granite xenoliths.

274.2 ± 13.4 Ma

9N25-8 Very fine-grained dolerite with plagioclase, hornblende ± pyroxene.

136.9 ± 2.6 Ma Cut by high angle fault with right separation.

305/90

Peak 1180 144°22' W76°56'S Site 6 in Figure 2

L9N27-1 Fine-grained dolerite with biotite after hornblende. Contains granitic xenoliths.

342/86

Mt Dolber 145°28'W77°07'S Site 1 in Figure 2 225/75

9D9-1 Fine-grained dolerite with biotite after hornblende.

110.6 ±3.1 Ma Dyke is cut by a chloritic shear zone, shown with cross-hatch pattern.

96.15 ± 2.08 Ma

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Fig. 4. Structural data for dykes and a shear zone that cut granitoids of the Ford Ranges. Attitude of dyke margin or of mylonitic foliation are summarized on equal-area stereographic diagrams. The Crevasse Valley Glacier is used to subdivide the Ford Ranges in to southern and northern regions for purposes of structural analysis, to assess regional variations. For locations, refer to Figure 2. Diagrams in this and subsequent figures were prepared using Stereonet v. 6.2.X by R. W. Allmendinger. (a) Southern Ford Ranges stereoplot and rose diagram; (b) northern Ford Ranges stereoplot and rose diagram, exclusive of Fosdick migmatite gneiss dome (see Fig. 5); (c) summary diagram of felsic porphyry dykes intruding Byrd Coast Granite, Phillips Mountains; (d) summary diagram of mylonitic foliation and mineral stretching lineation within narrow shear zones at Mt Cooper (Fig. 2). The 40Ar/39Ar biotite age of 97 Ma for a mylonite sample is presented in Table 2. (e) Orientations of mafic dykes that cut Byrd Coast Granite plutons; (f) summary rose diagram for the Ford Ranges showing the dominant regional dyke orientation of N16 W.

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Fig. 5. Equal-area stereographic diagrams for mafic dykes that cut foliation and folds in the Fosdick Mountains gneiss dome. The attitudes of dyke margins are summarized. The dykes trend WNW-ESE, with steep dips to both NNE and SSW. The Fosdicks dyke array is nearly orthogonal to systematic dykes of the broader Ford Ranges.

argon loss from matrix phases (e.g. Poland et al 1993). Complex heating behaviour for dyke samples was also noted by Richard et al. (1994). Textural evidence of breakdown of pyroxene and hornblende to retrograde products, exacerbated where faults overprint dykes (Table 1), is a sign of this alteration. Due to the effects of alteration, the preferred isochron or plateau ages within the range 142-97 Ma are considered to be minimum ages for most samples. Where MSWD analysis (cf. Mahon 1996) on each plateau segment or inverse isochron indicates that the single population criterion was not fulfilled at 95% probability, the errors on the weighted mean age or isochron age are adjusted (Tables 2-4). Although the 40Ar/39Ar age determinations are not of high quality due to minor amounts of alteration, the results nonetheless establish the time of emplacement of most dykes to be Albian, consistent with the observation that the dykes cut 102-95 Ma age Byrd Coast Granite and older rocks. Two dykes give older ages of 137 Ma and 146 Ma. Moreover, all of the dated dykes form part of a systematic structural array with a mean strike of N16W. Biotite or muscovite mineral separates were prepared from plutonic host rock bordering mafic dykes at four sites (Fig. 7), together with a biotite separate for a sample from a 3 m wide mylonite shear zone (Fig. 8) in the southern

Ford Ranges. The spectra for each mica are relatively well-behaved with only minor amounts of 40Ar* loss in the lowest temperature steps, probably the result of minor chloritization. The weighted mean of the flattest portion of each biotite and muscovite age spectrum is interpreted as the crystallization age of the sample. In the case of sample 9D7-4, the weighted mean (96.92 ± 0.34 Ma) is interpreted as the recrystallization age of the biotite within the mylonite zone. A similar age comes from an undeformed mafic dyke in the Fosdick Mountains gneiss dome, investigated in a previous study (Richard et al. 1994). 40Ar/39Ar hornblende results yielded concordant plateau and isochron ages of 97.8 ± 0.1 Ma and 96.4 ± 1.1 Ma (Table 2). A complex hornblende age spectra from a Chester Mountains sample (Fig. 2; Table 2) gave apparent ages varying irregularly between 148 Ma and 102 Ma, and a total gas age of 122 Ma. The complexity in 40Ar/39Ar results is attributable to complexly zoned hornblende with actinolitic rims (Richard et al. 1994). 40 Ar/39Ar weighted mean ages for biotite and muscovite separated from dyke wall rock samples range from 351 ± 2 Ma to 342.4 ± 1 Ma (Table 2; Fig. 7), a confirmation that the plutonic rock hosting the mafic dykes is Ford Granodiorite (cf. Pankhurst et al. 1998).

425

MARIE BYRD LAND KINEMATIC HISTORY Table 2. Summary of 40Ar/39Ar age determinations for Ford Ranges mafic dykes Sample number

Sample locality

Site on Figure 2

Preferred age*

n

%39Ar

MSWD

K-Ca

Ages determined from groundmass concentrates Mt Dolber, Sarnoff Range 9D9-1 9N17-1 Mt Little, west central Ford Ranges Mt Little, west central Ford Ranges 9N17-3 Mt Gilmour, Denfield Mountains 9N20-1 Mt Gilmour, Denfield Mountains 9N20-2 9N23-4 Mt Ralph, Denfield Mountains Ranney Nunatak, Gutenko Nunataks 9N25-8 Peak xllSO, Denfield Mountains 9N27-1 L9N16-2 Mt Little, west central Ford Ranges

1 2 2 3 3 4 5 6 2

isochron isochron isochron plateau isochron isochron isochron isochron plateau

8 9 9 3 9 5 8 7 6

Ages determined from mineral separates 9D7-4 Mt Cooper, Sarnoff Range L9N15-1 Mt Little, west central Ford Ranges L9N17-2 Mt Little, west central Ford Ranges L9N17-3 Mt Little, west central Ford Ranges L9N17-3 Mt Little, west central Ford Ranges

7 2 2 2 2

plateau, bt plateau, mu plateau, mu plateau, mu plateau, bt

9 7 8 8 3

Chester-2 Mutel-1

8 9

total gas plateau

Sample details are provided in Richard et al. (1994)

Chester Mountains Fosdick Mountains 40

88.1

50.3 68.7* 109* 20.1* 209* 0.8 21.4* 181* 6.1*

93.9 94.7 92.6 93.3 43.3

3.4* 2.1 0.8 3.2* 8.0*

71.6

4.20

0.83 21.6 368.7 28.4 21.6 12.2

Age

±2o

96.15 96.38 40.82 102.2 113.5 274.2 136.9 108.8 142.0

2.08* 3.54* 2.31* 1.7* 5.8* 13.4* 2.6* 7.0* 0.7*

96.92 342.4 348.2 347.3 350.8

0.34* 1.0* 0.9* 1.2* 2.1*

122 96.4

1.1

36

* Isochron ages, Ar/ Arj and MSWD values calculated from regression results obtained by the methods of York (1969). Decay constants and isotopic abundances follow Steiger & Ja'ger (1977). * Two-o errors. All final errors are reported at ±2o, unless otherwise noted. Weighted mean error is calculated using the method of Taylor (1982). * Mean standard weighted deviation (MSWD) outside 95% confidence interval. Weighted mean age is calculated by weighting each age analysis by the inverse of the variance. MSWD values are calculated for n - 1 degrees of freedom for plateau and preferred ages. Sample preparation: Groundmass concentrates, biotites and muscovites separated using standard techniques of crushing, sieving, magnetic separation and hand-picking. Groundmass concentrates were packaged and irradiated in machined Al discs for 7 hours in D-3 position, Texas A&M University Research Reactor. Biotite and muscovite mineral separates were packaged and irradiated in machined Al discs for 24 hours in L67 position, Ford Research Reactor, University of Michigan. The neutron flux monitor used is Fish Canyon Tuff sanidine (FC-1) of assigned age - 27.84 Ma (Deino & Potts 1990) relative to standard Mmhb-1 at 520.4 Ma (Samson & Alexander 1987). The Richard et al. (1994) horneblende ages are recalculated using the 27.84 flux monitor age, for conformity with NMGRL data. Sample procedures: The instrument used is a Mass Analyzer Products 215-50 mass spectrometer on-line with automated allmetal extraction system. All samples were step-heated in Mo double-vacuum resistance furnace. Duration of heating is 6-8 minutes. Reactive gases were removed with 2 SAES GP-50 getters, one operated at c. 450 °C and one at 20 °C, during the 6-minute reaction. Gases were also exposed to a W filament operated at c. 2000 °C. Analytical parameters: Electron multiplier sensitivity is averaged 3.06 X 10~16 moles pA"1 for samples analysed by the laser. The total system blank and background for the step-heated samples averaged 1040, 2.7,1.3,1.1, 3.4 X 10~17 moles. J-factors were determined to a precision of ± 0.1% by CO2 laser-fusion of four single crystals from each of four or six radial positions around the irradiation tray. Correction factors for interfering nuclear reactions were determined using K-glass and CaF2 and are as follows: Texas A&M: Michigan:

(40Ar/39Ar)K = 0.0002 ± 0.0003 (36Ar/37Ar)Ca = 0.00028 ± 0.00001 (39Ar/37Ar)Ca = 0.00089 ± 0.00003 (40Ar/39Ar)K = 0.0262 ± 0.0003 (36Ar/37Ar)Ca = 0.00028 ± 0.00001 (39Ar/37Ar)Ca = 0.00078 ± 0.00003

Discussion Throughout most of the Ford Ranges, mafic dykes record regional ENE-WSW stretching orthogonal to dyke margins (Fig. 4f). The observation that dykes cutting Byrd Coast Granite of 102-95 Ma age (Fig. 4c) and dykes emplaced within older rocks (Figs 4a, b) are co-parallel

makes it unlikely that dykes commonly intruded upon pre-existing fractures unrelated to the middle Cretaceous state of stress (e.g. Best 1988). The exception is NE-striking, moderately steeply dipping dykes. These intrude brittle faults with shallowly plunging striae, an indication that dykes exploited pre-existing structures. The mesoscopic normal-sense mylonitic shear

426

C. S. SIDDOWAY ETAL.

MARIE BYRD LAND KINEMATIC HISTORY

427

Fig. 6. 40Ar/39Ar age spectra and inverse isochron diagrams for representative groundmass concentrate separates for dolerite dykes. Each of the nine groundmass concentrates yielded slightly to severely discordant age spectra; thus both plateaux or weighted mean ages and inverse isochron ages are calculated. For most samples, the inverse isochron ages are analytically identical to plateau ages. The preferred age for each sample is noted here and summarized again in Figure 9. Diagrams for samples 9N17-3, 9N20-2 and 9N23-4 are not shown. The samples produced very disturbed spectra indicative of extensive alteration, have unacceptable error in 2o and yield results of overall low quality, (a, b) Sample L9N16-2; (c, d) sample L9N27-1; (e, f) sample L9N25-8; (g, h) sample L9N20-1; (i, j) sample L9N17-1; (k, 1) sample 9D9-1. All errors are reported at 2o.

zones at Mt Cooper (Fig. 2) record stretching orientated c. N75E (Fig. 4d) at 96.92 ± 0.34 Ma, according to the 40Ar/39Ar age on biotite (Fig. 6) forming part of the mylonitic fabric. Within the Fosdick gneiss dome, folds, ductile fabrics and AMS strain axes for migmatite gneisses indicate a finite strain axis orientated N65E during high temperature metamorphism and dome emplacement (Siddoway et al 20045). The stretching direction is oblique to the N80Wtrending Balchen Glacier Fault that bounds the dome on its north side (Fig. 2). The consistent results from brittle and ductile criteria from differing levels of crustal exposure provide firm evidence for a NE to ENE maximum finite strain direction across the Ford Ranges at the time of dyke emplacement. The orientation of felsic and mafic dykes (Figs 4c, e) cross-cutting Cretaceous Byrd Coast Granite plutons in the

