Geodynamic evolution of East Antarctica: a key to the East-West Gondwana connection

Geological correlations of East Antarctica with adjoining continents have been puzzling geologists ever since the conce...

Author: M. Satish-Kumar | Geological Society of London

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Geological correlations of East Antarctica with adjoining continents have been puzzling geologists ever since the concept of a Gondwana supercontinent surfaced. Despite the paucity of outcrops because of ice cover, difficulty of access and extreme weather, the past 50 years of Japanese Antarctic Research Expeditions (JARE) has successfully revealed vital elements of the geology of East Antarctica. This volume presents reviews and new research from localities across East Antarctica, especially from Dronning Maud Land to Enderby Land, where the geological record preserves a history that spans the Archaean and Proterozoic. The reviews include extensive bibliographies of results obtained by geologists who participated in the JARE. Comprehensive geological, petrological and geochemical studies, form a platform for future research on the formation and dispersion of Rodinia in the Mesoproterozoic and subsequent assembly of Gondwana in the Neoproterozoic to Early Palaeozoic.

Geodynamic Evolution of East Antarctica: A Key to the East –West Gondwana Connection

The Geological Society of London Books Editorial Committee Chief Editor

BOB PANKHURST (UK) Society Books Editors

JOHN GREGORY (UK) JIM GRIFFITHS (UK) JOHN HOWE (UK) PHIL LEAT (UK) NICK ROBINS (UK) JONATHAN TURNER (UK) Society Books Advisors

MIKE BROWN (USA) ERIC BUFFETAUT (FRANCE ) JONATHAN CRAIG (ITALY ) RETO GIERE´ (GERMANY ) TOM MC CANN (GERMANY ) DOUG STEAD (CANADA ) RANDELL STEPHENSON (NETHERLANDS )

Geological Society books refereeing procedures The Society makes every effort to ensure that the scientific and production quality of its books matches that of its journals. Since 1997, all book proposals have been refereed by specialist reviewers as well as by the Society’s Books Editorial Committee. If the referees identify weaknesses in the proposal, these must be addressed before the proposal is accepted. Once the book is accepted, the Society Book Editors ensure that the volume editors follow strict guidelines on refereeing and quality control. We insist that individual papers can only be accepted after satisfactory review by two independent referees. The questions on the review forms are similar to those for Journal of the Geological Society. The referees’ forms and comments must be available to the Society’s Book Editors on request. Although many of the books result from meetings, the editors are expected to commission papers that were not presented at the meeting to ensure that the book provides a balanced coverage of the subject. Being accepted for presentation at the meeting does not guarantee inclusion in the book. More information about submitting a proposal and producing a book for the Society can be found on its web site: www.geolsoc.org.uk.

It is recommended that reference to all or part of this book should be made in one of the following ways: SATISH -KUMAR , M., MOTOYOSHI , Y., OSANAI , Y., HIROI , Y. & SHIRAISHI , K. (eds) 2008. Geodynamic Evolution of East Antarctica: A Key to the East– West Gondwana Connection. Geological Society, London, Special Publications, 308. TOYOSHIMA , T., OSANAI , Y. & NOGI , Y. 2008. Macroscopic geological structures of the Napier and Rayner Complexes, East Antarctica. In: SATISH -KUMAR , M., MOTOYOSHI , Y., OSANAI , Y., HIROI , Y. & SHIRAISHI , K. (eds) Geodynamic Evolution of East Antarctica: A Key to the East –West Gondwana Connection. Geological Society, London, Special Publications, 308, 139–146.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 308

Geodynamic Evolution of East Antarctica: A Key to the East – West Gondwana Connection

EDITED BY

M. SATISH-KUMAR Shizuoka University, Japan

Y. MOTOYOSHI National Institute of Polar Research, Japan

Y. OSANAI Kyushu University, Japan

Y. HIROI Chiba University, Japan and

K. SHIRAISHI National Institute of Polar Research, Japan

2008 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of over 9000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society’s fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society’s international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists’ Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies’ publications at a discount. The Society’s online bookshop (accessible from www.geolsoc.org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London W1J 0BG: Tel. þ44 (0)20 7434 9944; Fax þ44 (0)20 7439 8975; E-mail: [email protected]. For information about the Society’s meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society’s Corporate Affiliates Scheme, write to [email protected] Published by The Geological Society from: The Geological Society Publishing House, Unit 7, Brassmill Enterprise Centre, Brassmill Lane, Bath BA1 3JN, UK (Orders: Tel. þ44 (0)1225 445046, Fax þ44 (0)1225 442836) Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. # The Geological Society of London 2008. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of the Copyright Licensing Agency, 90 Tottenham Court Road, London W1P 9HE. Users registered with the Copyright Clearance Center, 27 Congress Street, Salem, MA 01970, USA: the item-fee code for this publication is 0305-8719/08/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 978-1-86239-268-7 Typeset by Techset Composition Ltd., Salisbury, UK Printed by MPG Books Ltd., Bodmin, UK Distributors North America For trade and institutional orders: The Geological Society, c/o AIDC, 82 Winter Sport Lane, Williston, VT 05495, USA Orders: Tel. þ1 800-972-9892 Fax þ1 802-864-7626 E-mail [email protected] For individual and corporate orders: AAPG Bookstore, PO Box 979, Tulsa, OK 74101-0979, USA Orders: Tel. þ1 918-584-2555 Fax þ1 918-560-2652 E-mail [email protected] Website http://bookstore.aapg.org India Affiliated East-West Press Private Ltd, Marketing Division, G-1/16 Ansari Road, Darya Ganj, New Delhi 110 002, India Orders: Tel. þ91 11 2327-9113/2326-4180 Fax þ91 11 2326-0538 E-mail [email protected]

Preface International collaboration to study geophysical phenomena in the polar region dates back to 1882, which was designated as the first International Polar Year (IPY). The second IPY was organized 50 years later in 1932. Thereafter, in view of the progress in observation potentiality and scientific demands, the third IPY was arranged 25 years after the second IPY, and it was renamed as the International Geophysical Year (IGY) and ran during the period from July 1957 to December 1958. Japan commenced its scientific activity in the Antarctic in 1957 as one of the 64 participating nations in IGY, and established ‘Syowa Station’ on the Ongul Islands, Lu¨tzow-Holm Bay in East Antarctica by the 1st Japanese Antarctic Research Expedition (JARE-1) on 29 January 1957. Since then, Japan has been undertaking observation and research projects on various disciplines of natural sciences extending over 50 years in the Antarctic. Geological survey around Syowa Station, a region where no human being had ever set foot, started from the beginning in 1957 by JARE-1. During the past 50 years, a total number of nearly 100 geologists has joined in JARE to conduct geological surveys in the Lu¨tzow-Holm Bay region, Prince Olav Coast, the Yamato Mountains, the Belgica Mountains, the Sør Rondane Mountains, and Enderby Land, in East Antarctica. Despite a paucity of outcrops because of ice cover, difficulty of access and extreme weather, the JARE has successfully revealed vital elements of the geology of East Antarctica, and this led us to the attempt to clarify the origin and evolution of continents and their dynamics in the Earth’s history. Geological correlation of East Antarctica with adjoining continents has been a major topic of discussion among geologists. However, in the early 1990s, JARE succeeded in revolutionizing our understanding of East Antarctic geodynamics by the discovery of a Cambrian mobile belt in Lu¨tzow-Holm Bay. A couple of decades after, through this Special Publication, we attempt to compile reviews and new research from localities across East Antarctica, especially from Dronning Maud Land and Enderby Land. Reviews provide extensive bibliographies of results obtained by geologists who participated in the JARE geological activities. Comprehensive geological, petrological

and geochemical studies will potentially form a platform for future research on the geodynamics of amalgamation of Gondwana in the Neoproterozoic to Early Palaeozoic. In addition, the coincidence of Gondwana amalgamation with major global-scale climatic, environmental and biological changes in Late Neoproterozoic to Early Palaeozoic times implies a close connection between largescale tectonic events and global change, which needs to be confirmed in the future. The success of geological studies of JARE is indebted to the dedicated logistic support of the crews of icebreakers Soya, Fuji and Shirase. The journey is continuing with the commissioning of a new vessel. The last 50 years of dedicated group work of the Japanese Antarctic Research Expeditions in Antarctica is commendable, in this extreme environment of hazardous weather and travel conditions. We hope this Special Publication will enthuse the young generation and be a new starting point for the next 50 years research activity on the Antarctic geosciences.

This publication is a part of the Japanese contribution to IPY 2007–2008. M. SATISH -KUMAR , Y. MOTOYOSHI , Y. OSANAI , Y. HIROI & K. SHIRAISHI

Acknowledgements The editors would like to thank the following scientists around the world who gave generously of their time and expertise in reviewing the manuscripts submitted to this Special Publication. Makoto Arima Sotaro Baba Wilfried Bauer Mike Brown Chris Carson Somnath Dasgupta Prelevic Dejan Christoph Dobmeier Dennis Eberl David Ellis Mike Flowerdew Reinhardt Fuck Geoff Grantham Ed Grew Richard Hanson Joerg Hermann Tomokazu Hokada Julie Hollis Jan-Marteen Huizenga Kiyoshi Ito Masahiro Ishikawa Hideo Ishizuka Joachim Jacobs Simon Johnson Hiroo Kagami Hiroyuki Kagi Masaki Kanao

Ken-ichi Kano Dave Kelsey Tony Kemp Kare Kullerud Axel Liebscher Victor Melezhik Akira Miyake Anand Mohan Tomoaki Morishita Hans Mueller Takashi Nakajima Atsushi Okamoto Masaaki Owada Bob Pankhurst Konstantin Podlesskii H. M. Rajesh V. Ravikant K. Sajeev M. Santosh Rajesh Shrivastava Robert Stern Fabrizio Storti Bob Thomas Nobutaka Tsuchiya Carlos Villaseca Yue Zhao

Finally, we are grateful to Phil Leat, the editor-in-charge of the Special Publication, and Angharad Hills of the Geological Society Publishing House for continuous support and encouragement throughout the editing process.

Contents Preface

vii

Acknowledgements

viii

SATISH -KUMAR , M., HOKADA , T., KAWAKAMI , T. & DUNKLEY , D. J. Geosciences research in East Antarctica (08E–608E): present status and future perspectives

1

SHIRAISHI , K., DUNKLEY , D. J., HOKADA , T., FANNING , C. M., KAGAMI , H. & HAMAMOTO , T. Geochronological constraints on the Late Proterozoic to Cambrian crustal evolution of eastern Dronning Maud Land, East Antarctica: a synthesis of SHRIMP U –Pb age and Nd model age data

21

JACOBS , J., BINGEN , B., THOMAS , R. J., BAUER , W., WINGATE , M. T. D. & FEITIO , P. Early Palaeozoic orogenic collapse and voluminous late-tectonic magmatism in Dronning Maud Land and Mozambique: insights into the partially delaminated orogenic root of the East African– Antarctic Orogen?

69

GRANTHAM , G. H., MACEY , P. H., INGRAM , B. A., ROBERTS , M. P., ARMSTRONG , R. A., HOKADA , T., SHIRAISHI , K., JACKSON , C., BISNATH , A. & MANHICA , V. Terrane correlation between Antarctica, Mozambique and Sri Lanka; comparisons of geochronology, lithology, structure and metamorphism and possible implications for the geology of southern Africa and Antarctica

91

ISHIZUKA , H. An overview of geological studies of JARE in the Napier Complex, Enderby Land, East Antarctica

121

TOYOSHIMA , T., OSANAI , Y. & NOGI , Y. Macroscopic geological structures of the Napier and Rayner Complexes, East Antarctica

139

SATISH -KUMAR , M., MIYAMOTO , T., HERMANN , J., KAGAMI , H., OSANAI , Y. & MOTOYOSHI , Y. Pre-metamorphic carbon, oxygen and strontium isotope signature of high-grade marbles from the Lu¨tzow-Holm Complex, East Antarctica: apparent age constraints of carbonate deposition

147

MIYAMOTO , T., SATISH -KUMAR , M., DUNKLEY , D. J., OSANAI , Y., YOSHIMURA , Y., MOTOYOSHI , Y. & CARSON , C. J. Post-peak (,530 Ma) thermal history of Lu¨tzow-Holm Complex, East Antarctica, based on Rb – Sr and Sm–Nd mineral chronology

165

ISHIKAWA , M., SHINGAI , E. & ARIMA , M. Elastic properties of high-grade metamorphosed igneous rocks from Enderby Land and eastern Dronning Maud Land, Antarctica: evidence for biotite-bearing mafic lower crust

183

SUZUKI , S., ISHIZUKA , H. & KAGAMI , H. Early to middle Proterozoic dykes in the Mt. Riiser-Larsen area of the Napier Complex, East Antarctica: tectonic implications as deduced from geochemical studies

195

SUDA , Y., KAWANO , Y., YAXLEY , G., KORENAGA , H. & HIROI , Y. Magmatic evolution and tectonic setting of metabasites from Lu¨tzow-Holm Complex, East Antarctica

211

OWADA , M., BABA , S., OSANAI , Y. & KAGAMI , H. Geochemistry of post-kinematic mafic dykes from central to eastern Dronning Maud Land, East Antarctica: evidence for a Pan-African suture in Dronning Maud Land

235

HOKADA , T., MOTOYOSHI , Y., SUZUKI , S., ISHIKAWA , M. & ISHIZUKA , H. Geodynamic evolution of Mt. Riiser-Larsen, Napier Complex, East Antarctica, with reference to the UHT mineral associations and their reaction relations

253

CARSON , C. J. & AGUE , J. J. Early Palaeozoic metasomatism of the Archaean Napier Complex, East Antarctica

283

TSUNOGAE , T., SANTOSH , M., DUBESSY , J., OSANAI , Y., OWADA , M., HOKADA , T. & TOYOSHIMA , T. Carbonic fluids in ultrahigh-temperature metamorphism: evidence from Raman spectroscopic study of fluid inclusions in granulites from the Napier Complex, East Antarctica

317

vi

CONTENTS

HIROI , Y., MOTOYOSHI , Y., ISHIKAWA , N., HOKADA , T. & SHIRAISHI , K. Origin of xenocrystic garnet and kyanite in clinopyroxene –hornblende-bearing adakitic meta-tonalites from Cape Hinode, Prince Olav Coast, East Antarctica

333

KAWAKAMI , T., GREW , E. S., MOTOYOSHI , Y., SHEARER , C. K., IKEDA , T., BURGER , P. V. & KUSACHI , I. Kornerupine sensu stricto associated with mafic and ultramafic rocks in the Lu¨tzow-Holm Complex at Akami Point, East Antarctica: what is the source of boron?