Phillips, Clark, Allegheny Mountains (Fig. 2) help support this interpretation. The timing of regional dyke emplacement and NE to ENE stretching is well constrained by cross-cutting relationships of the parallel array of dykes that intrude both Byrd Coast Granite and Ford Granodiorite (Figs 4e, a, respectively) together with the 40Ar/39Ar results for Ford Ranges dykes (Figs 6, 9). The geological and thermochronological data show that the prevalent dyke ages are Aptian, with two older results in the latest Jurassic and earliest Cretaceous. The structural data presented here provide a clear record of the regional kinematic setting at that time and prove that the c. N70E regional stretching direction was not orthogonal to the (present-day) rifted margin of western Marie Byrd Land as has been thought previously (e.g. Richard et al

C. S. SIDDOWAY ETAL.

428

Table 3. Results of 40Ar/39Ar analysis of groundmass concentrate separates for mafic dyke samples 40

Ar/39Ar

37

Ar/39Ar

36

Ar/39Ar X 10'3r

39 ArK K-Ca X 10-16 mol

40

Ar*

39

Ar

Age (Ma)

±lo (Ma)

9D9-1, E6:130, 19.10 mg whole rock, J =: 0.00076235 ± 0.10%, NM-130, Lab# = 51769-02 703 -0.027 7.3 -7.5 0.0 A V 625 192.7 0.0697 33.7 0.2 0.477 0.25 B V 700 213.7 1.838 450 — -0.007 — 0.2 7779 C V 750 69.17 0.0000 7.27 96.5 7.0 1.139 0.45 9.350 D V 800 77.55 0.57 0.429 95.6 7.2 0.8915 4.026 E V 875 80.13 0.41 97.5 13.8 1.255 6.467 22.0 F 975 73.51 0.65 0.7900 2.360 35.9 99.1 34.4 G 1075 72.97 0.31 104.0 99.0 94.0 1.665 2.889 H 1250 73.45 39.77 700.0 0.069 57.9 7.365 70.5 I V 1650 91.84 0.38 174.5 total gas age 71 = 9 0.40 steps F-H plateau MSWD = ll.lt 161.9 92.8 n =3 40Ar/36Ar = 500 ± 124* isochron MSWD - 50.33t n =8

-20.5 96.5 — 700.25 705.6 96.11 96.91 97.50 705.40 97.91 96.85 96.15

737.2 6.7 — 0.75 2.2 0.21 0.20 0.21 0.37 0.42* 0.84* 2.08*

9N17-1, Gl:130, 27.33 mg whole rock, J = 0.00075944 ± 0.10%, NM-130, Lab# = 51779-01 — 0.760 — 7.3 0.7 0.0000 A V 625 — 0.7719 0.66 70.5 7.9 B V 700 833.9 2525 4.25 796.7 7.2 4.45 55.0 3.5 C V 750 138.1 0.4424 77.4 5.0 707.7 0.94 9.65 D V 800 110.4 0.5426 7.2 57.07 0.945 74.6 5.4 E V 875 101.0 0.4202 83.74 92.5 75.0 48.3 0.5203 0.98 F 975 99.00 39.81 53.5 0.86 86.1 71.4 0.5945 G 1075 84.29 75.24 52.1 0.18 76.4 93.9 H 1250 93.18 2.873 14.1 0.078 55.5 100.0 198.7 I 1650 130.7 6.574 231.7 0.71 total gas age n =9 n =4 steps F-I 212.2 0.69 91.6 plateau MSWD - 26.9* 40Ar/36Ar = 311 ±27* n =9 isochron MSWD - 68.7*

— 776.4 706.6 704.94 700.4 99.07 96.84 95.22 97.29 98.30 96.94 96.38

— 6.6 7.0 0.44 7.6 0.32 0.22 0.29 0.56 18.80* 1.51* 3.54*

9N17-3, E5:130, 19.86 mg whole rock, J -0.000761894 ±0.10%, NM-130, Lab# = 51768-01 — 0.077 -7.6 0.0 0.0000 55557 A V 625 — 0.557 752.7 5.0 7.2 0.33 B V 700 251.2 1.527 230.4 7.0 79.5 5.2 0.4943 2.95 C V 750 84.89 0.20 77.0 9.6 3.22 602.2 D V 800 199.8 2.513 0.8472 3.30 0.60 42.5 74.2 709.5 E V 875 56.59 50.3 42.6 0.3844 23.07 20.5 7.3 F V 975 34.52 1.4 79.5 64.6 0.3737 26.78 16.1 G 1075 38.36 23.7 0.42 74.9 97.0 1.206 34.95 H 1250 40.69 113.4 0.098 46.8 100.0 2.19 I 1650 62.22 5.189 73.1 0.90 total gas age n =9 57.4 0.77 42.0 plateau MSWD - 1.4 n =3 steps G-I 40Ar/36Ar = 278 ± 33* isochron MSWD - 109* n =9

— 27.6 23.0 30.7 33.00 37.74 41.43 41.44 39.77 38.55 41.41 40.82

— 4.4 7.7 7.9 0.70 0.75 0.16 0.15 0.98 5.42* 0.34* 2.31*

9N20-1, G3:130, 21.37 mg whole rock, J = 0.000758201 ±0.10%, NM-130, Lab# = 51781-01 4.7 0.1077 77022 7.65 7.3 0.6 A V 625 5094 0.2834 75.4 7.5 595.5 26.5 7.7 B V 0 700 240.4 765.4 72.7 7.7 63.0 77.4 C V 0 750 134.3 0.4443 777.7 9.59 0.50 70.4 74.5 D V 0 800 117.2 1.012 5.64 0.65 63.5 77.9 0.7859 750.0 E V 875 121.3 29.2 3.4 61.5 28.3 0.1523 167.0 F 975 128.1 65.64 65.8 6.3 0.0806 79.8 51.7 G 1075 95.79 106.4 127.9 3.1 66.9 89.5 H 1250 114.1 0.1642 2.0 47.7 700.0 350.6 29.6 / V 1650 193.00.2572 281.4 3.4 total gas age n =9 201.4 71.6 4.2 plateau MSWD - 20.1* n =3 steps F-H 40Ar/36Ar = 308 ± 9* isochron MSWD = 7.5* n =6

56.0 56.0 772.79 709.57 702.43 104.64 101.60 101.54 706.95 102.1 102.2 100.5

96.9 7.2 0.56 0.57 0.67 0.44 0.25 0.39 0.75 2.0* 1.7* 1.7*

ID

T

429

MARIE BYRD LAND KINEMATIC HISTORY Table 3. (continued) 40

Ar/39Ar

37

Ar/39Ar

36

Ar/39Ar x 10-3r

39 ArK K-Ca x 10-16 mol

40Ar*

39

Age (Ma)

±lo (Ma)

9N20-2, H9:130, 20.85 mg whole rock, J -0.000751144 ±0.10%, NM-130, Lab# = 51788-01 0.0432 748.7 29.6 25.2 74.5 A V 625 295.6 77.8 18.4 92.5 70.3 B 700 93.24 0.0278 23.68 113.9 15.94 91.2 750 91.27 0.0473 42.6 10.8 94.8 C 800 85.67 79.3 93.8 D V 0.7786 59.89 5.46 2.9 875 700.0 48.2 0.7400 775.5 7.30 3.6 94.5 E V 7.87 95.4 975 793.2 484.7 7.9 26.0 F V 0.2683 653.7 7.77 0.87 74.3 95.9 7075 225.4 G V 0.6308 73.7 97.9 H V 7250 379.5 0.6975 939.8 4.05 0.73 700.0 0.59 73.4 I V 7650 360.0 0.8578 7055 4.26 204.2 14.4 total gas age n =9 156.6 76.7 plateau MSWD = 1.2 n =2 steps B-C 16.3 40Ar/36Ar - 262 ± 29 * isochron MSWD - 209^ n =9

98.0 113.25 113.66 89.79 64.7 66.8 43.2 55.8 64.3 107.2 113.5 113.5

7.5 0.28 0.24 0.90 7.9 3.7 4.8 3.0 3.3 1.3* 0.6* 5.8*

9N23-4, G6:130, 26.11 mg whole rock, J -0.000757059 ±0.10%, NM-130, Lab# - 51784-01 0.77 0.2 A V 625 28978 4.633 97322 0.223 0.6 4044 2.72 72.7 7.9 B V 700 7360 3.429 0.75 4.0 7828 2.46 0.080 6.8 C V 750 2480 6.351 800 7747 3.09 0.080 76.4 6.5 D V 6.390 3230 70.4 4.75 0.72 875 623.9 4.163 7423 32.6 E V 51.2 F 0 975 1.052 49.4 0.49 87.7 244.7 101.9 0.34 89.1 69.0 G 0 1075 246.8 1.517 91.10 21.5 87.4 799.7 75.7 0.72 80.2 H V 0 7250 295.6 4.203 0.092 85.4 700.0 272.0 5.547 22.5 7 V 0 7650 736.0 121.1 0.30 total gas age n-9 n =2 0.44 plateau MSWD - 46.5* steps F-G 70.9 58.5 40Ar/36Ar = 286 ± 5* isochron MSWD - 0.8 n-5

207 3460 272.9 75.5 33.2 276.8 240.5 70.8 4.7 259.5 0.72 271.90 278.23 0.59 298.90 0.99 293.52 0.80 276.8 16.8* 6.2* 275.7 13.4* 274.2

9N25-8, G2:130, 19.16 mg whole rock, J -0.000760011 ±0.10%, NM-130, Lab# = 51780-01 3.758 77057 0.74 2.7 0.9 A V 625 3336 0.886 700 434.7 7.377 7257 4.04 0.37 B V 0 74.6 4.9 750 404.7 3.657 0.74 22.7 C V 7058 2.93 7.8 4.427 898.9 0.72 77.3 D V 800 357.8 3.46 25.9 875 284.0 2.406 640.7 4.76 0.27 76.0 E V 33.5 25.4 41.2 F 975 125.3 0.9303 84.16 0.55 80.2 1075 122.4 0.6457 62.00 85.1 64.1 G 23.1 0.79 H 1250 143.6 1.938 145.9 0.26 70.1 21.3 85.3 2.094 0.24 I 1650 133.9 102.6 14.8 77.5 100.0 total gas age n =9 100.6 0.45 n =4 plateau MSWD = 36.8* steps F-I 84.6 0.49 84.0 40Ar/36Ar - 282 ± 13 * isochron MSWD = 21 .4^ n =8

94.4 85.0 722.0 723.0 726.7 132.88 137.47 133.19 137.15 131.4 135.4 136.9

773.6 3.8 3.8 3.7 7.7 0.38 0.34 0.55 0.47 3.6* 2.5* 2.6*

9N27-1, G4:130, 19.58 mg whole rock, J - 0.00075582 ± 0.10%, NM-130, Lab# = 51782-01 625 7057 0.6491 3337 0.7 A V 7.05 0.79 6.9 766.2 B V 0 700 706.3 0.6236 9.92 0.82 7.7 53.8 750 103.2 C 0.2780 59.62 83.0 14.1 8.97 1.8 D 800 96.64 0.3160 40.86 1.6 87.5 19.9 8.28 E 875 95.96 0.2382 2.1 25.6 42.25 8.05 87.0 F 975 95.41 0.4372 25.37 16.2 1.2 92.2 37.1 G 1075 86.23 0.6087 20.90 29.8 0.84 92.9 58.1 H 1250 95.42 40.57 94.4 2.546 87.7 51.5 0.20 85.57 757.5 7.775 83.4 700.0 7 V 0 7650 7.88 0.30 n =9 141.6 0.79 total gas age plateau MSWD = 153t n =6 steps C-H 122.8 0.83 86.7 40 isochron MSWD = 181t n=l Ar/36Ar - 340 ± 124*