351

YOSHIMURA , Y., MOTOYOSHI , Y. & MIYAMOTO , T. Sapphirine þ quartz association in garnet: implication for ultrahigh-temperature metamorphism in Rundva˚gshetta, Lu¨tzow-Holm Complex, East Antarctica

377

GOTO , S. & IKEDA , T. Crystal size distribution of garnet in quartzo-feldspathic gneisses from the Lu¨tzow-Holm Complex at Skallen, East Antarctica

391

BABA , S., OWADA , M. & SHIRAISHI , K. Contrasting metamorphic P–T path between Schirmacher Hills and Mu¨hlig-Hofmannfjella, central Dronning Maud Land, East Antarctica

401

KAWASAKI , T. & OSANAI , Y. Empirical thermometer of TiO2 in quartz for ultrahightemperature granulites of East Antarctica

419

SATO , K., MIYAMOTO , T. & KAWASAKI , T. Fe2þ –Mg partitioning experiments between orthopyroxene and spinel using ultrahigh-temperature granulite from the Napier Complex, East Antarctica

431

Index

449

Geosciences research in East Antarctica (088E – 6088E): present status and future perspectives M. SATISH-KUMAR1, T. HOKADA2, T. KAWAKAMI3 & DANIEL J. DUNKLEY2 1

Institute of Geosciences, Shizuoka University, Oya 836, Suruga-ku, Shizuoka 422-8529, Japan (e-mail: [email protected]) 2

3

National Institute of Polar Research, Kaga, Itabashi-ku, Tokyo 173-8515, Japan

Department of Geology and Mineralogy, Kyoto University, Kitashirakawa-oiwake-cho, Sakyo-ku, Kyoto 606-8502, Japan Abstract: In both palaeoenvironmental and palaeogeographical studies, Antarctica plays a unique role in our understanding of the history of the Earth. It has maintained a unique geographical position at the South Pole for long periods. As the only unpopulated continent, the absence of political barriers or short-term economic interests has allowed international collaborative science to flourish. Although 98% of its area is covered by ice, the coastal Antarctic region is one of the wellstudied regions in the world. The integrity and success of geological studies lies in the fact that exposed outcrops are well preserved in the low-latitude climate. The continuing programme of the Japanese Antarctic Research Expedition focuses on the geology of East Antarctica, especially in the Dronning Maud Land and Enderby Land regions. Enderby Land preserves some of the oldest Archaean rocks on Earth, and the Mesoproterozoic to Palaeozoic history of Dronning Maud Land is extremely important in understanding the formation and dispersion of Rodinia and subsequent assembly of Gondwana. The geological features in this region have great significance in defining the temporal and spatial extension of orogenic belts formed by the collision of proto-continents. Present understanding of the evolution of East Antarctica in terms of global tectonics allows us to visualize how continents have evolved through time and space, and how far back in time the present-day plate-tectonic regime may have operated. Although several fundamental research problems still need to be resolved, the future direction of geoscience research in Antarctica will focus on how the formation and evolution of continents and supercontinents have affected the Earth’s environment, a question that has been addressed only in recent years.

The formation and evolution of continents has always been an intriguing topic in Earth Science studies. The complexity of continental evolution largely results from the protracted and recurring nature of geological processes that have taken place in the continental lithosphere. Decoding billions of years of complex history recorded in the continental crust is a daunting task. However, geologists have made great progress in understanding the processes involved in continental formation and their evolution through time. The Antarctic continental lithosphere is an important crustal fragment that provides us with an abundance of information on the formation of continents, and the temporal and spatial relationships involved in the assembly and dispersion of supercontinents. The significance of Antarctica lies not only in its unique geographical position, whereby it has gained due importance in palaeoenvironmental studies, but also in its geological stability since incorporation in the supercontinent Gondwana at the beginning of the Phanerozoic Era. This is primarily because the Antarctic lithospheric plate has been surrounded by

mid-ocean ridges since the Mesozoic (Fig. 1 inset), with the exception of the Antarctic Peninsula, which is its only active convergent plate margin with transform faults dividing the Antarctic Plate and Scotia Plate. This means that the Antarctic Plate is currently expanding relative to the surrounding plates. This is a feature that is at variance with most other lithospheric plates, and makes the Antarctic continent exceptionally stable, and isolated from all regional tectonic events during the Mesozoic and Cenozoic. Consequently, the older geological history of East Antarctica can be considered as one of the least overprinted records of crustal evolution in the Earth’s history, preserved in a natural ‘cold storage’. Its geological history records the formation of early Archaean protocontinents, and continues throughout the Proterozoic, until the amalgamation of East and West Gondwana at the beginning of the Palaeozoic. Therefore, the geological record in East Antarctica is an invaluable record of the origin and evolution of continents and supercontinents, and for understanding the secular changes in metamorphic conditions in orogenic belts (Brown 2007).

From: SATISH -K UMAR , M., MOTOYOSHI , Y., OSANAI , Y., HIROI , Y. & SHIRAISHI , K. (eds) Geodynamic Evolution of East Antarctica: A Key to the East– West Gondwana Connection. Geological Society, London, Special Publications, 308, 1 –20. DOI: 10.1144/SP308.1 0305-8719/08/$15.00 # The Geological Society of London 2008.

2 M. SATISH-KUMAR ET AL. Fig. 1. Index map of geographical regions and localities in East Antarctica corresponding to the contributions in this Special Publication. Inset shows a topographic map of Antarctica and surrounding oceans. Red indicates topographically elevated places; blue indicates ocean floor. (Data source: Department of Commerce, National Oceanic and Atmospheric Administration, National Geophysical Data Center, 2006, 2-minute Gridded Global Relief Data (ETOPO2v2), http://www.ngdc.noaa.gov/mgg/fliers/ 06mgg01.html).

GEOSCIENCES RESEARCH IN EAST ANTARCTICA

It is beyond the scope of this book to update the reader with the voluminous literature that has been produced in the past few decades on the geology of East Antarctica. However, we make an attempt to integrate the results of some recent studies from the eastern region of the Antarctic continent, where the Japanese Antarctic Research Expedition (JARE) has, over the past 50 years, conducted extensive investigations. We introduce the general geology of the region and summarize what is known to date, and in the process introduce the contributions in this volume. The contributions in the volume are related to the outcrops that are situated between 08E and 608E in Dronning Maud Land and Enderby Land of East Antarctica (Fig. 1). In addition, this paper also attempts to lay down ‘a vision for future’, based on the current status of geological knowledge.

East Antarctica: an integral part of Gondwana The challenge of developing tectonic scenarios for the formation of the ice-covered Antarctic continent is uniquely difficult; no other continent presents such a blank sheet on which geological terranes can be drawn by inference only. Virtually all understanding of the geological architecture is drawn from intensive studies of coastal outcrops and mountain ranges near the continental margins. A full 1808 arc of coastline, encompassing East Antarctica, provides an array of outcrops that almost exclusively share a Precambrian origin. This reflects the intracontinental nature of the East Antarctic coast in the supercontinent Gondwana, after its formation at the end of the Proterozoic. The stability of the continent throughout the Phanerozoic has also led to the concept of an East Antarctic Shield, one of the large areas of cratonized crust on Earth. The ‘shield’ concept also influenced tectonic interpretations of coastal geology before the formation of Gondwana. It was recognized that most localities in East Antarctica are represented by areas of high tectonic activity, dominated by moderate- to high-temperature metamorphic belts, shear zones, and regions of Proterozoic crustal growth, and that Archaean granite– greenstone and metamorphic terranes are mostly restricted to small discrete localities. This led to the development of tectonic models of a ‘cratonized’ East Antarctic Shield with extensive mobile belts, such as the c. 1 Ga Circum-Antarctic Mobile Belt of Yoshida (1992) and the Wegener –Mawson Mobile Belt of Kamenev (1991). These models implied the existence of a coherent Antarctic continent that was amalgamated during the formation of the supercontinent Rodinia at 1.3– 0.9 Ga

3

(Hoffman 1991). However, subsequent years have seen a steady increase in the volume and detail of tectonic and geochronological research from all areas of East Antarctica that has shown a more complex story of the diverse origins of various sectors of the East Antarctic margin, challenging the ‘shield’ paradigm. Late Mesoproterozoic metamorphic terranes located along the Antarctic coast at 308W–358E (the 1100– 1000 Ma Maud Belt), 458E– 708E (the 1000–900 Ma Rayner Complex) and 1008E–1208E (the 1300– 1100 Ma Wilkes Province), were found not only to differ subtly in age, but also to be separated by areas of c. 600–500 Ma moderate- to high-temperature metamorphism and tectonism at Lu¨tzow-Holm Bay (408E) and Prydz Bay (708E; Fitzsimons 2000). Thus, instead of representing a continuous marginal mobile belt, each of the Mesoproterozoic metamorphic terranes could be correlated with discrete mobile belts in South Africa (Namaqua–Natal Belt), India (Eastern Ghats) and South Australia (Albany– Fraser Orogen). Furthermore, it was recognized that a large section of the Maud Belt was reworked by late Neoproterozoic metamorphism and deformation that could be correlated with the extensive East African Orogen, produced by the amalgamation of East and West Gondwana (Jacobs et al. 2003a). Recognition of unrelated pre-Rodinian cratons in East Antarctica was also achieved, with the correlation of the Mawson continent and the Gawler Craton in South Australia (Fanning et al. 1996), and the geochronological characterization of Archaean terranes south of a c. 550 Ma suture zone in the southern Prince Charles Mountains adjacent to Prydz Bay (Boger et al. 2001; Mikhalsky et al. 2001, 2006; Phillips et al. 2006). New studies (e.g. Kelsey et al. 2008) continue to develop the latest paradigm of the assembly of East Antarctica from disparate continental bodies during the late Neoproterozoic formation of Gondwana. In particular, the complexity of crustal development in the sector between 08E and 708E, namely Dronning Maud Land, Enderby Land, Kemp Land and Mac Robertson Land, is the focus of recent and current research. Shiraishi et al. attempt to synthesize a large amount of new geochronological data obtained from eastern Dronning Maud Land, and discuss the variations in age distributions between the lithological units. Magmatic and metamorphic events between 1200 and 500 Ma are identified from zircon geochronology in different regions, providing insights into the formation and assembly of crustal fragments, Neoproterozoic sedimentation, and late Neoproterozoic to Cambrian episodes of metamorphism and magmatism. Shiraishi et al. further consider the geodynamic evolution of eastern Dronning Maud Land on the basis of published and new Nd model ages, which

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M. SATISH-KUMAR ET AL.

Fig. 2. Neoproterozoic Gondwana showing the cratonic regions and surrounding mobile belts. Simplified after Gray et al. (2008) and modified taking into consideration the Lawyer et al. (1998) tight-fit Gondwana configuration. SL, Sri Lanka; MD, Madagascar; WA, western Australia; EA, eastern Australia; SA, southern Australia; SF, Saõ Franscisco; RP, Rio de la Plata.

indicate juvenile extraction of Mesoproterozoic crust in the Sør Rondane Mountains, in contrast to the mixed Archaean and Proterozoic derivation of continental crust in the Lu¨tzow-Holm Complex. In a Gondwanan perspective, these results will shift the attention of geodynamic modelling to eastern Dronning Maud Land, to clarify the complex orogenic processes involved in the amalgamation of East and West Gondwana (Fig. 2). The significance of voluminous plutonic activity in Dronning Maud Land and northern Mozambique is discussed by Jacobs et al., who consider lateral southward extrusion and extensional collapse as the preferred tectonic scenario, potentially as a result of crustal delamination. The correlation of temporal variations with distinct shifts in geochemical affinities of magmatic regimes form the basis of a delamination model in association with orogenic collapse and escape tectonics, as proposed recently by Jacobs & Thomas (2004). The model remains to be tested in relation to Gondwanan geodynamics. An intriguing conundrum of correlation of terranes in Gondwana is examined by Grantham et al. through a detailed comparison of lithological, structural, metamorphic and geochronological data from Mozambique with Sri Lanka and Dronning Maud Land. As a follow-up study to Grantham et al. (2003), a continuing ambitious mapping project of Mozambique has led these workers to

propose a model dominated by nappe tectonics during the 590– 550 Ma period of Gondwana assembly. Tectonic windows in Sri Lanka and Dronning Maud Land are considered as expressions of the c. 600 km displacement of crust from northern Gondwana by mega-thrusts. This concept will be tested by future developments.

Geological outline of East Antarctica (088E – 6088E) Since the proposal of Gondwana and Rodinia reconstruction models by Dalziel (1991) and Hoffman (1991), geoscientists have conceived of the Antarctic continent as a single stable entity between 1000 and 500 Ma. However, recent palaeomagnetic, geological and geochronological studies have recognized several distinct Neoproterozoic orogenic events within the East Antarctic shield, and a new concept has emerged of East Antarctica as a collage of three distinct Mesoproterozoic provenances: the Wilkes (1330–1130 Ma), Maud (1090–1030 Ma) and Rayner (990 –900 Ma) Provinces (Fitzsimons 2000; Meert 2003; and references therein). Moreover, two distinct age populations of 650–550 Ma and 580– 500 Ma have emerged in extensive geochronological datasets from the so-called ‘Pan-African orogeny’ (Fig. 2),


involving several discrete crustal blocks in East Antarctica and regions surrounding the East African –Antarctic Orogen (e.g. Jacobs et al. 2003a, b; Meert 2003; Hokada & Motoyoshi 2006). Provinces of Archaean age in East Antarctica are found at ‘Annadagstopane’ in Grunehogna, western Dronning Maud Land; the Napier Complex and Oygarden Islands in Enderby Land; the southern Prince Charles Mountains, Vestfold Hills and

5

Rauer Islands in Mac Robertson Land and Princess Elizabeth Land; the Denman Glacier in Queen Mary Land, and the Mawson Block in Terre Ade´lie (Fig. 3). The Grunehogna terrane is considered as a part of the Archaean Kalahari Craton in southern Africa (Jacobs et al. 1993b), and the Mawson Block has been correlated with the Gawler Craton in southern Australia (Fanning et al. 1996). Altogether, the terranes of East

Fig. 3. Continents that surrounded Antarctica in the Neoproterozoic. Geological entities within East Antarctica are also shown.