96.6 76.44 113.12 111.84 110.43 116.14 106.10 110.83 764.79 111.1 110.6 108.8

24.3 0.56 0.42 0.39 0.39 0.30 0.23 0.24 0.58 1.0* 3.1* 7.0*

ID

T

Ar

C. S. SIDDOWAY ETAL.

430

Table 3. (continued) ID

T

40

Ar/39Ar

37

Ar/39Ar

36

Ar/39Ar X 10-3r

39 ArK X 10-16 mol

K-Ca

40Ar*

39

Ar

L9N16-2, G5:130, 18.34 mg whole rock, J = 0.000755249 ± 0.10% , NM-130, Lab# = 51783-01 3.433 A V 625 4403 14625 0.985 0.15 1.8 0.5 376.4 B V 9.29 700 213.1 1.635 0.31 47.9 5.5 69.07 C V 750 127.2 1.446 11.8 0.35 11.9 84.0 D 122.2 1.124 17.9 800 48.77 0.45 88.3 21.5 0.6184 E 57.15 23.7 0.82 875 125.9 86.6 34.3 120.4 F 0.4278 38.81 44.7 975 90.5 58.3 1.2 23.72 1075 115.5 0.4663 38.1 1.1 94.0 G 78.8 H 1250 1.390 29.99 31.8 0.37 92.5 116.3 96.0 I 3.185 1650 121.8 50.00 7.50 0.16 88.1 100.0 185.8 total gas age n=9 0.77 steps D-I plateau MSWD = 6.1* n=6 163.7 0.83 88.1 40Ar/36Ar = 281± 11* isochron MSWD = 9.9* n =9

Age (Ma)

±la (Ma)

107.5 134.1 140.24 141.47 142.89 142.73 142.24 141.07 140.92 141.4 142.0 142.6

90.7 1.0 0.48 0.36 0.35 0.28 0.26 0.30 0.50 1.7* 0.7* 1.3*

Notes: Preferred ages are shown in boldface; n, number of heating steps; analyses in italics are excluded from final age calculations; V, analyses excluded from weighted mean age; 0, analyses excluded from inverse isochron age; K-Ca molar ratio calculated from reactor-produced 39ArK and 37ArCa; isotopic ratios corrected for blank, radioactive decay and mass discrimination; not corrected for interfering reactions; individual analyses show analytical error only; plateau and total gas age errors include error in J and irradiation parameters; discrimination = 1.0069 ± 0.00099 a.m.u. t MSWD outside of 95% confidence interval. * 2a error.

1994; Storey et al 1999). The rifted margin is delimited sharply by the shelf-edge break where there is an abrupt increase in ocean depths to >1500 m (Fig. 1). It trends N58E. The N65E to N74E stretching direction is highly oblique to the rifted margin in western Marie Byrd Land and to crustal structures that were active there in middle Cretaceous time. One prevalent structural trend in the Ford Ranges is ESE, as exemplified by the Balchen Glacier Fault (Fig. 2; Siddoway et al 20046), and another is c. N30E (Fig. 2). The regional pattern of dextral and sinistral faults (Fig. 2) and the regional stretching direction from dykes are compatible with transcurrent strain in the Ford Ranges in Early Cretaceous time (Fig. 8). As reported by Luyendyk et al (2003), analysis of brittle faults in the Ford Ranges has revealed E-W-orientated dextral and NE-striking sinistral fault domains separating regions of predominantly normal faulting. Faults probably correspond with high amplitude, steep gradient aerogeophysical lineaments on Edward VII Peninsula (Ferraccioli et al 2002) and parallel to the Ice Stream E (ISE) margin (Luyendyk et al 2003). The ISE structure trends N30E toward the Ross/ Amundsen Province boundary at Land Glacier, identified by Pankhurst et al (1998) as the eastern limit of Palaeozoic, Ross-type crust in Marie Byrd Land. Potentially, the feature is a wrench fault that displaces the province

boundary in Marie Byrd Land, as a rift-stage modification of the Mesozoic convergent margin. The sole structural evidence for c. N-S stretching orthogonal to the western Marie Byrd Land margin comes from mafic dykes emplaced in the Fosdick Mountains dome (Fig. 5). The dyke orientations indicate c. NNE-SSW stretching (Richard et al 1994), in conflict with the broad regional pattern determined from the variety of structural data presented in this paper. The dykes are late discordant features, judging from the sharp planar contacts and lack of foliation (Richard et al 1994). They overprint kilometre-scale folds and penetrative ductile fabrics that do record NE extension; consequently, the Fosdick Mountains dykes are thought to have been emplaced during localized stretching and arching perpendicular to the dome axis as dome rocks were translated up into the brittle realm (Siddoway et al 20046). The single available dyke age of 96.4 ±1.1 Ma (Table 2, sample Mutel-1; Richard et al 1994) is consistent with the youngest dyke ages obtained in the broader Ford Ranges. The broadly coeval 40Ar/39Ar results for Ford Ranges mafic dykes (this study), 40Ar/39Ar mineral ages for migmatite gneisses of the Fosdick Mountains (Richard et al 1994) and 40 Ar/39Ar and apatite fission track ages for shear zone rocks offshore in the Ross Sea (Fitzgerald & Baldwin 1997; Siddoway etal 20040) indicate

431

MARIE BYRD LAND KINEMATIC HISTORY Table 4. Results of 40Ar/39Ar analysis of biotite and muscovite mineral separates

ID

T

40

Ar/39Ar

37

Ar/39Ar

36

Ar/39Ar

39 ArK K-Ca X 10-16 mol

40

39

Ar

Age (Ma)

±lo (Ma)

9D7-4, #21:131, 2.74 mg biotite, J = 0.003906654 ± 0.10%, NM-131, Lab# = 51800-01 0.0447 26.7 0.6 650 85.82 3.05 77.4 A V 0 34.61 5.9 750 29.7 85.6 B V 16.66 0.0245 8.020 20.8 26.7 850 14.36 0.0115 112.8 44.5 97.6 C 1.109 40.4 D 920 14.36 0.0202 74.1 25.2 98.7 0.5398 97.7 48.5 E 1000 14.49 0.0273 1.042 43.8 18.7 14.42 0.9642 57.4 71.6 F 1075 0.0089 125.3 97.9 98.9 84.0 G 1110 14.30 0.0127 0.4755 67.3 40.3 91.2 14.48 97.6 H 1180 0.0438 1.130 38.8 11.7 I 14.36 0.0793 0.6810 15.1 6.4 98.5 94.0 1210 14.30 0.3072 98.4 98.0 1250 0.8008 21.6 1.7 J 99.8 K 14.56 0.6420 1.437 10.2 0.79 97.3 1300 700.0 33.90 70.27 0.879 0.79 39.3 L V 0 7675 2.678 total gas age n = l2 542.2 7.8 93.9 plateau MSWD = 3.4t n =9 steps C-K 509.2 21.6 40Ar/36Ar = 313 ± 29 * isochron MSWD = 4.5t n = 10

63.9 97.86 96.11 97.28 97.17 96.81 97.01 96.93 97.03 96.54 97.20 97.7 96.66 96.92 96.79

2.6 0.34 0.18 0.19 0.19 0.20 0.17 0.20 0.31 0.27 0.41 5.9 0.46* 0.34* 0.52t

L9N15-1, #23:131, 2.26 mg muscovite, J = 0.003896088 ±0.10%, NM-131, Lab# = 51798-01 0.2 98.72 204.2 0.943 3.7 38.5 A V 0 600 0.1650 7.07 7.7 89.0 0.3 57.76 0.4693 79.37 B V 0 650 96.7 0.7 52.64 7.77 5.838 7.6 C V 0 700 0.3218 95.2 7.4 D V 56.00 0.0319 8.969 4.19 76.0 775 56.84 47.0 3.2 E V 0.0124 8.370 9.89 95.6 825 F 54.93 4.522 20.1 6.8 875 0.0015 330.0 97.5 54.27 28.2 98.9 11.8 G 900 0.0008 1.847 662.8 H 0.8559 293.4 99.5 33.7 950 53.81 0.0017 122.8 53.58 0.0014 0.7323 99.1 362.0 99.5 51.4 I 1010 — 54.06 0.7795 99.5 61.9 1050 0.0000 58.9 J K 0.2644 492.4 99.8 1100 53.79 0.0010 141.3 87.1 L 53.92 60.2 202.2 99.7 97.8 1150 0.0025 0.3801 55.45 97.8 98.8 M V 7225 0.0133 4.105 5.30 38.5 97.4 54.69 2.42 7.4 99.2 N V 7350 0.0686 4.798 700.0 7680 58.48 91.9 0 V 0.3175 15.98 4.36 7.6 total gas age n = 15 560.6 17.1 MSWD = 2.1 n-1 steps F-L 530.7 368.7 94.7 plateau 40Ar/36Ar = 320 ± 48 * isochron MSWD = 4.3* n = l2

247.7 298.0 326.5 340.6 346.59 341.95 342.72 341.74 340.63 343.31 342.67 343.22 345.7 340.7 343.7 342.0 342.4 342.4

8.3 3.7 2.2 7.2 0.95 0.71 0.60 0.62 0.73 0.54 0.64 0.63 7.2 7.9 7.5 1.4* 1.0* 1.6t

L9N17-2, #22:131, 1.59 mg muscovite, J = 0.003905434 ±0.10%, NM-131, Lab# = 51799-01 296.4 722.7 0.0287 0.503 0.7 A V 600 77.8 28.6 B V 650 57.96 65.42 7.9 66.6 0.4 0.2635 0.948 50.47 7.67 C V 700 1.046 73.32 0.49 92.3 0.8 55.26 93.7 D V 775 0.2712 72.96 3.50 7.9 7.6 56.06 7.37 96.7 E V 825 0.0322 7.290 75.8 3.4 F 3.684 875 55.51 0.0202 15.7 25.2 98.0 7.3 1.724 27.4 99.0 14.1 G 900 55.11 0.0137 37.1 H 950 54.62 0.6608 80.4 58.2 99.6 0.0088 34.0 99.5 I 1010 54.63 0.0170 0.7568 49.1 29.9 46.1 0.0141 J 1050 54.92 0.9275 23.6 36.1 99.5 52.0 K 54.89 1100 0.0113 0.8607 61.0 99.5 67.0 45.1 L 90.4 85.2 99.5 89.4 1150 54.80 0.0060 0.8668 M 55.81 4.337 26.7 97.7 1225 0.0779 6.5 96.0 54.70 77.82 5.84 0.4885 N V 1350 7.0 93.6 97.5 0.2291 70.2 2.2 92.8 700.0 1720 56.75 O V 73.75 404.1 total gas age n = 15 13.8 n -8 steps F-M 374.2 28.4 plateau MSWD = 0.8 92.6 40Ar/36Ar = 185 ± 63 * isochron MSWD = 62.5* n = l5

237.5 253.5 307.9 330.3 344.6 347.56 348.56 347.57 347.44 348.84 348.78 348.25 348.18 325.6 334.7 346.7 348.2 348.0

76.7 5.2 2.6 7.5 7.0 0.79 0.65 0.51 0.60 0.73 0.62 0.73 0.67 7.5 7.7 1.4* 0.9* 5.7*

X 10-3r

Ar*

432

C. S. SIDDOWAY ETAL.

Table 4. (continued) ID

T

40

Ar/39Ar

37

Ar/39Ar

36

Ar/39Ar X 10-3r

39 ArK K-Ca X 10-16 mol

40

Ar*

39

Ar

Age (Ma)

±lo (Ma)

L9N17-3, #26:131, 1.07 mg muscovite, J = 0.003885327 ± 0.10%, NM-131, Lab# = 51795-01 A J 0 99.45 236.4 2.863 0.563 0.2 600 0.18 30.0 68.72 B V 0 1.914 0.627 650 102.5 0.27 56.1 0.5 C V 0 700 60.19 0.572 3.315 33.33 0.15 84.1 0.9 0.8047 D V 60.98 20.76 1.8 775 0.63 90.0 2.15 E 0.0569 8.093 5.13 825 57.31 9.0 95.8 4.0 F 0.0264 10.2 875 56.63 6.500 19.3 96.6 8.3 G 55.37 0.0111 39.4 925 2.371 46.1 98.7 25.1 H 975 55.27 0.0075 1.547 35.0 67.7 99.1 40.0 55.34 0.0344 I 1.837 11.1 1010 14.8 99.0 44.7 J 1050 55.05 0.0141 1.821 21.1 36.2 99.0 53.7 0.0044 K 54.99 1.866 1150 88.7 116.6 99.0 91.5 L 0 1250 58.17 0.0750 8.29 13.85 6.8 92.9 95.0 0.7182 0.71 90.2 M V 0 1400 58.58 19.51 3.67 96.6 N V 60.66 0.4963 21.79 8.02 89.4 700.0 1650 1.0 total gas age n = l4 234.7 2.0 MSWD = 3.2t n =8 steps E-L plateau 218.9 21.6 93.3 40Ar/36Ar = 289 ± 26,* isochron MSWD = 2.?t n =9

198.2 252.2 324.5 348.9 348.8 347.58 347.32 348.19 348.07 346.42 346.01 344.0 337.7 345.1 345.9 347.3 347.3

77.2 7.0 4.7 2.1 1.2 0.83 0.63 0.58 0.74 0.59 0.58 1.0 7.6 7.7 1.5* 1.2* 1.4*

L9N17-3, #25:131, 3.18 mg biotite, J = 0.003886951 ± 0.10%, NM-131, Lab# - 51796-01 26.20 75.7 15.1 3.4 2.2 650 0.1509 21.53 B 5.410 56.70 0.1181 56.5 4.3 97.2 750 10.6 0.4764 C 850 55.10 0.0082 172.7 62.1 99.7 36.2 D 0.3149 63.2 99.8 920 55.60 0.0071 71.8 45.5 0.7312 0.0256 67.3 79.9 99.6 E V 1000 57.50 54.6 144.1 0.4203 35.7 99.7 F V 56.46 0.0145 75.9 1075 33.4 99.7 G V 1110 55.81 0.0153 0.4660 69.5 86.2 55.69 0.7395 7.6 H V 1180 0.0673 42.5 99.6 92.5 56.14 2.109 17.4 I V 1210 0.1015 5.0 98.9 95.1 0.4494 0.6459 14.2 55.39 99.7 97.2 J V 1250 1.1 1.170 99.4 56.11 0.4601 10.0 1.1 98.7 K V 1300 8.406 8.99 57.48 0.95 95.7 100.0 1675 0.5389 L V 5.2 total gas age n - 12 675.7 12.2 MSWD = 8.0* n =3 steps B-D 292.5 43.3 plateau 40Ar/36Ar = 191 ± 104* isochron MSWD - 43.4* « = 11

733.92 350.08 349.19 352.35 362.55 357.73 353.30 352.18 352.48 350.87 354.17 349.77 348.5 350.8 354.2

0.69 0.59 0.67 0.53 0.62 0.76 0.55 0.59 0.57 0.69 0.81 0.87 1.3* 2.1* 4.9*

A V 0

Notes: Preferred ages are shown in boldface; n, number of heating steps; analyses in italics are excluded from final age calculations; V, analyses excluded from weighted mean age; 0, analyses excluded from inverse isochron age; K-Ca molar ratio calculated from reactor produced 39ArK and 37ArCa; isotopic ratios corrected for blank, radioactive decay and mass discrimination; not corrected for interfering reactions; individual analyses show analytical error only; plateau and total gas age errors include error in J and irradiation parameters; discrimination = 1.0069 ± 0.00099 a.m.u. t MSWD outside of 95% confidence interval. * 2a error.

dynamic events in western Marie Byrd Land at c. 100 Ma to 95 Ma, and the kinematic data now available for the Ford Ranges suggest that they involved transcurrent motions, presumably linked to activity along the Gondwana margin (Fig. 10). The kinematics and timing are consistent with structural and age evidence for dextral shearing at c. 103 Ma in Ellsworth Land to the east (Vaughan & Storey 2000; Vaughan et al 2002). Thus, the onshore record in West Antarctica, based on the western Marie Byrd Land and Ellsworth Land sites, lends support to

the model for dextral oblique convergence along the Cretaceous Gondwana margin, proposed by Sutherland & Hollis (2001). The 40Ar/39Ar results for Ford Ranges mafic dykes are broadly coeval with 40Ar/39Ar mineral ages for migmatite gneisses of the Fosdick Mountains determined by Richard et al (1994). Some samples from Marie Byrd Land sites (Edward VII Peninsula, Lisker & Olesch 1998; and Fosdick Mountains, Richard et al. 1994) yield apatite fission track data that also reflect an event at c. 100 Ma; however, in the region

MARIE BYRD LAND KINEMATIC HISTORY

433

aFig. 7. 40Ar/39Ar age spectra and inverse isochron diagrams for muscovite and biotite mineral separates from samples of wall rock of the Ford Granodiorite suite, (a, b) Sample L9N17-3, muscovite; (c, d) sample L9N17-3, biotite; (e, f) sample L9N17-2, muscovite; (g, h) sample L9N15-1, muscovite. Preferred age is noted within the inverse isochron frames. All errors are reported at 2a.

434

C. S. SIDDOWAY ETAL.

Fig. 8. 40Ar/39Ar results for biotite sample 9D7-4 comes from a mylonitic shear zone cutting Ford Granodiorite. The biotite is recrystallized dynamically within the shear zone and is considered to be syn-tectonic in the zone, (a) Age spectra; (b) inverse isochron diagram. Errors are reported at 2o.

there is fairly clear apatite fission track evidence for rapid transit through the partial annealing zone at c. 80-70 Ma (Richard et al 1994; Lisker & Olesch 1998; Siddoway et al 20040). Thus, there are indications of a technically quiet interval followed by a denudation and rapid cooling event at 71-75 Ma (Richard et al. 1994; Lisker & Olesch 1998; Siddoway et al. 20040). The younger apatite fission track ages seem to

be the sole indication, in Marie Byrd Land, of differential fault movements at the time of rifting between Marie Byrd Land and the Campbell Plateau of New Zealand during Late Cretaceous time. Seafloor spreading commenced around 79 Ma (Stock & Cande 2002). The fairly rapid onset of seafloor spreading and the abrupt rifted margin defined by the -1500 m bathymetric contour in Marie Byrd

MARIE BYRD LAND KINEMATIC HISTORY

435

Fig. 9. Summary diagram of all 40Ar/39Ar ages. From top to bottom, the figure shows first a group of mica ages of >342 Ma; these correspond to host rocks of the Ford Granodiorite suite. The remaining 40Ar/39Ar biotite age in line 5 is the result from biotite involved in mylonitic shear zones at Mt Cooper. The nine groundmass concentrate ages are for dolerite dykes. Those labelled 'RS' yielded initial anomalously young apparent ages at the lowest temperature steps during analysis, followed by monotonically increasing ages at the intermediate to highest temperature steps. Those labelled 'CS' have complex saddle-, hump-, or stair-shaped spectra. For both 'RS' and 'CS' spectra, high K-Ca ratios generally coincide with older apparent ages.

Land and Campbell Plateau (Fig. 1) are possible signs that seafloor spreading commenced upon a pre-existing structure such as a transcurrent fault. The early phase of ENE stretching, recorded by the new kinematic data from western Marie Byrd, is attributed to a transcurrent strain field associated with dextral strike-slip along the Cretaceous Gondwana margin. It is hypothesized that the sharp rifted margin and the principal structural grain on land were established through strike-slip parallel to the Gondwana margin. The stretching direction determined from dyke geometries is kinematically consistent with the interpreted movements upon dextral c. E-W faults and sinistral NW-SE faults, active at the same time, at c. 100 Ma. The orthogonal rifting between the Campbell Plateau of New Zealand and Marie Byrd Land did occur c. 20 Ma later. Thus, modification of the Gondwana margin was accomplished in two stages (cf. Richard et al 1994).

Conclusions Mafic dykes in the Ford Ranges have 40Ar/39Ar ages of 142 Ma to 96 Ma and record crustal stretching orientated ENE-WSW (azimuth c. 070-250), orthogonal to dyke margins. Although some age determinations have large errors due to argon loss caused by alternation of groundmass, the 40Ar/39Ar results nonetheless establish the middle Cretaceous timing of mafic dyke emplacement with some certainty. Diverse types of structural data support the kinematic interpretation of transcurrent strain along and inboard of the Gondwana margin. The recognition of transcurrent structures in Marie Byrd Land preceding break-up of the margin may explain key features of the Marie Byrd Land rifted margin, namely: (1) the rifted margin cuts at a high angle across Ross Sea basins (Lawver & Gahagan 1994); (2) the rifted margin is exceptionally linear (Sutherland 1999); and (3) no evidence is found in the Ford

436

C. S. SIDDOWAY ETAL.

Fig. 10. Tectonic model of a transcurrent setting for the Gondwana margin and western Marie Byrd Land, associated with oblique convergence. The ENE to NE stretching direction interpreted from mafic dykes and other Ford Ranges structures is consistent with dextral transcurrent deformation along the margin. Abbreviations are as follows: CP, Campbell Plateau; MBL, Marie Byrd Land; TI, Thurston Island; EM, Ellsworth Mountains; WM, Whitmore Mountains. Dextral shear indication in Antarctic Peninsula is after Vaughan etal. (1999).

Ranges for distributed deformation or magmatism at the time of break-up (Siddoway et al 20040). The absence of contractional structures and 'rift' sediments is also consistent with a transcurrent strain environment and an oblique orientation of the minimum finite strain axis with respect to major structures. To test this hypothesis, future work will involve integration of geological results with the airborne geophysics dataset for the region and comparison of kinematic data from the Ford Ranges to more distant sites in Marie Byrd Land (e.g. Storey et al 1999), on the Antarctic Peninsula (e.g. Vaughan & Storey 2000; Vaughan et al. 1999), and on the western Ross Sea margin (e.g. Wilson 1995; Salvini et al. 1997). The longitudinal convergence problem' means that such a task will not be straightforward. Direct comparison to structural/geophysical data from other Antarctic regions can be accomplished best when the data are converted to the coordinates of the Antarctic navigational grid (Airforce & Navy 1973) according to the modular conversion: AGN strike = MOD [(latitude position at study site + strike of structure with respect to true north), 360]. The formula is generally applicable, irrespective of longitude. An advantage of this normalization is

that Antarctic navigational grid data can be compared directly with aeromagnetic and other geophysical trends plotted in polar stereographic grid coordinates. The geochemistry of the western Marie Byrd Land dyke array should also be investigated, in order to address the hypothesis that a mantle plume was active in the region and influenced the rifting process in Marie Byrd Land (Weaver et al. 1994; Storey et al. 1999). Funding is from United States National Science Foundation grant OPP-9615282 to C. Siddoway. S. Richard, A. Whitehead, S. Cowdery, M. Roberts & R. Meyer contributed to collection of field data. V. DiVenere, B. Storey, A. Vaughan & I. Dalziel provided helpful reviews. Stereographs were prepared using Stereonet v. 6.2.X [1988-2002] by R. W. Allmendinger, and Faultkin v. 4.1.0 software [2002] by R.W. Allmendinger, R. A. Marrett & T. Cladouhos.