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Antarctica preserve a protracted crustal history from the oldest in the Napier Complex (c. 3850 Ma) to the last episode of post-collision magmatism (c. 450 Ma) in the waning stages of Gondwana amalgamation. Enderby Land lies between longitudes 458E and 608E, and comprises Archaean to early Proterozoic crustal sequences, representing a continental core complex surrounded by Proterozoic– Cambrian mobile belts. To the east of Enderby Land, more than 1000 km of coastal areas in Dronning Maud Land (58E– 458E) comprise late Proterozoic to Cambrian mobile belts (650– 500 Ma) (Fig. 4). This mobile belt has been extrapolated to Prydz Bay (Boger et al. 2001) and as far as western Australia (Fitzsimons 2000; Meert 2003). The only recent equivalent of such an extensive mobile belt is the Cenozoic Alpine– Himalayan orogenic belt. How far are these two orogens comparable, and where do they differ? Although there are countless similarities between the two, the former lacks the expression of low-temperature/ high-pressure metamorphic belts, which would provide clear equivalents to present-day subduction, accretion and collision-related tectonic settings that presumably would predate the amalgamation of continental blocks by the orogen. Within Dronning Maud Land, the lithological contrast between the inland mountain chains of central–eastern Dronning Maud Land (including the Sør Rondane, Belgica and Yamato Mountains) and outcrops along the Soya and Prince Olav

Coasts further east is striking (Fig. 4). The former region is dominated by felsic (granitic, granodioritic and syenitic) orthogneisses and post-tectonic plutons, with lesser mafic lithologies and metasedimentary sequences. In contrast, the latter region (Lu¨tzow-Holm and western Rayner Complexes) consist of voluminous metasedimentary rocks with mafic and calcareous rocks, and relatively little granitic material. It is important that any regional tectonic model accounts for this transition in the makeup of the mobile belt.

The Napier Complex The Napier Complex is one of several Archaean cratonic terranes (Fig. 3) in the East Antarctic continent (e.g. the Grunehogna terrane, the Ruker terrane in the southern Prince Charles Mountains, the Vestfold Hills, the Mawson Block in Terre Ade´lie, the Miller Range and the Shackleton Range), but is unique in being entirely composed of high-temperature granulites. Early Archaean (.3850 Ma) protolith ages have been obtained from tonalitic orthogneisses (Black et al. 1986; Harley & Black 1997; Kelly & Harley 2005), which are the oldest in Antarctica and close to the age of the Earth’s oldest known orthogneiss, the Acasta Gneiss in Canada (4000 Ma, Bowring et al. 1989). These 3850 Ma tonalitic orthogneisses occur at least in two localities (Mt. Sones and Gage Ridge; Harley & Black 1997; Kelly & Harley 2005), and subsequent tonalitic –granodioritic

Fig. 4. Tectonic units in Dronning Maud Land and Enderby Land, East Antarctica, showing salient geological and geochronological features. Dashed lines represent suspect boundaries between the units.


magmatism is observed from 3270 Ma (Mt. RiiserLarsen; Hokada et al. 2003) to 2630 Ma (Tonagh Island; Carson et al. 2002). The Napier Complex consists predominantly of tonalitic –granodioritic orthogneiss, but also includes mafic to ultramafic orthogneisses, garnet-bearing peraluminous granitic gneisses, and subordinate quartzo-feldspathic, siliceous and aluminous paragneisses. This lithological diversity indicates a complex and progressive development of the proto-metamorphic terrane, and provides insights into the development of continental crust during the Archaean. However, it is still unclear when and how the various crustal components were brought together, and what types of tectonic processes were functional in the Archaean. What is known is that the crustal components of the Napier Complex shared a common history after 2850 Ma, the timing of the first major regional magmatic–metamorphic event. Following extensive field and laboratory work by geologists of the Australian National Antarctic Research Expedition (ANARE), who established the geological structure and history of this area (see Sheraton et al. 1987, and references therein), JARE has carried out geological fieldwork intermittently throughout the 1980s and 1990s. Ishizuka reviews the voluminous results obtained by various JARE expeditions to the Napier Complex. In addition to the preparation of detailed geological maps (Ishikawa et al. 2000; Osanai et al. 2001), the expeditions focused on the geochemical characterization of different lithological units within the Napier Complex, which represent an admixture of Archaean components with sedimentary, granite– greenstone and tonalite–trondhjemite–granite (TTG) affinities. The review emphasizes important results obtained in subject areas such as the processes of ultrahigh-temperature (UHT) metamorphism, stages of protolith formation and geochemical studies of dykes, to provide constraints on modelling the tectonic evolution of the region. Our basic knowledge of the regional structural features in the Napier and Rayner Complexes is based on the mapping results carried out in the 1960s by the Soviet Antarctic Expedition (SAE; see Kamenev 1972, 1975) and further extensive geological mapping by ANARE until the late 1970s (Sheraton et al. 1987). Toyoshima et al. construct a regional form-line map based on structural data from published maps in the Napier and Rayner Complexes. They identify potential boundaries between different regions, based on the convergence of several structural parameters. The location of major tectonic boundaries is supported by geophysical evidence, as well as detailed field geological data from representative areas. In addition to the geological information from outcrops, the nature and properties of lower crustal

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materials deduced from geophysical studies are critical in imaging the present-day crust. Ishikawa et al. examine the seismic and elastic properties of lower crustal rocks from Enderby Land and Dronning Maud Land to provide insights into the lower crust of East Antarctica. They suggest a possible predominance of biotite-bearing continental crust. The Napier Complex experienced unusually high temperatures during metamorphism of 900–1100 8C on a regional scale, providing the first recognized instance of UHT metamorphism. Mineral parageneses diagnostic of UHT metamorphism, including sapphirine þ quartz, orthopyroxene þ sillimanite þ quartz and osumilite (Harley & Hensen 1990), have been recognized over a 200 km by 100 km area. The widespread distribution of UHT metamorphism, with estimated peak metamorphic temperatures in excess of 1120–1150 8C at relatively shallow crustal depths of 20–30 km (e.g. Harley & Motoyoshi 2000; Ishizuka et al. 2002; Harley 2004), requires explanation by unusual tectonic models to reasonably explain these crustal conditions. The terrane attracts great interest in how such extreme temperatures can be achieved in the mid- to lower crust, and represents a metamorphic end-member at the opposite extreme from the ultrahigh pressure (UHP) metamorphism found in continent–continent collision zones. UHT metamorphism in the Napier Complex is a phenomenon that never has been found on such a scale anywhere else in the world. Hokada et al. model the thermal and barometric behaviour of the lower continental crust. Based on an extensive analysis of petrological, structural and geochronological data, they estimate the lateral and vertical extent of UHT lithologies, and discuss the difficulties in providing models that can sustain a .1000 8C thermal regime for crustal thicknesses of 4–5 km. It is stressed that an enormous quantity of heat is necessary for achieving this, and that modelling requires an active role for asthenospheric input. Experimental and empirical studies on various chemical systems in metamorphic mineral assemblages are essential in determining the temperatures and pressures prevailing under UHT conditions. The solubility of titanium in quartz under UHT conditions is evaluated with the help of experiments by Kawasaki & Osanai on samples from Bunt Island, from which they develop an empirical geothermometer, and they test it by applying it to selected localities in Enderby Land and Dronning Maud Land. This method should find application in many future studies, as quartz and titaniumbearing minerals such as rutile and ilmenite are common constituents in high-grade metamorphic rocks. Sato et al. examine the partitioning behaviour of Fe2þ and Mg between orthopyroxene and

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spinel from UHT assemblages. Although the partitioning does not seem to record UHT conditions because of retrograde exchange, the results are reliable indicators of post-peak conditions. Fluid composition is a critical factor that controls metamorphism under UHT conditions without melting the rock. It is essential that the rocks should be anhydrous when the UHT conditions are attained. According to experimental studies (e.g. Johannes & Holtz 1996), even 1 wt% of water in a muscovite granite will lead to complete melting at UHT conditions. Therefore, the precursor rocks should either be essentially dehydrated during prograde metamorphism, or should have been previously anhydrous (by earlier metamorphism). The anhydrous mineral assemblages under UHT conditions were probably sustained by the presence of dry CO2-rich fluid. Characterization of fluids in UHT rocks from the Napier Complex is elegantly carried out by Tsunogae et al., who apply Raman spectroscopy to obtain the precise chemical composition of the fluids that were present during peak metamorphism. The ubiquitous presence of CO2 is demonstrated. Minor amounts of CH4 and N2 are also identified. Intriguingly, carbonate minerals present within the fluid inclusions further provide a unique window into the evolution of fluids during UHT metamorphism. CO2-rich fluid has an important, if not instrumental, role in UHT metamorphism, because it can be an effective heat transfer medium. High-T carbonic fluids from asthenospheric mantle to crust can effectively transfer heat into the crustal rocks, much faster and more easily than thermal conduction or convection. Enderby Land is also characterized by multiple episodes of dyke emplacement (Sheraton et al. 1987). The geochemical and tectonic significance of post-tectonic dykes is studied by Suzuki et al., who identify two distinct generations of dykes at Mt. Riiser-Larsen that exhibit contrasting source characteristics. An earlier 1.9–2.0 Ga generation of dykes is considered to have derived from a mantle wedge source, with possible connections with the continental crust formation of Rayner Complex. The less prominent 1.2 Ga dyke suite has ocean island basalt (OIB) or enriched mid-ocean ridge basalt (E-MORB) affinities. In addition to the emplacement of dykes, Enderby Land is also intruded by early Palaeozoic pegmatites. Carson & Ague evaluate geochemical element mobility associated with the infiltration of aqueous fluids in association with pegmatites, and model the depth of wall-rock metasomatism. They also suggest that the source for pegmatitic melts and aqueous fluids might be the underplating of sedimentary rocks by convergent tectonism between the Rayner Complex and the Napier

Complex, implying an early Palaeozoic timing for the juxtaposition of these terranes in the western part of Enderby Land.

The Rayner Complex The Rayner Complex was originally named for Proterozoic metamorphic lithologies adjacent to the Archaean Napier Complex (Kamenev 1972). It is made up of amphibolite- to granulite-facies orthogneisses and paragneisses, including pelitic, mafic, ultramafic and calcareous layers and boudins. Although the Rayner Complex was originally defined by lithologies south of the Napier Complex in Enderby Land, the main extent of the terrane is recognized to the east in Kemp Land and Mac Robertson Land. In the latter, the terrane is terminated eastwards by the late Neoproterozoic to Cambrian granulites of Prydz Bay, and southwards by metamorphic rocks and granitoids of similar age in the southern Prince Charles Mountains (Boger & Wilson 2005). The Rayner Complex involves the 990– 900 Ma granulite-grade reworking of supracrustal lithologies, deposited on a basement that is mostly Palaeoproterozoic in the eastern section, with an Archaean component closer to the Napier Complex (Kelly et al. 2002; Halpin et al. 2005). Extensive intrusions of charnockite were emplaced during metamorphism along the eastern margin of the Rayner Complex (Young & Black 1991). The grade and timing of metamorphism and charnockitic magmatism, along with the nature of protolithic crust, are shared with the Eastern Ghats of the Indian peninsula, and the two terranes are now regarded as having been a single tectonic entity attached to the cratonic core of India before the Neoproterozoic (e.g. Dobmeier & Raith 2003). Metamorphic reworking of the eastern Rayner –Eastern Ghats terrane in the late Neoproterozoic is limited to its margins (east and south in the Rayner Complex, north in the Eastern Ghats; Mezger & Cosca 1999). The geological evolution of the western part of the Rayner Complex is more problematic. A predominance of early Cambrian ages (Shiraishi et al. 1997; Motoyoshi et al. 2006) suggests that this area was reworked simultaneously with the metamorphism of the adjacent Lu¨tzow-Holm Complex. Metamorphic conditions, involving isothermal decompression after UHT peak metamorphism at Forefinger Point, are similar to those at Rundva˚gshetta, at the opposite end of the Lu¨tzowHolm Complex. However, major differences between the Lu¨tzow-Holm and western Rayner complexes include the observation of prograde kyanite inclusions in garnet, and the eastwarddecreasing grade of metamorphism in the former terrane. In addition, 800–700 Ma ages obtained in


the western Rayner Complex (Asami et al. 1997, 2005; Shiraishi et al. 1997) have not been found in the Lu¨tzow-Holm Complex. Regardless, the extensive Cambrian reworking leads us to redefine this region as the ‘Western Rayner Complex’, in contrast to the main body of the Rayner Complex and the Lu¨tzow-Holm Complex (Fig. 4). The geological significance of the Western Rayner Complex is a subject of continuing and future research, and further field and analytical studies are required to understand this complicated section of East Antarctica.

The Lu¨tzow-Holm Complex The Lu¨tzow-Holm Complex, located in eastern Dronning Maud Land (Fig. 4), is a late Neoproterozoic orogenic belt bounded by the late Mesoproterozoic Rayner Complex to the east and by the late Neoproterozoic to early Palaeozoic Yamato– Belgica Complex to the west (Shiraishi et al. 1992, 1994, 2003). It is a significant area for the investigation of the final collision between East and West Gondwana, because the Lu¨tzow-Holm Complex is considered to be a southern extension of the suture between them (e.g. Shiraishi et al. 1994; Fitzsimons 2000). The geology of this complex has been reviewed in several earlier studies (Hiroi et al. 1983, 1986, 1987, 1991; Shiraishi et al. 1994, 2003). The Lu¨tzow-Holm Complex is composed of high-grade metamorphic rocks, including pelitic– psammitic gneisses, mafic to intermediate basic gneisses, subordinate lenses of ultramafic gneiss, marbles and calc-silicate rocks. Felsic pegmatitic dykes discordantly intrude the metamorphic rocks. Ultramafic lenses that were probably derived from oceanic crust are distributed across the central and southwestern part of the complex (Hiroi et al. 1986). Hiroi et al. (1991) postulated that the ultramafic lenses represent dismembered fragments of an ophiolite complex derived from the missing oceanic crust between older continents, now represented by the Yamato– Belgica and Rayner Complexes. The detailed structural evolution of the Lu¨tzow-Holm Complex has not yet been fully understood, although some parts of the complex have been structurally described in several studies (e.g. Kizaki 1962, 1964; Ishikawa 1976; Yoshida 1977, 1978; Matsumoto et al. 1979, 1982; Ishikawa et al. 1994; Motoyoshi & Ishikawa 1997; Ikeda & Kawakami 2004; Kawakami & Ikeda 2004a, b; Michibayashi et al. 2004; Osanai et al. 2004; Okamoto & Michibayashi 2005). The metamorphic grade of the complex progressively increases from upper amphibolite facies on the Prince Olav Coast to granulite facies in Lu¨tzow-Holm Bay (Hiroi et al. 1991), with a