References ADAMS, C.J., SEWARD, D. & WEAVER, S.D. 1995. Geochronology of Cretaceous granites and metasedimentary basement on Edward VII Peninsula, Marie Byrd Land, West Antarctica. Antarctic Science, 7, 265-277. AIR FORCE, DEPARTMENT OF, AND NAVY, DEPARTMENT OF, 1973. Grid Navigation, Air Navigation,

MARIE BYRD LAND KINEMATIC HISTORY Volume AFM 51-40; NAVAIR 00-80V-49: Flight Training. Air Training Command, United States Government Printing Office, Washington, D.C., Chapter 19,1-14. BEST, M.G. 1988. Early Miocene changes in direction of least principal stress, southwestern United States: conflicting inferences from dikes and metamorphic core-detachment fault terranes. Tectonics, 7, 249-259. BRADSHAW, J.D., ANDREW, P.B. & FIELD, B.D. 1983. Swanson Formation and related rocks of Marie Byrd Land and a comparison with the Robertson Bay Group of northern Victoria Land. In: OLIVER, R.L., JAMES, P.R. & JAGO, IB. (eds) Antarctic Earth Science. Cambridge University Press, Cambridge, 274-279. BRADSHAW, ID., PANKHURST, R.I, WEAVER, S.D., STOREY, B.C., MUIR, R.I & IRELAND, T.R. 1997. New Zealand superterranes recognized in Marie Byrd Land and Thurston Island. In: RICCI, C.A. (ed.) The Antarctic Region: Processes and Evolution. Terra Antartica Publications, Siena, 429-436. DALZIEL, I.W.D. & ELLIOT, D.H. 1982. West Antarctica: problem child of Gondwanaland. Tectonics, 1, 3-19. DEINO, A. & POTTS, R. 1990. Single-Crystal 40Ar/39Ar dating of the Olorgesailie Formation, Southern Kenya Rift. Journal of Geophysical Research, 95, 8453-8470. DIVENERE, V., KENT, D.V. & DALZIEL, I.W.D. 1994. MidCretaceous paleomagnetic results from Marie Byrd Land, West Antarctica: A test of post100 Ma relative motion between East and West Antarctica. Journal of Geophysical Research, 99, 15 115-15 139. FERRACCIOLI, F, Bozzo, E. & DAMASKE, D. 2002. Aeromagnetic signatures over western Marie Byrd Land provide insight into magmatic arc basement, mafic magmatism and structure of the eastern Ross Sea Rift flank. Tectonophysics, 347, 139-165. FITZGERALD, P.G. & BALDWIN, S. 1997. Detachment fault model for evolution of the Ross Embayment. In: RICCI, C.A. (ed.) The Antarctic Region: Processes and Evolution. Terra Antartica Publications, Siena, 555-564. POLAND, K.A., FLEMING, T.H., HEIMANN, A. & ELLIOT, D.H. 1993. Potassium-Argon dating of finegrained basalts with massive Ar loss: application of the 40Ar/39Ar technique to plagioclase and glass from the Kirkpatrick Basalt, Antarctica. Chemical Geology, 107, 173-190. IRELAND, T.R., BRADSHAW, ID., MUIR, R.I, WEAVER, S.D. & ADAMS, C.I 1994. Zircon age distributions in granites, greywackes, and gneisses from the southwest Pacific Gondwana Region. U.S. Geological Survey Circular, 1107,151. LAWVER, L.A. & GAHAGAN, L.M. 1994. Constraints on timing of extension in the Ross Sea Region. Terra Antartica, 1, 545-552. LISKER, F. & OLESCH, M. 1998. Cooling and denudation history of western Marie Byrd Land, Antarctica, based on apatite fission-tracks. In: VAN DEN

437

HAUTE, P. & DE CORTE, F. (eds) Advances in Fission-Track Geochronology. Kluwer Academic Publishers, The Netherlands, 225-240. LUYENDYK, B.P., RICHARD, S.M., SMITH, C.H. & KIMBROUGH, D.L. 1992. Geological and geophysical investigations in the northern Ford Ranges, Marie Byrd Land, West Antarctica. In: YOSHIDA, Y., KAMINUMA, K. & SHIRAISHI, K. (eds) Recent Progress in Antarctic Earth Science. Terra Publications, Tokyo, Japan, 279-288. LUYENDYK, B.P, CISOWSKI, S., SMITH, C.H., RICHARD, S.M. & KIMBROUGH, D.L. 1996. Paleomagnetic study of the northern Ford Ranges, western Marie Byrd Land, West Antarctica: a middle Cretaceous pole, and motion between West and East Antarctica? Tectonics, 15,122-141. LUYENDYK, B.P, SORLIEN, C.C., WILSON, D., BARTER, L., ELY, G, SIDDOWAY, C.S. & ZELLMER, K. 2001. Structural and tectonic evolution of the Ross Sea rift in the Cape Colbeck region, Eastern Ross Sea, Antarctica. Tectonics, 20, 6, 933-958. LUYENDYK, B.P, WILSON, D. & SIDDOWAY, C.S. 2003. The eastern margin of the Ross Sea Rift in western Marie Byrd Land: Crustal structure and tectonic development. Geochemistry, Geophysics, Geosystems (G3), 4 (10), 1090, doi:10.1029/ 2002GC000462. MAHON, K.I. 1996. The New 'York' regression: Application of an improved statistical method to geochemistry. International Geology Review, 38, 293-303. MUIR, R.I, IRELAND, T.R., WEAVER, S.D. & BRADSHAW, ID. 1994. Ion microprobe U-Pb zircon geochronology of granitic magmatism in the Western Province of the South Island, New Zealand. Chemical Geology, 113,171-189. PANKHURST, R.I, WEAVER, S.D., BRADSHAW, ID., STOREY, B.C. & IRELAND, T.R. 1998. Geochronology and geochemistry of pre-Jurassic superterranes in Marie Byrd Land, Antarctica. Journal of Geophysical Research, B2, 103, 2529-2547. RICHARD, S.M., SMITH, C.H., KIMBROUGH, D.L., FITZGERALD, P.G. & LUYENDYK, B.P. 1994. Cooling history of the northern Ford Ranges, Marie Byrd Land, West Antarctica. Tectonics, 13, 837-857. SALVINI, F, BRANCOLINI, G, BUSETTI, M., STORTI, F, MAZZARINI, F. & COREN, F. 1997. Cenozoic geodynamics of the Ross Sea Region, Antarctica: Crustal extension, intraplate strike-slip faulting and tectonic inheritance. Journal of Geophysical Research, 102, 24 669-24 696. SAMSON, S.D. & ALEXANDER, E.G., JR. 1987. Calibration of the interlaboratory 40Ar/39Ar dating standard, Mmhb-1. Chemical Geology, 66, 27-34. SIDDOWAY, C, BALDWIN, S., FITZGERALD, P.G, FANNING, C.M., & LUYENDYK, B.P. 20040. Ross Sea mylonites and the timing of intracontinental extension within the West Antarctic rift system. Geology, 32, 1, 57-60. SIDDOWAY, C., RICHARD, S., FANNING, C.M. & LUYENDYK, B.P. 2004/?. Origin and emplacement mechanisms for a middle Cretaceous gneiss dome, Fosdick Mountains, West Antarctica. In: WHITNEY, D., TEYSSIER, C.T. & SIDDOWAY, C.

438

C. S. SIDDOWAY ETAL.

(eds) Gneiss domes and orogeny. Geological Society of America Special Paper 380, 267-294. STEIGER, R.H. & JAGER, E. 1977. Subcommission on geochronology: Convention on the use of decay constants in geo- and cosmochronology. Earth and Planetary Science Letters, 36, 359-362. STOCK, J.M. & CANDE, S. C. 2002. Tectonic history of Antarctic seafloor in the Australia-New ZealandSouth Pacific sector: Implications for Antarctic continental tectonics. In: GAMBLE, J., SKINNER, D., HENRYS, S. & LYNCH, R. (eds) Proceedings of 8th International Symposium on Antarctic Earth Sciences. Royal Society of New Zealand Bulletin, 35, Wellington, New Zealand, 251-259. STOREY, B.C., DALZIEL, I.W.D., GARRETT, S.W., GRUNOW, A.M., PANKHURST, RJ. & VENNUM, W.R. 1988. West Antarctica in Gondwanaland: crustal blocks, reconstruction and break-up processes. Tectonophysics, 155, 381-390. STOREY, B.C., LEAT, P.T., WEAVER, S.D., PANKHURST, ID. & KELLEY, S. 1999. Mantle plumes and Antarctica-New Zealand rifting: evidence from mid-Cretaceous mafic dykes. Journal of the Geological Society, London, 156, 659-671. SUTHERLAND, R. 1999. Basement geology and tectonic development of the greater New Zealand region: An interpretation from regional magnetic data. Tectonophysics, 308, 341-362. SUTHERLAND, R. & HOLLIS, C. 2001, Cretaceous demise of the Moa plate and strike-slip motion at the Gondwana margin. Geology, 29, 279-282. TAYLOR, IR. 1982. An Introduction to Error Analysis: The Study of Uncertainties in Physical Measurements. University Science Books, Mill Valley, California, 270. TSUNAKAWA, H. 1983. Simple two-dimensional model of propagation of magma-filled cracks. Journal of Volcanology and Geothermal Research, 16, 335-342. VAUGHAN, A.P.M. & STOREY, B.C. 2000. The eastern

Palmer Land shear zone: a new terrane accretion model for the Mesozoic development of the Antarctic Peninsula. Journal of the Geological Society, London, 157,1243-1256. VAUGHAN, A.P.M., MILLAR, I.L. & THISTLEWOOD, L. 1999. The Auriga Nunataks shear zone: Mesozoic transfer faulting and arc deformation in northwest Palmer Land, Antarctica. Tectonics, 18, 911-928. VAUGHAN, A.P.M., PANKHURST, RJ. & FANNING, C.M. 2002. A mid-Cretaceous age for the Palmer Land event, Antarctic Peninsula: implications for terrane accretion timing and Gondwana palaeolatitudes. Journal of the Geological Society, London, 159, 113-116. WEAVER, S.D., BRADSHAW, ID. & ADAMS, CJ. 1991. Granitoids of the Ford Ranges, Marie Byrd Land, Antarctica. In: THOMSON, M.R.A., CRAME, I A. & THOMSON, IW. (eds) Geological Evolution of Antarctica. Cambridge University Press, Cambridge, 345-351. WEAVER, S.D., ADAMS, C.I, PANKHURST, RJ. & GIBSON, I.L. 1992. Granites of Edward VII Peninsula, Marie Byrd Land: anorogenic magmatism related to Antarctic-New Zealand rifting. Transactions of the Royal Society of Edinburgh, 83, 281-290. WEAVER, S.D., STOREY, B.C., PANKHURST, RJ., MUKASA, S.B., DlVENERE, VI & BRADSHAW, ID.

1994. Antarctica-New Zealand rifting and Marie Byrd Land lithospheric magmatism linked to ridge subduction and mantle plume activity. Geology, 22, 811-814. WILSON, T 1995. Cenozoic transtension along the Transantarctic Mountains-West Antarctic rift boundary, southern Victoria Land, Antarctica. Tectonics, 14, 531-545. YORK, D. 1969. Least squares fitting of a straight line with correlated errors. Earth and Planetary Science Letters, 5, 320-324.

Index Page numbers in italic refer to figures, page numbers in bold refer to tables.

accretion 1,6, 76-77 APGR 103,105 Delamerian Cambrian forearc 69,70 Famatina Complex 252-253 Jurassic, southwestern Gonwana 233-234 Mesozoic, episodicity 143-166 New England Orogen 61-62 Proterozoic-Palaeozoic, South America 307 Tabberabberan cycle 51, 54, 74 Victoria Land 280, 284, 285, 286-287 Adavale Basin 53, 57, 58 Adelaide Fold-Thrust Belt 27 Adelaide Rift Complex 25, 27, 32,33, 37, 67, 68 Albemarle arc 104-105,104 Altiplano-Puna, Cenozoic, Central Andes 257,258, 265-269 Amazonia block 105,306, 310, 311, 320, 329,331, 343 Ambato, metasedimentary provenance study 397, 398, 399 Amundsen Province, Antarctica 205-206 Anaiwan terrane 52,59, 62, 64 Anakie Inlier 26, 29, 30, 31, 38, 44, 47,59, 60, 62, 67 Andean Orogen 257, 265-269 Andes, Central Cenozoic Altiplano-Puna 257, 265-269 Palaeozoic Orogen 257-265, 268-269 Colombian 332-343 andesite Delamerian 39, 41, 69 New England Orogen 60, 63 anoxia, Mesozoic 151-152,159 Antarctica 10-11, 359, 360-361 archaeocyathan limestone 351 Gondwana-Pacific margin 113,114,115,117 metasediments 117-118 Nd-Sr isotope studies 119,122-123,131 Gondwana-Pangaea margin, Triassic-Jurassic deformation 145,147-148 lithospheric magmatism 362-374, 370 lithospheric mantle domains 359-374 mafic magmas 359-360,361 Marie Byrd Land, kinematic history 417-436 mid-Cretaceous deformation 152-154, 417 seismic structure 297,298, 301 source rocks 205-206 tectonism 360,361 terranes 10-11, 360-361 Victoria Land terranes 275-287,276,278 Antofalla terrane 308, 316-318,317, 322 Appalachian peri-Gondwanan realm (APGR) 97-108, 98 reconstruction 97,106-108 apparent polar wander path 160, 309, 310, 314, 315 Ar/Ar ratios Antarctic lithospheric magmatism 362, 363, 367 Ford Ranges mafic dykes 421-436

Proterozoic inliers, Colombian Andes 332, 337, 339 arc continental margin 60 collision 69, 77 island Famatina 245 intra-oceanic 38, 39, 43, 51, 54, 60, 65, 46 magmatic 230-231,276, 277, 279, 280, 282, 284, 286 Famatina-Eastern Puna 311-312, 321 Garzon Complex 341, 342, 343 Median Tectonic Zone 183,184 migration, Jurassic, southwestern Gondwana 232-233 Neoproterozoic-Cambrian 99,104-108 Puna-Altiplano 258 Ross Orogen 276, 283, 284, 286-287 see also terrane, magmatic arc archaeocyaths 347, 349, 350-356,350,354-355 Arequipa Massif 316, 322 Argentina Cenozoic Altiplano-Puna 265-269 Famatina complex 241-253,242 Nd-Sr isotope studies 123-124,131-132 Palaeozoic Orogen 257-265, 268-269 Proterozoic-Palaeozoic evolution 305-322 Puncoviscana complex 381-410,382 Argentine Precordillera see Cuyania terrane Arthur River Complex 187,189,190 Asia, mid-Cretaceous deformation 155-156 Aspiring terrane 116,120, 726,129-130 asthenosphere 359-360, 363 Auburn arch 58, 60-61 Australia 9 Gondwana margin 113,114,115,119,121,126 mid-Cretaceous deformation 156 Tasmanides 25, 67 Australides 2,3 Avalon Zone 97, 98, 98,103,105,107-108 Avoca Fault 26, 27, 40 back-arc system Benambran cycle 47, 70-71 Las Termas belt 245-250,252-253 geodynamics of closure 250 late Devonian, Drummond Basin 58,59, 60 models 75 Tabberabberan cycle 51, 52-54 back-docking, island-arc 252-253 Bancannia Trough 25,28, 32, 39 Barnard Province 31, 38, 44 basalt Mesozoic LIPs 150-151,157-158 mid-ocean ridge (MORB) 28, 39, 45, 46, 53, 63, 64, 71,316,359, 363-364 ocean island (OIB) 359, 363-364

INDEX

440

see also magma, mafic basin filling, Triassic 65-66 basin formation back-arc Benambran cycle 47, 70-71 Bowen Basin 205 Las Termas belt 245, 247-250 forearc, New England Orogen 27, 60,61, 69 foreland Bowen-Gunnedah-Sydney 65, 69 Melbourne Trough 53 Puncoviscana 408-410 Torrens Hinge Zone 43 late Devonian 58, 60, 61 rift, early Permian 63 Tabberabbera cycle 57, 52-53, 73 basin inversion, Lachlan Orogen 55-56, 58 Batholith of Central Patagonia 218,219 batholiths 54, 56, 74,155,158,183 Patagonia 218,219-221 Beenleigh Block 58, 61 Bega terrane 42, 71, 72, 73, 77 Benambran cycle 41,45-47,49, 51,70-71,73 Benambran Orogeny 28, 41, 49 Bendigo terrane 42,71, 72,74 Bendigo Zone 29, 40, 49 Bindi Orogeny 28, 56 biotite 338-340 blueschist Benambran cycle 47, 69 mid-Cretaceous deformation 155,156 Bounty Trough 180,181 Bowen Basin 26, 29, 30-31,59, 63, 64, 65-66,205, see also Hunter-Bowen supercycle Bowers terrane 276, 277,278, 279-280,281, 282,

283-287,360,567

Broken River subprovince 31, 38, 53, 54, 56-57 Brook Street terrane 115,116, 111, 120,726,128-129, 191-192 Bucaramanga Gneiss 339-340, 342 Buller terrane 115,776,119,125,726,127,184,185, 187,188,199 Cambrian Delamerian convergent phase 38-39, 41, 44 Gondwanan palaeogeography 103 Piedmont zone, tectonothermal event 102 Puncoviscana Complex 381-410,382,383 Takaka terrane 187 Campbell Plateau 180,181 Campbell Magnetic Anomaly System 203 Cape River Metamorphics 38,44 Cape York Peninsula Batholith 53,54 Caples terrane 115,776,117,120,726,127,129,131, 193,195 carbonate deposition 53 Caribbean, mid-Cretaceous deformation 154-155 Carolina Zone 97-99,103,104-106,108 Cenozoic, Altiplano-Puna of Central Andes 257, 265-269 Central America Gondwana-Pangaea margin, Triassic-Jurassic deformation 149 mid-Cretaceous deformation 154-155 Central Atlantic Magmatic Province 150,152

Central Gondwana block 310, 311 Challenger Plateau 180,181 Charlotte terrane 99,104,104 Charters Towers Metamorphics 38 Chatham Rise 180,181 metasedimentary rocks 118 Chile, Nd-Sr isotope studies 122,123,132 Chilenia terrane 316, 322 Chon Aike Province 219, 231,232 clasts archaeocyathan limestone 347-356 conglomerate 196-197,199, 201, 204,205 coal measures, Sydney Basin 65 Cobb Igneous Complex 188 collision 1, 31 Benambran cycle 49 Delamerian cycle 37, 43-44, 69, 77 post-collisional phase 44-45 Tabberabberan cycle 54-57 Triassic, New England Orogen 66 see also deformation conglomerates, clast geochronology 196-197,199, 204, 205 Congo-Sao Francisco craton 307,309, 310 Connors arch 60-61 conodont identification 46, 51, 286 convergence 31 Benambran cycle 45-47, 48,49 Delamerian cycle 38-39, 41, 43 Hunter-Bowen supercycle 58, 60-62 New England Orogen 65-66 Tabberabberan cycle 51-54,73 Cordillera 2,3 Cordillera Oriental, metasedimentary provenance study 397, 398 Cretaceous, mid Gondwana-Pangaea margin deformation 143, 146-147,152-159 magmatism 157 Curnamona craton 25, 27, 28, 32, 39, 68 Cuyania terrane 306, 308, 311, 312-316, 319-320 Laurentian origin 313-314, 321-322 cycles, tectonic, Tasmanides 31-66 D'Aguilar Blocks 58, 61, 62 Darran Suite plutons 184,187,189,190,797,199 deformation 31 Cambrian, Delamerian cycle 43, 69 Carboniferous, New England Orogen 62,77 Cretaceous, palaeo-Pacific Ocean 143,146-147, 152-159,160 Devonian, Lachlan Orogen 28, 75 Benambran 73 Kanimblan 57-58 Tabberabberan 55-56 Triassic, New England Orogen 66, 77 Triassic-Jurassic, Gondwana-Pangaea margin 143-152,144-145,159-160 Victoria Land terranes 281-282,282, 286 see also collision Delamerian cycle 32-45 collisional phase 37, 43-44, 69 convergent phase 37, 38-39, 41, 43 post-collisional phase 37, 44-45, 69 rift phase 32-33,33, 37-38, 67-69

INDEX Delamerian Orogen 26, 27-28,34,35, 43-44,103,188 Benambran cycle 47, 51 Kanimblan cycle 57, 58 stratigraphy 34 Tabberabberan cycle 51, 53 deposition glacial, Delamerian cycle 32,33 turbidite, Delamerian cycle 38 Deseado Monzonite Suite 218, 219 Devil River Volcanics Group 188 Diamantina River Lineament 25,26, 49, 68 diamictite, glacigenic, Palaeozoic 348 Dibulla Gneiss 339, 342 Dimboola Igneous Complex 37, 39 Djungati terrane 51,52,59 docking see accretion Dronning Maud Land, magmatism 368-369, 370, 373 Drummond Basin 55, 58,58, 60, 66 Dun Mountain-Maitai terrane 115,116, 111, 120, 726, 129,131,192 Dundas Trough 43, 44 Dwyka Tillite, archaeocyathan limestone 351 dykes Gairdner Dyke Swarm 32, 68 Las Termas Belt 245 mafic, Antarctica 363, 366-368,371 Ford Ranges 417-436,419 earthquakes, seismic data 295, 301 Eastern Province, New Zealand 181,182,185,186, 191-195 metasediments 114,115,116,117,181,182 Nd-Sr isotope studies 118,120-121, 726,127-131 eclogite Charlotte Arc 104,105 Peel-Manning fault system 41 Electric Granite 184,199 Ellsworth Mountains magmatism 366, 370 metasediments 131,135 episodicity 143 erosion, crustal 294 Eurasia, Gondwana-Pangaea margin, Triassic-Jurassic deformation 149-150 extension Cretaceous, West Antarctica 417 Jurassic-Cretaceous, New Zealand 201-202 Permian, New England Orogen 62-65 post-compressional, mid-Cretaceous 158-159 see also rifting extinction, mass, Mesozoic 152,159 Falkland Islands 347,348,349 archaeocyathan limestone 347-356 magmatism 361,362, 366-367, 370 Famatina complex 241-253,242,244, 245,246 Famatina-Eastern Puna magmatic arc 306, 308, 311-312, 321 metasedimentary provenance study 398, 399 faults extensional 62,74 New England Orogen 62, 66 Tabberabberan collision phase 56 terrane boundaries 4 Fiordland, New Zealand 185,189,190