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‘thermal axis’ of maximum peak temperature estimated to lie at the southern end of Lu¨tzow-Holm Bay, near Rundva˚gshetta (Motoyoshi 1986). Several lines of evidence suggest that the Lu¨tzow-Holm Complex has experienced a typical ‘clockwise’ P –T path. These include prograde kyanite and staurolite as relict inclusions in garnet or plagioclase (Hiroi et al. 1983; Motoyoshi 1986; Kawakami & Motoyoshi 2004; Satish-Kumar et al. 2006b), and reaction textures in ultramafic rocks (Hiroi et al. 1986) are also significant. It has been observed that paragneisses from the Prince Olav Coast experienced the reaction staurolite ¼ garnet þ aluminosilicate þ spinel þ H2O within the sillimanite stability field, whereas those from Lu¨tzow-Holm Bay experienced the reaction in the kyanite stability field (Hiroi et al. 1983, 1987). This petrographical evidence is peculiar among high-grade metamorphic terranes in East Antarctica, as no obvious prograde P –T paths have been reported except for the Lu¨tzow-Holm Complex (Harley & Hensen 1990). UHT peak metamorphic conditions of about 1000 8C and 11 kbar, and subsequent isothermal decompression have been reported from Rundva˚gshetta (Kawasaki et al. 1993; Ishikawa et al. 1994; Motoyoshi & Ishikawa 1997). Yoshimura et al. present further petrological evidence for UHT metamorphism at Rundva˚gshetta (Fig. 1). The coexistence of sapphirine and quartz within garnet porphyroblasts, high Al contents of orthopyroxene and temperature estimates based on ternary feldspar thermometry suggest that the rocks in this region were metamorphosed above temperatures of 1000 8C. In the neighbouring Skallen region (Fig. 1), Goto & Ikeda present crystal size distributions (CSDs) of garnet in quartzo-feldspathic gneisses metamorphosed at above 800 8C. They attempt to provide reasons for the differences in garnet nucleation and growth between layers. Based on the crystal size distribution of garnet they predict less predominance of Ostwald ripening, even at granulite-facies conditions, in the absence of fluids. The timing of the peak regional metamorphism has been estimated by sensitive high-resolution ion microprobe (SHRIMP) U –Pb zircon dating at between 521 + 9 and 553 + 6 Ma (Shiraishi et al. 1992, 1994, 2003). Zircon from syn-deformational leucosome has a U –Pb age of 517 + 9 Ma, which is interpreted as a melt crystallization age (Fraser et al. 2000). Fraser et al. (2000) suggested from combined SHRIMP zircon analyses and Ar –Ar hornblende and biotite chronology that post-peak decompression and subsequent cooling to c. 300–350 8C took place within a time interval of c. 520–500 Ma. A summary of recent dating results, has been given by Nishi et al. (2002) and references therein. Recently, however, in situ

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monazite chemical Th –U –total Pb isochron method (CHIME) dating and zircon SHRIMP dating combined with the microstructural observation of monazite and zircon by Hokada & Motoyoshi (2006) yield ages of 650 –580 Ma and 550– 520 Ma for monazite in garnet-bearing felsic gneisses from the Skallen region. Based on the medium to heavy REE (MREE –HREE)-enriched nature of 650 –580 Ma monazite, Hokada & Motoyoshi interpreted the older ages as monazite growth under prograde, garnet-absent conditions, whereas the 550– 520 Ma age group represents monazite grown at peak metamorphism in the presence of garnet. Dunkley (2007), reporting a similar spread of ages from 600 to 500 Ma, interpreted the age range as also reflecting the progressive growth of zircon at various stages during a single clockwise P–T history of the complex. These contrasting interpretations will be tested in the near future by petrological and microstructural studies, to find out whether the unexpectedly long duration of a single metamorphism in the Lu¨tzowHolm Complex is feasible (Dunkley 2007). Miyamoto et al. review the chronology of events after peak metamorphism, and present new Sm– Nd and Rb – Sr ages for key metamorphic rocks in the southwestern Lu¨tzow-Holm Complex. Two possible explanations are put forward for postmetamorphic thermal perturbations in the region, involving either cooling and uplift of the terrane, or reheating by magmatic and associated metasomatic activity. Proterozoic and Archaean detrital cores of zircon grains from Rundva˚gshetta and West Ongul Island (Shiraishi et al. 1994; Fraser 1997) demonstrate ancient provenance in the metasediments of the Lu¨tzow-Holm Complex. Satish-Kumar et al. focus on isotopic compositions and geochemical characteristics of high-grade marbles from the Lu¨tzow-Holm Complex. From earlier studies the inferred depositional ages of sedimentary protoliths in the Mozambique Ocean that separated East and West Gondwana is some time between c. 630 Ma (the earliest metamorphic age reported by Hokada & Motoyoshi 2006) and the youngest Sm–Nd model age of c. 850 Ma (Shiraishi et al.). Carbon, oxygen and strontium isotopic compositions indicate that most metacarbonate rocks were altered by multiple episodes of fluid activity, related to pre-peak, peak and post-peak metamorphic events. By applying multiple geochemical criteria, nearpristine sedimentary signatures were identified in some layers, which when compared with the nonmetamorphic chemostratigraphic curves suggest a depositional age between 830 and 730 Ma. Along the Prince Olav Coast, Cape Hinode (Fig. 1) is an exceptional outcrop where the late Neoproterozoic ages are completely absent and

only a c. 1000 Ma age has been reported (Shiraishi et al. 1994, 2003). Grenvillian ages have been reported from three other localities from the Lu¨tzow-Holm Complex, including Skavsnes (Fraser 1997), Telen and Innhovde (Shiraishi et al. 2003). All of these represent inherited cores of zircon with magmatic zoning, and no c. 1000 Ma metamorphic overgrowths have been found. Therefore, Shiraishi et al. (2003) interpreted the c. 1000 Ma age as representing localized igneous activity. Hiroi et al. (2006) have suggested, on the basis of U – Th–Pb ages reported by Shiraishi et al. (1994, 2003) and Motoyoshi et al. (2004), that the gneisses of Cape Hinode are exotic to other parts of the Prince Olav Coast. Xenocrystic garnet and kyanite in adakitic trondhjemites and tonalities from Cape Hinode are treated by Hiroi et al. as phases that were entrained in Mesoproterozoic tonalitic magmas. Kyanite is a stable matrix phase in Cape Hinode metapelites, contrary to the mode of occurrence of kyanite as relic inclusions within garnet in other parts of the Lu¨tzow-Holm Complex. The lack of 600–500 Ma ages from Cape Hinode also supports the notion of an allochthonous block emplaced in the waning stages of amalgamation of East and West Gondwana. The major age population of 1080– 1000 Ma reported from Cape Hinode is comparable with that of the Maud Province to the west, rather than with the 990–900 Ma ages of the closer Rayner Complex. Continuation from Cape Hinode to the Vijayan Complex of Sri Lanka is possible, with extensions to Mozambique and the Natal Belt (Hiroi et al. 2006). Alternatively, Cape Hinode may represent an isolated block, as implied by gravity and geomagnetic data (Nogi et al. 2006). Pre- to syn-metamorphic granitic rocks are characterized by the irregular shape of the intrusive boundaries, intergradational contacts and intense deformation, and only one of them from Breidva˚gnipa has been dated at 576 Ma by the Rb –Sr whole-rock isochron method (Shimura et al. 1998). Post-tectonic granitic dykes intrude across gneissic fabrics throughout the Lu¨tzowHolm Complex. These granites have been dated by the Rb–Sr whole-rock isochron method as younger than 500 Ma (e.g. Nishi et al. 2002; Ajishi et al. 2004). Suda et al. (2006) carried out a geochemical study of metabasites (mostly garnet-absent) in the Lu¨tzow-Holm Complex. Suda et al. further studied the geochemical and isotopic composition of metamorphosed ultramafic and mafic rocks, and distinguish those of the eastern part of the Lu¨tzow-Holm Complex as derived from immature continental crust during the Mesoproterozoic, from those in the western part as products derived from a matured crust. These results further establish the changing tectonic


environment in eastern Dronning Maud Land during the Neoproterozoic. Because of the systematic and gradual southwestward increase of metamorphic grade from upper amphibolite facies to UHT conditions, the Lu¨tzow-Holm Complex provides a good example for study of the behaviour of melts, fluids and accessory minerals under these conditions. Satish-Kumar et al. (2006a) studied scapolite boudins from Skallen and presented detailed petrographical and geochemical evidence for changing fluid composition from scapolite phase equilibria. Kawakami et al. (2006) reported the mode of occurrence of sulphide minerals throughout the Lu¨tzow-Holm Complex and found that sulphide inclusions are completely different in composition and species from those in the rock matrix, retaining information from peak metamorphism. Inclusion sulphides were mostly restitic in composition, suggesting the loss of sulphide melt from the rocks of the Lu¨tzowHolm Complex during anatexis. Kawakami et al. characterize the occurrence of kornerupine in mafic and ultramafic rocks from Akarui Point. They propose possible sources for boron through aqueous fluids derived from sediments or hydrothermal alteration of protoliths by seawater.

The Yamato – Belgica Complex The Yamato –Belgica Complex is also thought of as a late Neoproterozoic to Cambrian orogenic terrane between the Lu¨tzow-Holm Complex and the Sør Rondane Mountains (Fig. 4). It consists of two inland mountain ranges, the Yamato and Belgica Mountains. The area is characterized by widespread granite and syenite intrusions with minor amphibolite-facies metamorphic rocks of quartzo-feldspathic and intermediate composition (Shiraishi et al. 1994). Rare granulite-facies rocks with peak metamorphic conditions of 700 –750 8C and ,5 kbar are found, but the relationship between amphibolite-facies and granulite-facies rocks is uncertain. Age constraints for this area are mainly from zircon SHRIMP data by Shiraishi et al. (1994, 2003) that range from 1000 to 500 Ma, with the exception of one spot yielding an age of c. 2500 Ma. Quartz monzonite and granitic gneiss from the Yamato Mountains yielded an age of 535 Ma, which is interpreted as the timing of amphibolite-facies metamorphism and magmatism. These events followed the widespread syenite magmatism of the area, but the actual timing is not well constrained. Although there are not enough data available to establish the Proterozoic– Cambrian history of this area, the lack of essential .1000 Ma ages suggests juvenile crustal formation in the late Mesoproterozoic, similar to the Sør

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Rondane Mountains and Central Dronning Maud Land to the west, and in marked contrast to the Lu¨tzow-Holm Complex.

The Sør Rondane Mountains Outcrops in the Sør Rondane Mountains are dominated by Mesoproterozoic crustal lithologies that vary from predominantly arc-related material to continental materials from north to south (Shiraishi et al. 1991; Grew et al. 1992; Osanai et al. 1992). A semi-ductile shear zone divides the region into a northeastern granulite-facies terrane and a southwestern amphibolite-facies terrane. Recently, Asami et al. (2007) estimated peak granulite-facies metamorphism at temperatures of 860– 895 8C and pressures of around 12 kbar for the NE terrane. Furthermore, they found evidence for retrograde metamorphism under amphibolite-facies conditions. Extensive geochronological results presented by Shiraishi et al. suggest that crustal formation in the Sør Rondane Mountains occurred in the late Mesoproterozoic, and that the NE and SW terranes were juxtaposed around c. 570 Ma under amphibolite-grade metamorphic conditions, subsequent to higher temperature metamorphism at c. 600 Ma that affected only the NW terrane. An exact picture of late Neoproterozoic to Cambrian terrane amalgamation and tectonic evolution of the Sør Rondane Mountains requires further field studies, which are being conducted by JARE between 2007 and 2010.

Central Dronning Maud Land High-grade metamorphic rocks intruded by voluminous igneous bodies form coastal and inland mountainous outcrops in central Dronning Maud Land (CDML), from 28 to 148E (Dallmann et al. 1990). Metamorphic rocks in this region comprise banded gneisses and migmatites, whereas igneous rocks are mainly of charnockitic, syenitic and granitic composition (Ohta 1999). Two tectonothermal events have been distinguished in the region, at c. 1100 Ma and between 560 and 490 Ma (Jacobs et al. 1998, 2003a; Paulsson & Austrheim 2003). The younger event is generally considered as part of the East African–Antarctic Orogeny and involves an early collisional event at c. 560 Ma followed by large-scale extension associated with voluminous granitic magmatism. A variety of rock types are found in the CDML, including pelitic granulites, garnet-, biotite- and/or hornblendebearing gneisses, charnockites, mafic granulites and calc-silicate rocks. An early Grenvillian age (c. 1150 Ma) for granulite-facies metamorphism, followed by amphibolite-facies metamorphism at c. 560 Ma, is ascribed to these rocks. In addition,

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c. 630 Ma ages have been obtained from the coastal outcrop at Schirmacher Oasis, suggesting a different evolution of this area in the late Neoproterozoic compared with that of the inland mountains. A recent study by Bisnath et al. (2006) proposed a two-stage collision model, involving an initial arc– continent collision followed by continent– continent collision. Baba et al. compare the metamorphic evolution of Schirmacher Hills with that of Mu¨hligHofmanfjella and find that, although there is no clear difference in peak P–T conditions, the retrograde P–T paths contrast between these two regions. They suggest that the Schirmacher Hills could be part of SE Africa, whereas the inland mountain regions were part of crust formed during the final amalgamation of East Gondwana. Owada et al. consider the geochemical characteristics of post-tectonic mafic dykes and find that the parental magma was derived from a metasomatized mantle source. Based on a detailed evaluation of Sr and Nd isotope systematics of CDML and the Sør Rondane Mountains, they suggest the possibility of a suture zone of East and West Gondwana transition between these two regions.