441

Fitzroy Tillite Formation 348-353 Fleurieu structural arc 37, 43, 69 fold-thrust belt Delamerian Orogen 43 New England Orogen 30, 62, 66 folds, orogen-scale see orocline formation Ford Ranges mafic dykes 417-436, 419 Fosdick Mountains gneiss dome 420, 424, 427, 430 fossils Cambrian 41,104,105,108, 279 archaeocyaths 347-356 Ordovician 281 Silurian-Devonian 51 gabbro, Benambran cycle 47 Gairdner Dyke Swarm 32, 68 Gander Zone 98 garnet 250,253, 332, 338-339 Garzon Complex 338-339, 341-343 Gaussberg, Antarctica, magmatism 369, 370, 371 Gawler Craton 25, 43, 69 geochemistry Mesozoic granites, Patagonia 224-228, 225 Puncoviscana complex 384,386-396,400-407, 409-410 geochronology Antarctic lithospheric magmatism 362-372 conglomerate clasts 196-197,199, 204 Grenvillian inliers 330-332,333, 334-337, 338-341, 343 Mesozoic granite, Patagonia 222, 223-224 Palaeozoic Central Andes 261-262,263-264 Puncoviscana complex 384 south Pacific margin metasedimentary rocks 113-136 Victoria Land terranes 279, 281, 284-286 geomagnetism, Mesozoic 151,159 glaciation, Gondwana, Late Carboniferous 352 gneiss, Proterozoic inliers, Colombian Andes 333, 334-337, 338-340 Gondwana APGR 98,102,103 apparent polar wander path 309-310,309 assembly, Proterozoic 307-322, 361 breakup, Mesozoic 231 Cambrian 103 Jurassic, magmatic arc migration 232-233 Large Igneous Province 150,152 Laurentia collision 11, 313-314 Pacific margin 114,179,180 isotopic signatures 113,118-136 Mesozoic terrane accretion 143-166 mid-Cretaceous deformation 146-147,152-159, 753,160, 417 plate boundary 38, 51, 66, 68, 70, 73, 75, 77 back-docking v. terrane accretion 252-253 magmatism 217-235 Ross Orogen 275 tectonic reconstruction 418 palaeogeographic reconstruction 198,199-208,232, 233-234,349, 352,362, 418, 436 Pangaea margin, Triassic-Jurassic, deformation 143-152,144-145 Panthalassan margin 181 palaeogeographical reconstruction 199-208

442

Proterozoic-Palaeozoic evolution 307-322,372,327 Puncoviscana Complex 381-410 Governor Fault zone 39, 40, 44, 57 granite Carboniferous, New England Orogen 60-62 Delamerian cycle 39, 43 Kanimblan cycle 57, 58 Mesozoic, Patagonia 219-222 geochemistry 224, 225,226, 227-228 geochronology 222, 223-224 isotope data 228-230 Permian-Triassic, New England Orogen 65 Tabberabberan cycle 51,53, 54, 56, 73,74 Granite Harbour magmatic arc 282-285 granitoids Famatina Complex 245,246, 248,249 northern Victoria Land terrane 281, 282, 284 Palaeozoic Orogen, Central Andes 261 granodiorite, Ford Ranges mafic dykes 418-436 gravity anomaly 318 Greenland Group 115,187,199 Grenvillian orogen 329,361 inliers Colombian Andes 329, 330, 343 Queensland 38 Guapoton-Moncagua Gneiss 338-339, 342 Gunnedah Basin 26, 30-31,59, 69, 64, 65, 66, 70 Gympie terrane 58, 65, 66 Haast Schist 195 Heathcote Fault zone 39, 40, 44 Hikurangi Plateau 180,181 Hodgkinson Formation 53-54 Hodgkinson subprovince 31,49, 53-54,56 Hunter-Bowen Orogeny 66 Hunter-Bowen supercycle 52,55, 58, 60-66, 63 Carboniferous convergence 60-62 late Devonian convergence 58, 60 hydrocarbons, Mesozoic 152,159 lapetus Ocean 703, 314, 319,327, 330, 343 Ida Fault, seismic structure 299,300 inliers, Proterozoic Andes 329-344,337 Queensland 31-38 see also Anakie Inlier island arcs see arc, island isotope data conglomerate clasts 197,199, 201 Cuyanian terrane 313 Mesozoic granite, Patagonia 228-230 Pacific margin metasediments 119-124,124-136 Palaeozoic central Andes 261-262,263-264 Japan, mid-Cretaceous deformation 155-156 Jojoncito Gneiss 340, 342 Jurassic, Early magmatism, Patagonia 219-221, 230-233 palaeogeography, Patagonia 233-234 see also Triassic-Jurassic K-Ar dating 125 Kalahari craton 307, 310, 311, 319 Kanimblan cycle 54, 57-58

INDEX deformation 57-58 post-collisional phase 58 rifting phase 57 Kanimblan Orogeny 28 Kanmantoo Fold Belt 27 Kanmantoo Trough 25, 27,28, 32,37-38,43, 69 Karoo mantle plume 231, 234 Koonenberry Belt 32,33,39,41,43, 45, 51, 53, 57 Lachlan Fold Belt, metasediments 115,119,121,125, 726,127 Lachlan Orogen 26, 27, 28-29, 30, 38, 39,40, 69,75-78 Benambran cycle 40, 41 deformation 28,44, 55-56 Delamerian cycle 38, 39, 40, 44 Devonian 55-56 Kanimblan cycle 40, 57-58 Lambie facies 55, 57 Ordovician, tectonism 29, 49 Ordovician terranes 42 stratigraphy 34,35 Tabberabberan granites 40, 51, 54, 56, 73, 74 Lachlan supercycle 45-47, 49, 51-58,70-76 Lambie facies 55, 57-58 lamproite 360, 368, 370-371, 372 lamprophyre 360,364,366-367,369 large igneous provinces (LIPs) 157-158 Las Margaritas Gneiss 338-339, 341 Las Termas belt 243 back-arc basin 245-250 Laurentia 11, 319-320,327 Grenville belt 329-330, 343 origin of Cuyania terrane 312-315, 321-322 separation 67, 68, 310, 313 Laurentia-APGR interaction 98-99,101,102,103, 105-108 limestone archaeocyathan 349, 350-352 Benambran cycle 47 lithosphere, break-up continental 164-165 oceanic 162-163 deep structure 294-302 evolution 294 magmatism 362-374 sub-continental, composition 360 Lolworth-Ravenswood Block 31, 38,47, 53, 54 Lord Howe Rise 117,180,181 Macquarie Arc 29, 42, 46, 49, 70, 71, 74 magma 7 asthenospheric 359-360, 363,371-372 mafic, Antarctica 359-374 sub-continental lithospheric 360 magmatism Cenozoic Andean 266-269 Early Jurassic, Patagonia 217-235,218 hot spot 160-165 Late Triassic-Early Jurassic 150-151 lithospheric, Antarctica 362-374,370 mid-Cretaceous 152,153,157-158 Neoproterozoic-Palaeozoic, APGR 104-105 Palaeozoic Central Andes 257,260, 261,268-269 post-Cambrian, New Zealand 188-189 see also superplumes; arc, magmatic

INDEX magnetic anomaly 203, 301 Maitai terrane 192 see also Dun Mountain-Maitai terrane mantle 74, 293, 301-302 lithospheric, Antarctica 359-374 Marie Byrd Land Ford Ranges mafic dykes 419, 420^36 geology 417-418, 418, 419, 420 magmatism 158, 365, 370, 373 metasediments 117,119,131 Martinsville Intrusive Suite 101,102 Median Batholith 116, 111, 179,183,184 Median Tectonic Zone (MTZ) 116, 111, 179,181-184, 187 Meguma Zone 98,108 melange zone 7, 8,100 Melbourne Trough 39, 52-53 Melbourne Zone 40, 49, 56 Mesozoic anoxia 151-152 geomagnetism 151 Gondwana-proto-Pacific plate rift 66 terrane accretion, Gondwana-Pacific margin 143, 144-145 volcano-sedimentary units 190-191 metamorphism APGR 100 Cenozoic, Andean 266-269 Palaeozoic, Central Andes 258-261,260 Proterozoic inliers, Colombian Andes 329-344 Puncoviscana complex 381 schist 195, 282-284, 286 superimposed, Famatina complex 243-245 Victoria Land terranes 275-278, 282, 283, 284 metasediment Las Termas belt 250 northern Victoria Land 277-278, 281 Pampean Ranges 243 Pampia 311 Puncoviscana, provenance studies 381-410 South Pacific margin 113-136 Moho discontinuity 294, 295-297 monzonite APGR 102,103 Deseado Massif 219 Subcordilleran belt 221 Moyston Fault 26, 27, 39, 40, 44 Mt Wellington Fault zone 39, 69, 70 Murihiku terrane 115,116, 111, 120,126,129,131, 192, 204 mylonite zones 243, 250, 252, 282,306, 311, 421, 423 Narooma terrane 42, 44, 47 Nd isotope signatures Antarctic lithospheric magmatism 362-372 Gondwana Pacific margin 113,118-136,184 igneous clasts 197,199 Mesozoic granites, Patagonia 228-229 Palaeozoic Central Andes 261,262,263 Proterozoic inliers, Colombian Andes 332, 336, 338-341 Nebine ridge 29-30 Nelson, New Zealand 183,184,188-190 Neoproterozoic-early Palaeozoic metasediments, Puncoviscana 381-410

443

palaeogeography, APGR 97,104 proto-Pacific ocean 38 rifting 67 New England Batholith 65 New England Fold Belt 204-205 New England Orogen 25,26, 28, 30,58, 69, 78 Benambran cycle 41, 47, 51, 71 Carboniferous convergence 60-62 Carboniferous deformation 62, 77 Delamerian cycle 38, 41 Early Permian extension 62-65 Hunter-Bowen cycle 52 late Devonian convergence 59, 60 metasediment Nd-Sr isotope data 121,726 orocline formation 64 Permian-Triassic convergence 65-66 stratigraphy 36 strike-slip model 71, 72, 73, 75 Tabberabberan cycle 51,52, 54-55,73 Triassic collision 66 New Zealand 9,179 continental crust 181 East Gondwana margin 113-117 Nd and Sr isotope studies 118-136,119-121 Gondwana-Pangaea margin, deformation 150 Gondwana-Panthalassan margin 181 Median Tectonic Zone 179,181-184,187 mid-Cretaceous deformation 156-157 palaeogeographic reconstruction 199-208 southwest Pacific margin 180 tectonostratigraphy 186,187-195 Norfolk ridge 180,181 North America Late Triassic-early Jurassic deformation 149 mid-Cretaceous deformation 155 North Patagonian Massif 218, 219 North Queensland Orogen 25,26, 28, 31, 38, 48,11, 78 Benambran cycle 41, 47 Lambie facies 57 late Devonian arc 60 stratigraphy 36 Tabberabberan cycle 51, 53, 56 granites 54, 73 north Victoria Land see Victoria Land, Antarctica obduction 43 ophiolite 143,154, 310 oceans, Mesozoic 151-152,159-160 oil see hydrocarbons Olepoloko Fault System 26, 27 ophiolites 7, 38,160,161,162,192 Cambrian, New England Orogen 71 Neoproterozoic, Pampia 310 Ordovician, Chilenia 316 Ordovician black shale 71 Lachlan Orogen 42 tectonism 49 orocline formation, New England Orogen 59, 64 orogenesis 1-2 accretionary 77 orthogneiss 190 Otago Schist 150 Ouachita Embayment 314, 315

444

INDEX

Pacific Ocean see palaeo-Pacific Ocean; Gondwana, Pacific margin Pahau terrane 115,116,121,126,130,131,194 conglomerate clasts 196-197 palaeo-Pacific Ocean 103 mid-Cretaceous, deformation 143,152-160 palaeo-Tethys Ocean 148,149,159,160 palaeogeography, reconstruction APGR 97-108 Patagonia 233-234, 318-319 Proterozoic-Palaeozoic, South America 265, 305, 319-322 palaeomagnetism, Proterozoic-Palaeozoic APGR 108 Cuyania terrane 314-315 Famatina-Eastern Puna magmatic arc 311-312 Patagonia 319,320 Puna 316-318 Rio de la Plata craton 306, 307,308-310,309 South America 305, 307, 308 Palaeozoic Gondwanan active margin 102-103 Orogen of Central Andes 257-265,267,268-269 Ross Orogen 275-286 Palmer Land event 143,152 Palmerville Fault System 24, 25,26, 31, 68 Pampean complex 244 Pampean cycle 241, 243-245 Pampean Orogen 102-103,103, 243, 252 Pampia 306, 310-311, 319, 320 Pan-African Orogeny 361, 368, 372 Pangaea 148,159-160 Gondwana margin, Triassic-Jurassic deformation 143-152,144-145,148,159-160 Panthalassan Ocean 148,160 Panthalassan plate margin 181 Papua New Guinea, Tasmanides 25 Patagonia evolution 318-319 palaeomagnetism 308, 319 Subcordilleran belt, Mesozoic magmatism 217-235, 218 Patagonian Batholith 219, 220, 231 Peel-Manning Fault System 30, 41, 47, 51,59, 60, 62, 63, 64, 69, 71 pellites 383, 384 Peninsula Orogeny 143,145 Peninsular Ranges batholith 155,158 peri-Gondwana, Appalachian (APGR) 97-108 Permian, New England Orogen 62-65 Permian-Triassic boundary 66 petrography, Puncoviscana 385, 399-400,407 Piedmont Zone 98, 99,106 Cambrian tectonothermal event 102 central Piedmont shear zone 99,105 plate tectonics boundary forces 157 models 75-76 superplume events 160-164 plumes 158 mantle 217, 231, 359, 373 see also superplumes plutonism APGR 99,104 Mesozoic 151,153,158,181,183,183,190, 217-235