Emerging thoughts and future perspectives Geophysical studies The continuing compilation of aeromagnetic, marine and satellite-based surveys by the Antarctic Digital Magnetic Anomaly Project (ADMAP) provides the best picture of the internal architecture of East Antarctica, with the latest versions of the East Antarctic magnetic anomaly map and the Antarctic Digital Magnetic Anomaly Map published by Golynsky (2007) and von Frese et al. (2007), respectively (http://www.geology.ohio-state.edu/ admap/). Geodynamic models of the assembly of various terranes between and during cycles of supercontinent formation need to take into account the regional-scale structural information that magnetic anomaly maps provide. A belt of high magnetic anomalies that curves around the coastal Grunehogna craton in western Dronning Maud Land is correlated with the c. 1.1 Ga Namaqua–Natal mobile belt in South Africa, which shows a similar pattern of anomalies (Golynsky & Jacobs 2001). In contrast, a broad area of low magnetic signature extends across central and eastern Dronning Maud, which corresponds well to the interpretation of this region as Mesoproterzoic felsic crust incorporated into the broad East African– Antarctic Orogen (Jacobs et al.; Shiraishi et al.). This geomagnetic domain

has an abrupt north –south trending termination against a region of positive magnetic anomalies, just east of the Yamato Mountains, that corresponds exactly to the terrane boundary between the Sør Rondane Mountains and Yamato– Belgica Complex and the Lu¨tzow-Holm Complex inferred by Shiraishi et al. However, in other key areas of Late Neoproterozoic geological activity, especially around Prydz Bay and Lu¨tzow-Holm Bay, there is a significant discrepancy between the latest tectonic models made on the basis of surface geology (field geology, petrography and geochronology) and geophysics (aeromagnetic mapping). Golynsky et al. (2002) and Golynsky (2007) suggested that the presence of intense east –west linear anomalies, which extend across the Lambert Graben from the northern Prince Charles Mountains to Prydz Bay, and from the southern Prince Charles Mountains to the Grove Mountains, implies a tectonic association of these areas that predates late Neoproterozoic activity in the region. These features were associated by Golynsky (2007) with paired east – west-trending belts of negative and positive anomalies that extend from Prydz Bay to Lu¨tzow-Holm Bay, where the boundaries of these belts rotate into a trend perpendicular to the Prince Olav Coast. These belts are interpreted as late Mesoproterozoic terranes, corresponding to the Rayner Complex, that suture together the Archaean terranes of the Napier Complex and the Ruker terrane. The model implies that late Neoproterozic metamorphism and magmatism observed in Prydz Bay and the southern Prince Charles Mountains is unrelated to that found in Lu¨tzow-Holm Bay. Counter to tectonic models by Boger et al. (2001) and Phillips et al. (2006) that involve the collision of an Indo-Antarctic continent with inner Antarctica during the formation of Gondwana, Golynsky (2007) attributed all late Neoproterozoic activity to within-plate processes, similar to the concept proposed for the Grenvillian circumAntarctic mobile belt (Yoshida 1992, 2007). However, such a interpretation neglects evidence of the heterogeneous nature of crust modified by late Proterozoic metamorphism in the Lu¨tzowHolm Complex, as indicated by a diversity of protolith and crustal model ages (Shiraishi et al.; Suda et al.), and continental to oceanic geochemical signatures (Satish-Kumar et al.; Suda et al.; Hiroi et al.). In the future, integration of geomagnetic data with surface petrology and geochronology should resolve these issues. Another technique to obtain basement geological data inland is ice core drilling that continues to reach basement rocks. There is a two-fold benefit involved in continental ice core drilling, as both palaeoenvironmental and palaeocontinental problems can be solved in a single project. Ice


core drilling projects at Vostok and Dome Fuji have returned promising results and technical know-how on pursuing drilling in subzero conditions. In fact, the Vostok drilling has succeeded in collecting sediment from the bottom of the ice sheet, and preliminary SHRIMP dating of zircon and monazite yielded a range of ages between 1.8 and 0.6 Ga, similar to those seen in coastal mobile belts and a further indication of the pervasive involvement of Mesoproterozoic and Neoproterozoic geological activity in the formation of East Antarctica (Rodionov et al. 2006). To solve the problems of palaeocontinental uncertainties relating to obscurity of the inland Antarctic continent it will be necessary to gather information from inland regions.

Field-based studies Because of the low-latitude climate, lack of rainfall, and absence of vegetation (excepting mosses and lichen), outcrops in Antarctica provide high-quality field information for geological studies. Mechanical

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weathering by the action of wind and glacial abrasion, and limited chemical or hydrothermal alteration, results in the exposure of fresh outcrops that are perfectly suited for multidisciplinary geological studies (Fig. 5). We have identified the following key localities in eastern Dronning Maud Land, which need further attention to improve our understanding of the Archaean to early Palaeozoic evolution of East Antarctica. Enderby Land is a potentially important area for studies not only for clarifying the tectonism in the Archaean but also for understanding lower crustal processes. This region can enlighten us further about: (1) the formation of continental crust in the Archaean; (2) the causes and consequences of unusually high-temperature (.1000–1150 8C) metamorphism in the Napier Complex; (3) Proterozoic suturing between the Archaean cratons of India (e.g. Dharwar –Napier) and Archaean terranes in the southern Prince Charles Mountains of Antarctica; (4) subcontinental mantle dynamics,

Fig. 5. Illustrative outcrops in East Antarctica, and their potential for future research. (a) Field photograph showing the regional distribution of UHT metamorphic rocks in the Napier Complex at Tonagh Island (JARE-38). The Napier Complex is a key area in understanding the crustal evolution in the Archaean. (b) Metacarbonate and paragneiss sequences at Skallevikshalsen (JARE-46), with potential for understanding the depositional environment of sediments between East and West Gondwana. (c) Partial melting and melt segregation as seen in the paragneisses at Skallevikshalsen (JARE-44); this is a topic of prime importance for understanding the generation, segregation and movement of melts in middle to deep continental crust. (d) The inland nunataks of the Sør Rondane Mountains (JARE-49). Geological evidence from these nunataks may clarify the history of amalgamation of East and West Gondwana.

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as revealed by Proterozoic dyke swarms; (5) the amalgamation of the Napier and Rayner Complexes with other terranes in the formation of Gondwana. Moving west, the Lu¨tzow-Holm Complex has well-preserved regional amphibolite to UHT metamorphic zones with classic clockwise P –T trajectories. The problems that remain to be solved include: (1) the unravelling of the earlier peak granulite to UHT metamorphism and later extensive rehydration; (2) the significance of dual 600– 550 and 550– 500 Ma events in a regional context of Gondwana amalgamation; (3) the provenance and tectonic setting of volcano-sedimentary sequences and basement lithologies. The Yamato and Belgica Mountains are constituted mostly of felsic orthogneisses and syntectonic plutonic rocks. There are fewer suitable lithologies for the detailed characterization of the metamorphic P–T evolution for comparison with the neighboring Lu¨tzow-Holm Complex and the Sør Rondane Mountains. However, preliminary information regarding the geochemical features of plutonic rocks needs to be further developed to determine if the Mesoproterozoic juvenile crust identified in central Dronning Maud Land and the Sør Rondane Mountains extends to this area. Further west, the Sør Rondane Mountains is an important area from a regional geological point of view. This area is seen as critical in finding solutions to longstanding problems on the suturing of East and West Gondwana. The variety of ages recorded in this region may be critical in distinguishing the 600– 550 and 550 –500 Ma conjugate tectonic belts and the order in which crustal fragments amalgamated to form Gondwana.

Understanding geological extremes The geology of East Antarctica not only has provided a regional framework for supercontinent correlation studies but also has been critical in understanding some of the most extreme geological phenomena in crustal regimes. One such extreme is UHT metamorphism. It is increasingly accepted that UHT conditions exist in the continental crust; however, it is still a challenge to understand the factors that control such unusual thermal regimes. In this perspective, more accurate physical parameters and much tighter temporal constraints of such extreme conditions need to be determined. It is also essential to understand the total heat budget, the quantity of heat added to a certain initial or steady-state condition, and whether other factors such as fluids played a role. Precise determination of physical conditions under extreme crustal metamorphism is essential in modelling crustal evolution. It is a challenge to duplicate these conditions in the laboratory,

although recent developments in experimental petrology can achieve this, except for the time factor. Microstructures in minerals, especially exsolution textures, are now recognized as powerful tools for recovering high-pressure and -temperature conditions prior to cooling and exhumation. Typical examples are the recovery of pigeonite compositions from orthopyroxene with Ca-clinopyroxene lamellae (e.g. Harley 1987; Ishizuka et al. 2002), recovery of single-phase compositions from ternary mesoperthitic feldspar (Hokada 2001) and Ti exsolution in quartz or in garnet (Kawasaki & Osanai). However, this technique needs caution in selecting suitable compositional ranges to recover such information and thermodynamic models to be applied for temperature estimation (e.g. Hokada & Suzuki 2006). The formation and preservation of UHT rocks in the crust is essentially controlled by the fluid regime during prograde metamorphism. Dehydration of rocks prior to partial melting is essential to restrict the melt fraction to a critical melting proportion, as larger proportions of melt can destroy the solid rock structure. In other words, UHT metamorphism should be observed only in rocks that are more or less in restitic nature, and potentially anhydrous UHT rocks may be widely distributed in the deepest continental crust worldwide. Composition of fluid also strongly controls the melt fraction. CO2-rich fluid flow from deeper (and hotter) crust transfers heat to shallower crust more effectively than conduction or convection. In addition, we also need to pay attention to the different cooling and uplifting processes that result in the exposure of extremely metamorphosed rocks without completely destroying the original parageneses. UHT metamorphism with subsequent isothermal decompression can be readily achieved by crustal uplift, and internal radioactive heat production in thickened crustal is a potential source of heat. In contrast, UHT metamorphism with isobaric cooling is problematic; that is a scenario that may be achieved when the heat source is local and magmatic (e.g. Bamble terrane in Sveconorwegian or Wilson Lake in Canada, where UHT metamorphic zones are developed around anorthosite bodies). Therefore, our fundamental understanding of extreme crustal processes remains primitive. The Napier Complex in East Antarctica is perfectly suited to understand the occurrence and importance of geological extremes.

Nanoscience and supercontinents: recent technological realms The past two decades have seen wide application of electron microprobe and ion microprobe techniques to investigate the chemical and isotopic


composition of minerals, especially the accessory phases, on a micrometre scale. Accessory phase behaviour with regard to trace element geochemistry (including REE, P, Zr, Ti, U and Th) is a major topic in the microanalytical world and has potential in resolving many problems relating to the evolution of continents. In recent years, barriers have been broken in linking the isotopic record with petrology in complex and multiply metamorphosed and deformed terranes (Rubatto 2002; Mu¨ller 2003; Vance et al. 2003). Within our reach is a new phase in accessory mineral research that will unravel complicated metamorphic and tectonic histories. Submillimetre-scale techniques are required to distinguish events in the multiply reactivated mobile belts of East Antarctica. Effective strategies include: (1) U –Th –Pb dating on a sub-grain, micrometre scale of zircon, monazite, apatite, titanite, rutile, perrierite and other accessory minerals by ion microprobe, electron microprobe, and laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS); (2) tying in the microstructural context of accessory phase growth with geochronology and chemistry, using highresolution secondary electron, back-scattered electron, and cathodoluminescence imaging under a scanning electron microscope; (3) experimental and empirical approaches to understanding the stability and chemical behaviour of accessory phases during deformation, metamorphism and partial melting (e.g. Harrison & Watson 1983; Watson & Harrison 1983); (4) integration of metamorphic and magmatic ages obtained by microbeam techniques with Lu –Hf isotope model ages for understanding the crustal extraction history (e.g. Kemp et al. 2006); (5) understanding Precambrian crustal evolution from non-radiogenic isotopes such as oxygen (e.g. Cavosie et al. 2005). Application of these advanced analytical techniques in East Antarctica will help in formulating reasonable geodynamic models of pre-Gondwanan supercontinent evolution.

Supercontinent cycle, global tectonics and Earth’s environment The Neoproterozoic to early Cambrian period was a time of extensive global tectonic activity that culminated in the amalgamation of the supercontinent Gondwana. This time span is also well known for phenomenal changes in climatic conditions that predated the Cambrian explosion in biodiversity. However, extreme climate change models invoking a ‘Snowball Earth’ (Hoffman et al. 1998) or drastic changes in Earth’s obliquity (Evans 2000) do not seem to satisfactorily explain the complex scenario (Meert 2007). Several lines of evidence have

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started to appear for the role of global tectonism in creating an environment conducive to biological activity (e.g. Maruyama & Santosh 2008; Meert & Liebermann 2008; Stern 2008). Undisputedly, a key factor that controls global environment is CO2 concentration in the atmosphere. Although volcanogenic-CO2 input to the atmosphere seems to be a potential source of sudden large-scale climatic variations, other sources such as CO2 transfer through orogenesis and oxidation of an earlier biosphere cannot be neglected. Variations in atmospheric CO2 in the past are clearly recorded by carbon isotope excursions in carbonate sediments that record conditions in palaeo-oceans. Examples of Neoproterozoic carbon isotopic excursions combined with geological evidence convincingly indicate two major and several minor glaciation events (e.g. Halverson et al. 2005). It still remains unclear how much the closure of oceans between the continents has a bearing on global climate change. Furthermore, it is perceived that the spatial extent of Neoproterozoic orogenic belts retained in the present day continental crust must have been a few orders of magnitude larger than what we see in the Cenozoic Alpine –Himalayan Orogeny. The impact of Neoproterozoic amalgamation of the Gondwana supercontinent on the global environment is yet to be clarified and information from East Antarctica is crucial in solving this problem.

Concluding remarks Studies on Antarctica have considerably refined our knowledge on the geodynamic evolution of continental crust. We envisage Antarctica as a model in Earth Science studies, in the advancement of science and for the peaceful living of mankind. However, because of its remoteness and extreme weather, Antarctica is still a difficult place to carry out geological fieldwork. As discussed above, many fundamental problems remain unsolved. However, progress through international collaboration can efficiently tackle this handicap. The future of Antarctic geoscience research seems bright through collective effort from different countries, and will be a driving force for the advance of our understanding of the history of the Earth. Enormous progress has been achieved in the past 50 years of geological research in East Antarctica. However, technology has overwhelmingly overtaken the pace of basic scientific research. The incongruity between basic scientific research and the momentum with which information is provided by the latest technology is challenging the whole world of science itself. Geoscience research is no exception to this trend. Being part of the natural

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sciences, geoscience research can act as a link between progress achieved in basic science and technology that can effectively transfer information for society. Henceforth, the keyword for the future is ‘Earth system science’, where natural science can sustain life and vice versa. Antarctica is an ideal place for resolving the complex problems of geoscience studies. MS-K acknowledges grant (No. 18740319) from the Ministry of Education, Culture, Sports, Science and Technology, Japan. Many of the ideas presented in this paper have evolved through discussions in symposia and informal meetings. The authors, therefore, cannot single out any individuals who might have influenced the contents. We thank the geoscientific community as a whole. The motive behind this paper was to explore new avenues for future geosciences research in the Antarctic continent; we suspect that we have hardly skimmed the topic. We owe gratitude to the numerous scientists who have taken part in JARE expeditions over the past 50 years. Readers who are interested in obtaining more information on JARE activities may visit the website http://ci.nii.ac.jp/organ/journal/ INT1000001377_en.html, where publications of the National Institute of Polar Research, Tokyo, are available for open access. We thank B. Pankhurst and B. Thomas for their constructive comments which helped improve the style and content of this contribution. Finally, the present paper would not have seen light without the constant support and encouragement to the senior author from the co-editors of the volume, K. Shiraishi, Y. Motoyoshi, Y. Hiroi and Y. Osanai, and in particular P. Leat, the Geological Society of London editor-in-charge of the volume. We express our sincere gratitude to all of them.

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Geochronological constraints on the Late Proterozoic to Cambrian crustal evolution of eastern Dronning Maud Land, East Antarctica: a synthesis of SHRIMP U – Pb age and Nd model age data KAZUYUKI SHIRAISHI1,2, DANIEL J. DUNKLEY1, TOMOKAZU HOKADA1,2, C. MARK FANNING3, HIROO KAGAMI4 & TAKUJI HAMAMOTO4,5 1

National Institute of Polar Research, 1-9-10, Kaga, Itabashi-ku, Tokyo 173-8515, Japan (e-mail: [email protected]) 2

Department of Polar Studies, Graduate School of Advanced Studies (SOKENDAI), Tokyo 173-8515, Japan

3

Research School of Earth Sciences, The Australian National University, A.C.T 0200, Australia 4

Graduate School of Science and Technology, Niigata University, Niigata 950-2181, Japan

5

Present address: Dia Consultants Co. Ltd, Kanayama-cho 1-6-12, Nagoya 456-0002, Japan Abstract: In eastern Dronning Maud Land (DML), East Antarctica, there are several discrete, isolated magmatic and high-grade metamorphic regions. These are, from west (c. 208E) to east (c. 508E), the Sør Rondane Mountains (SRM), Yamato– Belgica Complex (YBC), Lu¨tzowHolm Complex (LHC), Rayner Complex (RC) and Napier Complex (NC). To understand this region in a Gondwanan context, one must distinguish between Pan-African and Grenvillian aged magmatic and metamorphic events. Sensitive high-resolution ion microprobe U –Pb zircon ages and Nd model ages for metamorphic and plutonic rocks are examined in conjunction with published geological and petrological studies of the various terranes. In particular, the evolution of the SRM is examined in detail. Compilation of Nd model ages for new and published data suggests that the main part of eastern Dronning Maud Land, including the SRM, represents juvenile late Mesoproterozoic (c. 1000–1200 Ma) crust associated with minor fragments of an older continental component. Evidence for an Archaean component in the basement of the SRM is lacking. As for central DML, 1100– 1200 Ma extensive felsic magmatism is recognized in the SRM. Deposition of sediments during or after magmatism and possible metamorphism at 800–700 Ma is recognized from populations of detrital zircon in metasedimentary rocks. The NE Terrane of the SRM, along with the YBC, was metamorphosed under granulite-facies conditions at c. 600–650 Ma. The SW and NE Terranes of the SRM were brought together during amphibolite-facies metamorphism at c. 570 Ma, and share a common metamorphic and magmatic history from that time. High-grade metamorphism was followed by extensive A-type granitoid activity and contact metamorphism between 560 and 500 Ma. In contrast, TDM and inherited zircon core ages suggest that the LHC is a collage of protoliths with a variety of Proterozoic and Archaean sources. Later peak metamorphism of the LHC at 520–550 Ma thus represents the final stage of Gondwanan amalgamation in this section of East Antarctica.

The eastern part of Dronning Maud Land (DML) between 378E and 508E, which represents the original definition of Dronning Maud Land as a whole, was discovered by the Norwegian aviator Riiser-Larsen in 1930. Most of the coastal and inland outcrops were air-photographed by the Lars Christensen Expedition in 1937 and by Operation Highjump in 1946– 1947. For the International Geophysical Year (IGY) of 1957, the Japanese Antarctic Research Expedition (JARE) established the first wintering station in East Ongul Island, located in the mouth of Lu¨tzow-Holm Bay. Since then, the footprints of geologists have covered eastern DML from 208E to 508E.

Prior to JARE, the Soviet Antarctic Expedition started to investigate an extensive area of the Indian Sector of the East Antarctica and the Australian Antarctic Research Expedition (ANARE) worked eastwards from 458E. In particular, ANARE performed a large-scale geological investigation of Enderby Land between 1974 and 1980. Belgium started wintering during IGY on Princess Ragnhild Coast and conducted pioneer investigations in the Sør Rondane Mountains until 1961. Amongst the various national expeditions, geologists from many countries have participated in field programmes in eastern DML and Enderby Land.

From: SATISH -K UMAR , M., MOTOYOSHI , Y., OSANAI , Y., HIROI , Y. & SHIRAISHI , K. (eds) Geodynamic Evolution of East Antarctica: A Key to the East– West Gondwana Connection. Geological Society, London, Special Publications, 308, 21– 67. DOI: 10.1144/SP308.2 0305-8719/08/$15.00 # The Geological Society of London 2008.

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K. SHIRAISHI ET AL.

The results of a decade of geological surveys following IGY were compiled in the Antarctic Map Folio Series (Bushnell & Craddock 1969, 1970). The early stages of field investigations by JARE were devoted to geological mapping, resulting in 36 sheets of geological maps on the scales of 1:5000 to 1:250 000, published as the Antarctic Geological Map Series by the National Institute of Polar Research, Tokyo. Reviews of the basement geology of East Antarctica based on the studies before the 1980s were published by Ravich & Kamenev (1975), Grew (1982) and Tingey (1991). Since then, the accumulation of precise age data by thermal ionization mass spectrometry (TIMS) and sensitive high-resolution ion microprobe (SHRIMP), detailed petrological studies by electron probe micro analysis (EPMA) and intense field investigations have made it possible to describe the crustal evolution of East Antarctica in relation to the formation of Gondwana (e.g. Fitzsimons 2000; Yoshida et al. 2003; Frimmel 2004; Jacobs & Thomas 2004). In particular, studies on Dronning Maud Land have focused on continental plate reconstructions, and this region is now widely regarded to be the southern continuation of the late Neoproterozoic to early Palaeozoic East African Orogen (e.g. Jacobs et al. 1998, 2003a, b; Jacobs & Thomas 2002; Grantham et al. 2003; Paulsson & Austrheim 2003). In eastern Dronning Maud Land and adjacent areas, there are several discrete, isolated exposures

of magmatic and high-grade metamorphic terranes. These are, from west (c. 208E) to east (c. 508E), the Sør Rondane Mountains (SRM), the Yamato – Belgica Complex (YBC), the Lu¨tzow-Holm Complex (LHC), the Rayner Complex (RC) and the Napier Complex (NC) (Fig. 1). Hiroi et al. (1991) compiled the geology and petrology of the Lu¨tzow-Holm Complex, which extends along the eastern coastline of Lu¨tzow-Holm Bay and the Prince Olav Coast, and is characterized by a continuous increase in metamorphic grade and a clockwise prograde P – T path. The Yamato – Belgica Complex is characterized by widespread igneous activity and low-P/high-T type metamorphism, for which Hiroi et al. (1991) proposed a tectonic scenario involving continent – continent collision. Petrological studies on the LHC revealed ultrahigh-temperature metamorphism in the southern part of the complex (Motoyoshi & Ishikawa 1997). Shiraishi et al. (1992, 1994, 2003) reported c. 550–530 Ma SHRIMP zircon U –Pb ages for peak metamorphism in the LHC and suggested that the collision took place in the last stage of Gondwana construction during the Pan-African Orogeny (Fig. 2). Since then, Pan-African events (c. 600–500 Ma) involving extensive plutonic activity and high-grade metamorphism have been recognized in many parts of East Antarctica. Whether there was a single Grenville-aged Circum-East Antarctic mobile belt that was tectonically reactivated in Pan-African times (Yoshida et al. 2003) or a Pan-African

Fig. 1. Map of geological terranes in central–eastern Dronning Maud Land to Enderby Land, East Antarctica. The areas of Figures 2 and 3 are marked. NE, NE Terrane; SW, SW Terrane; wRC, Western Rayner Complex.

AGE CONSTRAINTS IN EAST ANTARCTICA

23

Fig. 2. Map of SHRIMP zircon U– Pb ages for rocks from the Lu¨tzow-Holm Complex. i, igneous age; m, metamorphic age (ages in million years). *Ongul Is. includes East Ongul Island, Nesøya, Fleynøya and Utholmen. Modified from Shiraishi et al. (2003).

juxtaposition and assembly of three unrelated Grenville-aged terranes into East Gondwana (Fitzsimons 2000), is a major point of contention. In this context, the correlation of the Pan-African Mozambique Suture with geological features in the SRM to the west of the LHC is a major focus of current research. Recent comprehensive compilations of geological history of East Antarctica reveal the essential involvement of East Antarctica in the tectonic development of Gondwana (e.g. Fitzsimons 2000; Jacobs & Thomas 2004; Meert & Lieberman 2008). In particular, studies indicate that eastern DML is at a critical location in terms of the intersection of the Mozambique belt and various Pan-African terranes from the Prince Charles Mountains in Mac Robertson Land to the Pinjara Orogen in Western Australia (e.g. Jacobs et al. 1998; Fitzsimons 2000; Grantham 2003; Jacobs & Thomas 2004; Bisnath et al. 2006; Mikhalsky et al. 2006; Meert & Lieberman 2007). Boger et al. (2001) proposed an early Palaeozoic orogenic belt extending from Western Australia to Mac

Robertson Land and suggested possible links with the Lu¨tzow-Holm Complex. A similar but more extensive orogenic belt, penetrating through the LHC and across central Africa, was proposed by Meert (2001, 2003), who combined geochronological and palaeomagnetic data to develop a polyphase model for eastern Gondwana assembly. In recent years, two major models have emerged for Neoproterozoic –early Palaeozoic tectonism involved in the formation of East Gondwana: (1) Pan-African structures are coeval over long distances, forming a broad linear orogen from East Antarctica to southern East Africa (East Africa– Antarctic Orogen (EAAO); e.g. Jacobs & Thomas 2002, 2004); (2) two overlapping orogens were involved in the amalgamation of East Gondwana: an older East Africa Orogeny (Stern 1994) and a younger Kuunga Orogeny (Meert 2003). Both models identify two stages of Pan-African orogenesis, at 750–620 Ma and 570–530 Ma (Jacobs et al. 2003b; Meert 2003; Grantham et al. 2008). Although comprehensive geochronological studies have been published on eastern DML, the

24

K. SHIRAISHI ET AL.

timing of key magmatic, deformational and metamorphic events has not been sufficiently resolved to place these events in a tectonic framework related to the formation of Gondwana. Pan-African v. Grenvillian orogenesis is uncertain. In this paper we present zircon and titanite SHRIMP U –Pb ages for 11 metamorphic and magmatic rocks from the Sør Rondane Mountains, the largest mountain range in eastern DML. In addition, new Nd model ages, mainly from the LHC, are compared with published ages from central–eastern Dronning Maud Land to Enderby Land. Geochronological data are examined in conjunction with geological and petrological studies of the respective terranes. The results provide constraints on the Neoproterozoic to Cambrian crustal history of eastern Dronning Maud Land, and present insights into a region that played a critical role in the formation of the Gondwanan supercontinent.

Geological outline of the Sør Rondane Mountains The Sør Rondane Mountains (SRM), located between 228 and 288E and 71.58 and 72.58S, form one of the largest mountain chains in East Antarctica. The SRM are underlain by medium- to highgrade metamorphic rocks with various intrusions of plutonic rocks and minor mafic dykes (e.g. Shiraishi et al. 1991, 1997a). The mountains are divided into the NE and SW Terranes by an inferred tectonic line, the Sør Rondane Suture (SRS: Osanai et al. 1992). The NW Terrane is mainly composed of granulite-facies metamorphic rocks of pelitic, psammitic and intermediate compositions, whereas the SW Terrane is composed of amphibolite-facies and lower-grade metamorphic rocks of mainly intermediate to basic composition, including a large volume of meta-tonalite (Fig. 3). Plutonic rocks and dykes of various sizes intrude metamorphic lithologies in both terranes, and consist of syn- to post-orogenic granite, syenite, diorite and alkaline mafic dykes. Metamorphosed basic and intermediate igneous rocks in the central part of the mountains indicate that magmatic protoliths have geochemical affinities with oceanic, island arc, accretionary complex and continental margin arc settings in modern plate-tectonic systems (Osanai et al. 1992). Ishizuka et al. (1996) reported that ultramafic gneiss in the NE Terrane evolved from a midocean ridge basalt (MORB)-like magma source. Meta-tonalite from the SW Terrane is characterized by relatively high Na2O/K2O, K/Rb, Sr/Y and (La/Yb)N, low CaO/Na2O and low initial Sr isotopic ratios, which have been attributed to magma

genesis from a hot subducting plate (Ikeda & Shiraishi 1998). Metamorphic P–T conditions of gneisses from the NE Terrane have been estimated by several workers (Asami & Shiraishi 1987; Shiraishi & Kojima 1987; Grew et al. 1989; Asami et al. 1990, 1992, 2007; Ishizuka et al. 1995). The NE Terrane was metamorphosed under granulite-facies conditions (800 8C and 7–8 kbar), with subsequent amphibolite-facies retrogression (530 –580 8C and 5.5 kbar) within the kyanite stability field. Subsequently, the rocks were recrystallized at lower pressures, as indicated by the presence of andalusite in certain lithologies (Asami & Shiraishi 1987; Asami et al. 1992, 1993). Recently, Asami et al. (2007) reported relict sapphirine þ kyanite and spinel þ kyanite associations in garnet from the eastern SRM; such assemblages have also been found in granulite-facies rocks from the LHC and Sri Lanka, which are located close to the SRM in the reconstruction of Gondwana (Asami et al. 2007). From these assemblages, the peak conditions of granulite-facies metamorphism was estimated to be 860–895 8C and 12 kbar based on the sapphirine–spinel thermometer and stability of kyanite, although the stability range of the kyanite þ spinel assemblage is not well defined (Asami et al. 2007). In contrast, such high-T metamorphism has not been recognized in the SW Terrane, where gneisses at Vengen ridge and northern Walnumfjelle experienced metamorphism estimated at upper amphibolite-facies conditions, followed by pervasive retrograde metamorphism in association with mylonitization (Shiraishi & Kojima 1987). Since pioneer studies in the 1960s, numerous Rb –Sr and K –Ar geochronological studies in both terranes of the SRM have yielded ages from c. 500 to 420 Ma, attributed to an intense thermal event associated with plutonic activity in the early Palaeozoic (Picciotto et al. 1964; Takahashi et al. 1990; Grew et al. 1992; Shiraishi & Kagami 1992; Shiraishi et al. 1997a). Older mineral and whole-rock isochron ages have also been reported. Shiraishi & Kagami (1992) considered the age of granulite-facies metamorphism in the NE Terrane to be c. 1000 Ma, on the basis of Sm– Nd and Rb –Sr whole-rock isochron ages from orthogneisses. In contrast, internal mineral isochrons from orthogneisses and paragneisses yielded 556 Ma and 624 Ma ages for Rb–Sr and Sm–Nd systems, respectively. These younger ages were attributed to a series of thermal events associated with granitic intrusions. On the basis of Nd model ages, it was also observed that the timing for Grenvillian (c. 1000 Ma) granulite-facies metamorphism requires a short time interval between crustal formation and orogenesis (Grew et al. 1992; Shiraishi & Kagami 1992).