New Zealand 183-184,187-188 Palaeozoic, Antarctica 284 polar wander path, apparent 160, 309, 310, 314, 315 post-collision Delamerian cycle 44-45, 69 Kanimblan cycle 58 Prince Charles Mountains, magmatism 369,370, 371 prism, accretionary 1, 54, 62, 71, 75 Proterozoic inliers Colombian Andes 329-344 Queensland 31 palaeogeographical reconstruction 305 palaeomagnetism 305-322 proto-Pacific plate boundary magmatism 38, 217-235 rollback 70, 73, 76, 77 tectonic models 75-76 provenance studies, Puncoviscana 399-410 Puna plateau 257-269,258 metasediments 123-124 provenance study 397, 398 Puncoviscana complex 258, 381-410,382 foreland basin model 408-410 geochemistry 384, 386-396,400-407,409-410 geology 383-384 petrography 385, 399-400, 407 Queensland New England Orogen 30,36 Tasmanides 25,24, 36 Thomson Orogen 29 see also North Queensland Orogen radiolaria 51 Rakaia terrane 115,116,121,126,130,131,194 conglomerate clasts 196, 203-205 recycling 202-203 source 205-206 Rangitata Orogeny 143,150,156 rare earth elements 362, 401, 402-403, 404, 405, 407-408, 410 Ravens wood Batholith 54 Rayner Province, Antarctica, seismic structure 297, 298 Rb-Sr isochron studies Famatina Complex 250, 252 Gondwana Pacific margin 118,125 Palaeozoic Central Andes 261,262 reconstruction, palaeogeographical APGR 97 Gondwana 198,200,307-322 Gondwana-Laurentia 68 New Zealand 198,199,200,201-208 Patagonia 233-234 Proterozoic-Palaeozoic, South America 265, 305-322 Rodinia 67-68 residence age, crustal 132-135 rhyolite magmatism 150,151,219, 231 rifting 31 APGR 103-104,105 Delamerian cycle 32-33,33, 37-38, 67-69 Hunter-Bowen cycle 63-64 Kanimblan cycle 57

INDEX Mesozoic 66 Tabberabberan cycle 52, 73-74 see also extension Rio de la Plata craton 306, 307, 308, 309-310 Robertson Bay terrane 276, 277,278, 280-287, 360, 361 Rodinia 67-68, 76, 305, 307,309, 310, 329, 343 rollback, proto-Pacific plate boundary 69,70,73,76, 77,157 Ross Orogen 38, 46,103,188, 275-277,276, 284-285, 286-287, 360,36,?, 368, 373 Ross Province, Antarctica 206, 365, 418 rotation, late Carboniferous, New England Orogen 62,64 Russia, Mid-Cretaceous deformation 155 sandstone, turbiditic, Benambran cycle 45 San Rafael, remagnetization 314 Schirmacher Oasis, magmatism 369, 370, 372 schist Haast schist 195 north Victoria Land terranes 282-284, 286 Otago schist 150 seismic velocity studies earthquake data 295 lithospheric characteristics of terranes 294 seismic receiver functions 295, 301 Separation Point Suite 181,189,190 serpentinite, New England Orogen 62-63 Shackleton Range, magmatism 367-368,370 shale, black Benambran cycle 45-46, 49, 71 Victoria 115 shear, TIPA shear zone 243, 250,251, 252 Shoalwater terrane 59, 61,117 shoshonite 29, 46, 57,151 Sierra de Famatina 241, 245,246 volcano-sedimentation 247,248 Sierras Pampeanas 245, 312, 384 Sm-Nd isotope studies 124,132-133, 332, 336, 338-340 Smith River allochthon 97, 98, 99,100-104,106 accretion to Laurentia 103 palaeogeographical model 106,107 rifting from Gondwana 103-104 source 102-103 South America 10 Gondwana Pacific margin metasediments 118 Nd-Sr isotope studies 122-123,131-132 Gondwana-Pangaea margin, Triassic-Jurassic deformation 149 mid-Cretaceous deformation 154 palaeogeographical evolution 319-322 palaeomagnetism 305-322 South Shetland Islands, 122,132 Southern Cross terrane 297,300 Southwest terrane 296-297,298 Sr isotope signatures Antarctic lithospheric magmatism 362-372 Gondwana-Pacific margin 113,118-136,184 igneous clasts 197,199 Mesozoic granites, Patagonia 228-229 Palaeozoic Central Andes 261,263 staurolite ages 102,103

445

Stavely Volcanic Complex 39, 44, 45, 69 Stawell Zone 26, 27, 28, 39, 40, 44 Stewart Island 185,187,189,190,191 strike-slip model 75, 76 APGR 105 New England Orogen 62, 64, 71, 72,73 sub-continental lithospheric mantle 293, 301-302, 360 Subcordilleran belt, Patagonia, Mesozoic magmatism 217,218, 219-221,220, 230-231 subduction Cambrian, Delamerian cycle 38,41, 69 Cambrian-Ordovician Benambran cycle 47, 49, 73 Famatina Complex 252 late Devonian 60 Mesozoic 195 New England Orogen 61-62 North Queensland Orogen 48 northern Victoria Land terrane 283-285, 286 plate tectonic models 75-76 Silurian-Devonian, Tabberabberan cycle 51,73 Sunsas belt 329, 343 superplumes 143,160,161,162-165,163,164 sutures 6 Lachlan Orogen 28 Ross Orogen 278,283,285 Suwanee terrane 98,108 Sydney Basin 26, 30-31,58,63, 64, 65-66 Tabberabbera zone 40 Tabberabberan cycle 50, 51-57,73-75 collisional phase 54-57 convergent phase 51-54 granites 51, 54, 56, 73, 74 Tabberabberan Orogeny 28, 56, 74, 77 Takaka terrane 115,116,119,125, 726,127,183,184, 185,187-188 Tamworth Trough 51, 60, 61, 62, 64 Tapley Hill Shale 68 Tasman Line 24, 25, 68-69 Tasmania 32 Delamerian cycle 44, 69 forearc collision 43, 69 post-collision 44 rifting from Australia 67 seismic structure 297,299 Tabberabberan cycle 53, 56 Tasmanides 9, 23,24,26, 76-78 boundaries 24, 25 rifting 67-69, 76 subdivisions 27-31 tectonic cycles 31-66, 67 tectonic models 75-76 tectonism Antarctica 360,367, 418 APGR 100,102,103,106 Jurassic, Gondwana-Patagonia 233-234 New Zealand 181-206,183,184 Ordovician Lachlan Orogen 49 North Queensland Orogen 48 Proterozoic, Colombian Andes 341-344 TIPA shear zone 250,257 see also plate tectonics Terra-Australis orogen 2,3

INDEX

446

terrane allochthonous see terrane, exotic analysis 8-9,11 authochthonous 284-285 collages 8 definition 2, 4,277, 293 exotic 179,275,277, 284, 285,287,293, 313 palaeogeography 100-108 magmatic arc 7, 99 northern Victoria Land 275-287,276,278 Ordovician, Lachlan Orogen 42 processes 6 rock types 6-8 sedimentary 7-8 seismic structure 296-302 sub-crustal lithospheric mantle 293 terrane boundaries 4 northern Victoria Land 282-284, 285 seismic structure 294, 297,299-300, 301 Tethys Ocean see palaeo-Tethys Ocean Texas terrane 59, 62, 64 Thomson Orogen 25,26,28,29-30,78 Benambran cycle 47,51 Delamerian cycle 38,39,41,44 Kanimblan cycle 57, 58 Tabberabberan cycle 53 thrust systems mid-Cretaceous 155,156 Wilson terrane 282 Thurston Island, magmatism 158, 365, 370 tillite see diamictite; Fitzroy Tillite Formation; Dwyka Tillite Tinogasta-Pituil-Antinaco shear zone see TIPA shear zone TIPA shear zone 243, 250,257, 252 tomography, seismic 294 Torlesse composite terrane 115, 776,117,118,130, 179,191,193-194, 205 Torrens Hinge Zone 25, 43, 67, 68 Transantarctic Mountains 131, 206,276, 284, 351, 352,

360,367,373,475

Triassic, collision, New England Orogen 66 Triassic-Jurassic Gondwana-Pangaea margin deformation 143, 144-145,148 Americas 148-149 Antarctica 145,147-148 Eurasia 149-150 New Zealand 150 magmatism, Patagonia 219 turbidites 6-7 Benambran cycle 45-46,49, 71 Delamerian cycle 38 post-collisional 45 Robertson Bay terrane 280-281 Tabberabberan cycle 53-54,74 Tutoku Complex 183,184,185,187,189-191,199 Tyennan Orogen see Delamerian Orogen U-Pb geochronology 125,187,188 APGR 102,103

Palaeozoic Central Andes 261,262,264 Proterozoic inliers, Colombian Andes 332,333, 334-335,338-340,342,343 Subcordilleran belt 222,223-224,224, 230 uplift, crustal 294 velocity, seismic see seismic velocity studies Vergel Granulites 338-339, 341 Vestfjella, Antarctica, magmatism 368-369,370 Victoria, Australia Delamerian Orogen 26, 27,53, 69 Lachlan Orogen 28, 29 seismic structure 297,299 Victoria Land, Antarctica magmatism 365-366, 370 terranes 275-287,276,275 volcanics, calc-alkaline 7, 39, 45, 61, 69,184,190 volcanism 7 APGR 98-99 Benambran cycle 71 Bowers terrane 279 Central Andes 266 Delamerian cycle 32-33, 37, 39,41, 44,45,67, 69, 70 Early Permian 63, 64, 65 Famatina Complex 245,246, 247-250 Kanimblan cycle 57 Lachlan Orogen 29 New England Orogen 60,61, 63, 64, 65, 66 Tabberabbera cycle 51-54, 55, 73 see also magmatism Wagga Basin 70-71 Waipapa terrane 115,776,117,121,726,130,131, 194-195 Wandilla terrane 59, 61,117 West African block 310, 311 West Tamar Fault Zone 26, 27 Western Province, New Zealand 181,182,184,185, 186,187-191, 206 magmatism 187-190 metasediments 113, 774,115 Nd-Sr isotope studies 118,119-120,125, 726,127 Tutoko Complex 183,184, 755,189 Wilson Cycle 1-2,143 Wilson 'terrane' 277-279,275,282, 283-287, 360,367 Yarrol Fault System 30,59,61 Yarrol Trough 51,59, 60, 61 zircon dating 11,125 Cambrian turbidites 38,188 Las Termas belt 248 Ordovician turbidites 46,187 Proterozoic inliers Colombian Andes 332,333,334-335, 338-340, 342, 343 Queensland 38 Puncoviscana complex 384 Rakaia sandstones 204-205, 206 Tabberabberan granite 54