Fig. 3. Simplified geological map and sample localities of the Sør Rondane Mountains. SRS, Sør Rondane Suture (Osanai et al. 1992). 25

26

K. SHIRAISHI ET AL.

A radically different interpretation was provided by Asami et al. (1996, 1997, 2005), on the basis of electron microprobe dating of monazite and zircon in granulites from the SRM as well as the LHC, RC and NC. Monazite belonging to granulite-grade metamorphic assemblages, including grains enclosed in garnet porphyroblasts, yielded chemical Th – U – total Pb isochron method (CHIME) ages that mostly range from 550 to 510 Ma. Asami et al. concluded that granulite-facies metamorphism and high-strain deformation took place at 540 – 530 Ma, in a Cambrian mono-metamorphic belt extending from 258E to 458E. It was also inferred that sediments that formed the protoliths of paragneisses were deposited during the Neoproterozoic. In this scenario, c. 1000 Ma ages must date the formation of protoliths, of either igneous or high-grade metamorphic lithologies. Although the monazite dating establishes the widespread significance of high-grade metamorphism during Pan-African orogenesis, it does not clearly address the possibility of multiple metamorphic events, the origins of metamorphic protoliths, or the relationship of deformational structures and magmatic intrusions to the geochronological data. There is also no clear explanation of the differences in metamorphic grade and lithological types between the NE and SW Terranes of the SRM. Sub-grain analysis of zircon by ion microprobe was required to resolve these issues.

SHRIMP U– Pb geochronology of the Sør Rondane Mountains Samples and analytical procedures The present study compiles SHRIMP U – Pb zircon and titanite analyses for 11 samples, obtained during 13 analytical sessions at the SHRIMP II facilities at the Australian National University, Canberra (ANU; sessions A1 – 7) and the National Institute of Polar Research, Tokyo (NIPR; sessions N1 – 6). Data reduction and processing was performed using the Excel add-in program SQUID (v.1.12a; Ludwig 2001) and plots were generated using ISOPLOT (v.3.50; Ludwig 2003). For zircon analysis, abundance of U was calibrated against standard SL13 (238 ppm), and U – Pb measurements were calibrated against 204 Pb-corrected (Pb/U)/(UO/U)2 values for standard AS3 (1099 Ma, Paces & Miller 1993). For each standard dataset, scatter on (Pb/U)/ (UO/U)2 ratios and external spot-to-spot errors are quoted with data from each sample and session in Tables 1 – 14.

All measurements on zircon were corrected for common Pb content using measured 204Pb and a Stacey & Kramers (1975) model for ages approximating those of standard and unknown zircon ages (see Ludwig 2001 for details). The procedure for titanite U – Pb calibration against standard KHAN (700 ppm U and 518 Ma, Kinny et al. 1994) was identical, except that the 207Pb correction for common Pb was used, which assumes concordance between radiogenic 206Pb/238U and 207 Pb/235U ages. Wherever possible, pooled ages were calculated from single analytical sessions using the concordia age function of SQUID, which has the advantage of providing a test of concordance between pooled 206Pb/238U and 207 Pb/206Pb ages. Mean 206Pb/238U ages for pooled data are also provided in Tera – Wasserberg plots. Errors on single spot ratios and ages are quoted at 1s, whereas pooled ages and concordia intercept ages are quoted at 95% confidence levels. Concordia ages are always calculated separately from single sessions with a standard calibration. In some samples, concordia ages for the same generation of zircon growth were obtained in duplicate sessions, to assess the reproducibility of results between the SHRIMP II facilities at ANU and NIPR. Where pooled ages were obtained using data from multiple sessions, mean 206Pb/238U ages were calculated incorporating errors from standard reproducibility in each session. Samples selected for analyses from the NE Terrane comprise four pelitic to semipelitic paragneisses, two hornblende – biotite gneisses (of probable volcaniclastic origin), one enderbitic orthogneiss and one granitic dyke. Samples from the SE Terrane comprise one granitic orthogneiss, one mylonitized granite and one tonalitic orthogneiss. Localities are shown in Figure 3. Zircon and titanite grains were separated from each sample, mounted in epoxy, polished and coated in high-purity gold using an evaporative coater. Sub-grain ion beam analysis of zircon with complex internal zoning requires careful attention to spot positioning, so backscattered electron (BSE) and cathodoluminescence (CL) imaging was performed with a scanning electron microscope before and after analysis, to identify spot positions overlapping multiple growth zones, grain edges, cracks or damaged zircon. Data from analyses or sessions with analytical problems (such as ion beam instability) are not included in the tables. On Tera – Wasserburg plots, data from spots overlapping multiple growth zones are marked with open error ellipses (68.3% confidence level), and analyses on cracked or damaged zircon or with strongly discordant isotopic ratios are marked with error crosses (1s).

Table 1. SHRIMP U–Pb data for zircon from sample 85020401C Spot

232

Th/ U

% 206Pb* Pbc (ppm)

+%

173 128 67

8.299 5.921 10.836

0.08 0.01 0.09 0.03 0.06 0.10 0.04 0.01 0.07

73 185 111 96 27 213 102 158 111

0.30 0.64 0.24 0.07 0.18 0.17 0.03 0.11 0.04

78 52 115 167 132 125 127 240 104

Th (ppm)

Session A2 1.1 2.1 3.1

1667 879 849

13 75 14

0.01 0.09 0.02

0.03 0.01 0.03

Session A3 4.1 5.1 6.1 7.1 8.1 9.1 10.1 11.1 12.1

865 1681 1079 1106 212 2327 1152 1472 1323

39 11 28 44 64 9 11 16 17

0.05 0.01 0.03 0.04 0.30 0.00 0.01 0.01 0.01

Session N3 13.1 13.2 14.1 15.1 16.1 17.1 18.1 19.1 20.1

1002 430 1189 1425 1772 1609 1236 2239 1281

12 50 8 97 29 21 44 9 16

0.01 0.12 0.01 0.07 0.02 0.01 0.04 0.00 0.01

238

206

Total Pb/ 206 Pb

+%

1.60 1.05 1.05

0.06297 0.07367 0.05905

10.187 7.791 8.325 9.867 6.807 9.369 9.732 8.003 10.279

1.32 1.86 2.37 1.39 1.75 1.59 2.06 2.75 1.04

10.997 7.063 8.912 7.313 11.525 11.066 8.340 8.020 10.560

0.58 0.88 1 0.56 0.79 0.89 1.00 1.50 0.60

238

238

Pb/ U age

1s error

207

Pb/ Pb age

1s error

% Discordant

0.06275 0.07361 0.05880

2.57 733.2 0.43 1006.0 0.74 568.9

11.1 9.7 5.7

699.8 1030.9 559.6

54.6 8.8 16.0

25 2 22

Inner rim Detrital Outer rim

1.33 1.86 2.37 1.39 1.75 1.59 2.06 2.75 1.04

0.05931 0.06683 0.06533 0.05999 0.07039 0.06037 0.05978 0.06713 0.06053

0.64 1.45 1.81 0.75 1.88 1.19 1.38 1.60 1.07

603.2 778.4 730.7 622.1 883.2 653.3 630.3 759.0 598.0

7.7 13.6 16.4 8.2 14.5 9.9 12.3 19.7 5.9

578.6 832.5 785.1 603.0 939.6 617.0 595.8 841.9 622.6

13.9 30.3 38.1 16.2 38.5 25.8 29.9 33.4 23.0

24 6 7 23 6 26 26 10 4

Outer rim Inner rim Inner rim Outer rim Mixed Mixed Outer rim Mixed Outer rim

0.59 0.91 1 0.57 0.79 0.90 1.00 1.50 0.60

0.05800 0.06490 0.06310 0.07040 0.05867 0.05903 0.06494 0.06500 0.06040

2.00 3.60 2.40 0.98 1.30 1.50 1.30 2.10 1.20

559.5 848.5 684.0 825.7 535.5 556.8 729.8 757.1 583.0

3.2 7.2 6.5 4.4 4.1 4.8 7.0 10.5 3.4

528.1 772.0 712.9 940.0 554.7 568.3 772.4 773.6 617.8

44.2 75.3 51.0 20.1 27.8 33.2 27.2 44.4 25.0

26 29 4 14 4 2 6 2 6

Outer rim Detrital crack Mixed Mixed Sector z rim Sector z rim Inner rim Rim Sector z core

U/ Pb*

+%

2.55 0.43 0.64

8.302 5.921 10.839

1.60 1.05 1.05

0.05993 0.06689 0.06606 0.06025 0.07085 0.06102 0.06010 0.06722 0.06113

0.56 1.45 1.77 0.68 1.76 1.18 1.35 1.60 1.01

10.194 7.792 8.332 9.870 6.811 9.376 9.736 8.003 10.287

0.06041 0.07020 0.06510 0.07094 0.06010 0.06038 0.06520 0.06590 0.06070

1.30 1.50 2.10 0.75 0.91 1.20 1.20 2.00 1.00

11.030 7.109 8.934 7.318 11.545 11.084 8.343 8.020 10.564

207

206

207

Pb*/ Pb*

+%

206

206

238

206

Notes


Total U/ 206 Pb

U (ppm)

(Continued )

27

28

Table 1. Continued Spot

Th (ppm)

235 440 1733 675 525 1241 967 645 297 937 293 258 1184 689 245 1581

101 181 26 74 78 25 16 59 130 16 27 34 22 65 38 9

232

Th/ U

238

0.44 0.42 0.02 0.11 0.15 0.02 0.02 0.09 0.45 0.02 0.10 0.14 0.02 0.10 0.16 0.01

% 206Pb* Pbc (ppm)

206

0.11 0.00 0.30 0.02 0.06 0.01 0.04 0.09 0.03 0.09 0.02 0.16 0.60 0.07 0.71

33 65 193 100 76 106 77 74 44 75 30 39 100 85 35 136

Total U/ 206 Pb

+%

6.030 5.800 7.720 5.810 5.926 10.080 10.840 7.489 5.867 10.790 8.340 5.720 10.210 7.000 6.057 10.010

2.40 2.70 2.90 2.70 1.10 1.10 1.10 1.20 1.10 1.10 2.10 6.60 2.40 2.60 1.10 1.40

238

Total Pb/ 206 Pb

+%

0.07850 0.06960 0.06740 0.07700 0.07325 0.05937 0.05932 0.06610 0.07373 0.05820 0.06740 0.06960 0.06630 0.07715 0.07170 0.06695

4.20 3.90 5.40 2.80 0.63 0.79 0.49 2.40 0.60 2.00 4.50 5.40 2.60 0.42 2.00 1.10

207

238

U/ Pb*

+%

6.030 5.800 7.740 5.810 5.922 10.090 10.840 7.492 5.872 10.790 8.340 5.720 10.220 7.040 6.061 10.080

2.40 2.70 2.90 2.70 1.10 1.10 1.10 1.20 1.10 1.10 2.10 6.60 2.40 2.60 1.10 1.40

206

207

Pb*/ Pb*

+%

0.07760 0.06960 0.06490 0.07690 0.07371 0.05892 0.05927 0.06570 0.07297 0.05800 0.06670 0.06940 0.06490 0.07224 0.07110 0.06114

4.30 3.90 5.60 2.80 0.68 0.84 0.50 2.40 0.73 2.00 4.60 5.50 2.70 0.89 2.00 1.50

206

206

Pb/ U age

1s error

207

Pb/ Pb age

1s error

% Discordant

988.5 1026.1 783.2 1024.2 1005.8 609.2 568.8 807.7 1013.7 571.5 729.8 1038.2 601.6 855.7 984.4 609.5

22.3 25.6 21.1 25.7 10.1 6.2 5.9 9.5 10.4 5.9 14.2 63.4 13.9 20.9 10.5 8.2

1137.3 916.1 771.5 1118.0 1033.5 564.2 577.2 798.3 1013.2 530.5 827.4 909.6 771.6 992.7 960.4 644.2

85.6 79.4 117.8 55.7 13.8 18.3 10.8 51.0 14.8 43.1 96.7 112.4 57.6 18.1 41.8 33.0

15 211 21 9 3 27 1 21 0 27 13 212 28 16 22 6

238

206

Notes

Detrital Detrital Crack Detrital Detrital Outer rim Outer rim Mixed Detrital Outer rim Detrital Detrital Rim Mixed Detrital Outer rim

All errors are quoted at 1s; Pbc and Pb* indicate the common and radiogenic portions, respectively. Common Pb corrected using measured 204Pb. External spot-to-spot errors from observed scatter in standard are 0.00%, 0.00%, 0.00% and 1.01% (respective sessions, included in the calculation of sample-spot errors). Errors in standard calibration are 0.31%, 0.28%, 0.59% and 0.41% (respective sessions, not included in above errors but required when comparing data from different sessions). % Discordance: positive indicates normal (i.e. Pb–Pb age . U –Pb age); negative indicates reverse discordance.

K. SHIRAISHI ET AL.

Session N4 8.2 9.2 21.1 21.2 22.1 23.1 24.1 24.2 25.1 26.1 26.2 27.1 28.1 29.1 29.2 30.1

U (ppm)

Table 2. SHRIMP U–Pb data for zircon from sample 84022004 Spot

U Th (ppm) (ppm)

232

Th/ U

238

% Pbc

206

206

Pb* Total (ppm) 238U/ 206 Pb

+%

+%

0.06970 0.06600 0.06627 0.06539 0.06541 0.06561 0.06560 0.06591 0.07040 0.07070

238 206

U/ Pb*

+%

Pb*/ Pb*

+%

206

Session A2 1.1 2.1 3.1 4.1 5.1 6.1 7.1 8.1 9.1 10.1

130 508 452 431 292 495 196 416 113 102

56 23 220 177 129 270 70 90 42 44

0.44 0.05 0.50 0.42 0.46 0.56 0.37 0.22 0.39 0.44

0.13 0.02 0.03 0.11 0.03 0.11 0.10 0.14 0.69

14 51 54 50 33 56 23 48 13 11

7.780 8.570 7.220 7.376 7.650 7.658 7.510 7.485 7.690 8.390

1.70 1.70 0.64 0.84 0.82 0.87 3.10 0.68 1.90 2.80

7.780 8.590 7.222 7.378 7.659 7.660 7.510 7.492 7.700 8.440

1.30 1.70 1.10 1.10 1.10 1.20 1.80 1.30 1.80 4.00

Session N2 3.2 4.2 7.2 10.2 10.3 17.1 18.1 18.2 19.1 20.1

433 444 254 267 122 238 315 99 263 527

209 258 69 179 54 37 207 34 216 127

0.50 0.60 0.28 0.69 0.46 0.16 0.68 0.36 0.85 0.25

0.24 0.12 0.23 0.29 0.25 0.08 0.44 0.34 0.09 0.15

49 51 39 38 15 24 47 9 38 57

7.615 1.20 0.06510 0.93 7.520 1.50 0.06799 0.81 5.630 2.20 0.07829 0.94 6.002 1.50 0.07398 1.00 6.820 14.00 0.07600 3.60 8.640 1.30 0.06860 2.80 5.765 1.4 0.07708 0.94 9.040 3.10 0.07260 2.70 5.878 1.30 0.07569 0.88 7.944 1.20 0.06859 1.20

7.633 7.530 5.640 6.020 6.840 8.650 5.790 9.070 5.883 7.956

1.20 1.50 2.20 1.50 14.00 1.30 1.4 3.10 1.30 1.20

1.30 1.70 1.10 1.10 1.10 1.20 1.80 1.30 1.80 4.00

207

Pb/ U age

1s error

779.9 710.3 836.0 819.3 791.1 791.0 805.5 807.6 787.2 721.5 793.6 803.6 1051.7 990.7 880.0 705.1 1027.0 674.4 1012.0 763.3

238

0.07040 3.40 0.06480 1.80 0.06606 0.67 0.06514 0.87 0.06442 1.20 0.06539 0.90 0.06460 3.30 0.06503 0.80 0.06920 2.20 0.06450 4.80 0.06308 0.06701 0.07640 0.07150 0.07390 0.06790 0.07340 0.06980 0.07490 0.06739

206

1.40 1.40 1.50 1.80 4.00 3.10 1.70 3.80 1.50 1.30

207

Pb/ Pb age

1s error

% Discordant

9.6 11.7 8.9 8.6 8.5 8.9 13.9 9.9 13.0 27.4

938.7 767.2 808.3 779.0 755.3 786.9 762.0 775.2 903.8 757.4

70.5 38.5 14.1 18.2 24.4 18.9 70.2 16.9 45.2 101.1

20 8 23 25 25 21 25 24 15 5

Igneous Unz. rim Igneous Igneous Igneous Igneous Ig. rim Igneous Igneous Ig. rim

9.2 11.3 21.7 14.1 114.3 9.0 13.7 19.7 12.0 8.7

711.0 838.2 1104.4 972.4 1039.0 865.7 1024.1 922.7 1066.5 849.9

30.0 28.3 29.1 35.7 81.2 63.4 34.4 78.0 29.3 28.0

210 4 5 22 18 23 0 37 5 11

Igneous Igneous Inh. core Inh. core Ig. rim Unz. rim Inh. core Ig. rim Ig. core Ig. crack

206

Notes


Total Pb/ 206 Pb

207

All errors are quoted at 1s; Pbc and Pb* indicate the common and radiogenic portions, respectively. Common Pb corrected using measured 204Pb. External spot-to-spot errors from observed scatter in standard are 1.05% and 1.11% (respective sessions, included in the calculation of sample-spot errors). Errors in standard calibration are 0.40% and 0.44% (respective sessions, not included in above errors but required when comparing data from different sessions). % Discordance: positive indicates normal (i.e. Pb– Pb age . U –Pb age); negative indicates reverse discordance.

29

30

Table 3. SHRIMP U–Pb data for zircon from sample 85011503D Spot

200 327 239 93 201 117 256 407 113 114 97 70 82 79 547 825 241 751 281 113 95 217 175 419 169 478 165

42 65 66 16 39 22 47 134 22 22 19 17 15 15 175 281 33 299 88 21 22 69 34 124 38 121 32

232

Th/ U

238

0.21 0.20 0.28 0.17 0.20 0.19 0.18 0.33 0.19 0.20 0.19 0.23 0.19 0.19 0.32 0.34 0.14 0.40 0.31 0.19 0.23 0.32 0.19 0.30 0.23 0.25 0.19

% Pbc

206

0.15 0.01 0.10 0.19 0.02 0.32 1.00 0.25 0.06 0.04 0.09 0.24 0.23 0.10 0.07 0.22 0.03 0.09

206

Pb* (ppm)

22 49 31 13 27 11 26 67 10 11 9 7 7 7 83 127 37 111 44 10 12 27 24 62 27 54 24

Total U/ 206 Pb

238

+%

Total Pb/ 206 Pb

207

+%

238

U/ Pb*

+%

8.974 6.542 7.696 7.003 7.190 10.111 9.493 6.050 10.523 9.858 10.422 9.534 10.502 10.675 6.643 6.558 6.263 6.969 6.398 10.719 7.627 8.118 6.960 6.788 6.157 8.682 6.666

2.66 2.41 4.34 2.85 2.62 2.60 2.54 2.34 2.85 2.90 2.81 3.50 3.09 2.69 2.40 2.32 2.44 2.40 3.83 2.60 3.12 2.45 2.44 2.34 2.73 2.40 2.66

206

207

Pb*/ Pb*

+%

0.06551 0.07118 0.07312 0.07143 0.06926 0.06150 0.06140 0.07079 0.06178 0.05818 0.06191 0.05832 0.06057 0.06053 0.07091 0.07025 0.07214 0.06935 0.07226 0.05991 0.06689 0.06833 0.06546 0.07170 0.06961 0.06717 0.06725

1.54 1.24 7.60 2.21 2.06 1.71 1.58 0.86 1.44 2.94 1.53 6.60 3.42 2.89 0.62 0.73 0.89 0.52 5.31 2.65 2.87 1.93 1.16 0.78 1.72 1.15 1.44

206

206

Pb/ U age

1s error

207

Pb/ Pb age

1s error

% Discordant

Notes

681.1 916.9 787.6 860.5 839.4 608.0 645.6 986.1 585.2 622.8 590.7 643.0 586.4 577.2 903.9 914.9 954.9 864.4 936.1 575.0 794.2 748.9 865.4 885.9 970.3 702.8 901.0

17.2 20.6 32.3 23.0 20.7 15.1 15.6 21.4 16.0 17.2 15.9 21.4 17.3 14.9 20.3 19.8 21.7 19.5 33.5 14.3 23.4 17.4 19.8 19.4 24.6 16.0 22.4

790.8 962.5 1017.4 969.6 906.6 656.7 653.3 951.4 666.5 536.6 670.9 541.9 624.2 622.6 954.8 935.6 989.9 909.2 993.4 600.4 834.4 878.6 789.1 977.6 916.7 843.1 845.6

32.8 25.6 162.1 45.8 43.2 37.2 34.2 17.7 31.2 65.5 33.2 151.4 75.6 63.8 12.8 14.9 18.3 10.7 111.9 58.5 60.9 40.6 24.6 16.1 36.0 24.1 30.2

14 5 23 11 7 7 1 24 12 216 12 219 6 7 5 2 4 5 6 4 5 15 210 9 26 17 27

Unzoned Igneous Ig crack Ig crack Ig core Simple Mix Igneous Simple zoned Simple zoned Rim Rim Rim Simple zoned Igneous Igneous Igneous Ig core Igneous Rim Mixed Ig crack Mixed Igneous Igneous Mixed Igneous

238

206

All errors are quoted at 1s; Pbc and Pb* indicate the common and radiogenic portions, respectively. Common Pb corrected using measured 204Pb. % Discordance: positive indicates normal (i.e. Pb–Pb age . U–Pb age); negative indicates reverse discordance.

K. SHIRAISHI ET AL.

Session A1 1.1 2.1 2.2 3.1 3.2 4.1 5.1 5.2 6.1 7.1 8.1 9.1 10.1 11.1 12.1 13.1 14.1 14.2 15.1 15.2 16.1 17.1 18.1 18.2 19.1 19.2 20.1

U Th (ppm) (ppm)

Table 4. SHRIMP U–Pb data for zircon from sample 9032401A Spot

430 796 363 806 392 651 242 382 627 581 596 1113 234 347 95 520 543 117 185 469 720 79 267 344 213

59 31 37 27 101 52 112 208 216 242 45 10 81 10 36 39 0 48 72 339 393 46 35 123 121

232

Th/ U

238

0.14 0.04 0.11 0.03 0.27 0.08 0.48 0.56 0.36 0.43 0.08 0.01 0.36 0.03 0.39 0.08 0.00 0.42 0.40 0.75 0.56 0.60 0.14 0.37 0.59

206 Pb* Total % Pbc (ppm) 238U/ 206 Pb

+%

206

0.70 0.16 0.56 0.51 0.67 0.44 0.19 0.50 0.33 0.30 0.58 0.09 0.86 0.63 1.53 0.52 0.33 0.27 0.12 0.27 0.14 1.49 0.73 0.26 0.42

38 71 33 73 35 57 22 62 91 95 53 193 21 30 13 47 39 67 65 79 118 11 26 45 39

9.610 1.10 9.645 1.00 9.570 1.20 9.509 10.00 9.580 1.20 9.740 1.00 9.490 1.30 5.260 2.30 5.897 1.30 5.260 2.40 9.600 1.10 4.965 1.30 9.630 1.40 10.000 1.20 6.210 1.60 9.530 1.10 11.930 1.10 1.514 1.40 2.464 1.30 5.105 1.90 5.254 0.98 6.010 1.70 8.780 1.20 6.549 1.10 4.736 1.90

Total Pb/ 206 Pb

207

0.06636 0.06400 0.06595 0.06215 0.06440 0.06407 0.06530 0.08076 0.07837 0.07835 0.06516 0.08009 0.06700 0.06490 0.08290 0.06350 0.06034 0.23750 0.22530 0.07722 0.07632 0.08900 0.06727 0.07547 0.08270

+%

238 206

U/ Pb*

+%

1.40 9.680 1.10 0.94 9.661 1.00 1.40 9.620 1.20 1.20 9.558 1.00 2.50 9.640 1.20 1.10 9.790 1.00 1.70 9.500 1.30 0.92 5.290 2.30 0.79 5.917 1.30 0.76 5.280 2.40 1.30 9.660 1.10 0.52 4.969 1.30 1.70 9.710 1.40 1.60 10.070 1.30 1.90 6.310 1.70 1.10 9.580 1.10 1.30 11.970 1.10 1.00 1.518 1.40 1.10 2.467 1.30 0.89 5.118 1.90 0.66 5.261 0.98 2.00 6.100 1.80 1.50 8.840 1.20 1.10 6.566 1.10 1.50 4.756 1.90

207

Pb*/ Pb*

+%

0.06060 0.06265 0.06130 0.05800 0.05890 0.06050 0.06380 0.07660 0.07560 0.07580 0.06040 0.07931 0.05990 0.05970 0.07020 0.05920 0.05760 0.23510 0.22430 0.07490 0.07518 0.07660 0.06130 0.07330 0.07910

3.40 1.30 3.20 2.10 5.50 2.20 1.80 1.60 1.40 1.40 2.40 0.60 6.00 3.70 6.30 2.90 2.70 1.10 1.10 1.40 0.98 7.60 3.40 1.70 2.30

206

206

Pb/ U age

1s error

207

633.6 634.9 637.3 641.5 636.3 627.2 644.9 1116.0 1006.6 1118.0 635.0 1181.9 631.6 610.5 948.5 640.0 517.3 3261.6 2193.6 1150.4 1121.8 979.2 690.8 913.9 1230.3

6.9 6.1 7.2 6.1 7.6 6.2 7.9 23.3 12.4 24.9 6.4 13.9 8.6 7.3 14.9 6.6 5.4 35.5 24.5 19.7 10.1 16.8 8.1 9.5 21.3

624.6 696.5 650.5 528.5 563.6 621.6 733.5 1109.9 1084.4 1090.5 619.5 1180.0 601.3 593.6 933.4 575.5 515.6 3086.9 3012.1 1067.1 1073.3 1111.7 648.2 1023.3 1175.6

238

Pb/ Pb age

206

1s % error Discordant

73.2 26.8 69.4 46.5 120.2 46.8 39.2 32.4 27.4 27.1 50.7 11.9 128.9 79.2 129.1 62.2 60.2 17.9 18.4 28.1 19.7 151.2 72.5 34.4 45.3

21 10 2 218 211 21 14 21 8 22 22 0 25 23 22 210 0 25 37 27 24 14 26 12 24

Notes

Unzoned Unzoned rim Unzoned rim Core Rim Rim Rim Inherited core Igneous rim Igneous core Sector z rim Core Zoned rim Mixed Ig. crack Core Rim Igneous core Igneous rim Igneous Igneous Ig crack Mixed Igneous core Igneous core


Session A4 1.1 2.1 3.1 4.1 4.2 5.1 5.2 6.1 6.2 7.1 7.2 8.1 8.2 10.1 11.1 12.1 12.2 14.1 14.2 15.1 16.1 17.1 18.1 19.1 20.1

U Th (ppm) (ppm)

All errors are quoted at 1s; Pbc and Pb* indicate the common and radiogenic portions, respectively. Common Pb corrected using measured 204Pb. External spot-to-spot error from observed scatter in standard is 0.83% (included in the calculation of sample-spot errors). Error in standard calibration is 0.33% (not included in above errors but required when comparing data from different sessions). % Discordance: positive indicates normal (i.e. Pb–Pb age . U –Pb age); negative indicates reverse discordance.

31

32


U Th (ppm) (ppm)

232

Th/ U

238

206 % Pb* Pbc (ppm)

206

+% Total U/206Pb

238

Total +% Pb/206Pb

207

238

U/ +% Pb*

206

207

Pb*/ Pb*

+%

206

206

1s Pb/ U age error

238

207 206

1s % Pb/ Pb age error Discordant

Notes

190 30 802 72 82 115 114

141 10 78 34 26 56 40

0.74 0.33 0.10 0.47 0.32 0.49 0.35

1.39 1.85 0.14 1.21 1.78 0.83 1.49

32 5 103 12 12 17 20

5.214 5.204 6.011 5.191 5.393 5.585 4.584

2.38 3.67 1.42 2.91 2.47 3.20 5.07

0.08152 0.09703 0.07863 0.08794 0.08143 0.08428 0.08977

2.34 2.82 0.78 2.41 1.60 2.38 4.39

5.288 5.302 6.020 5.255 5.491 5.632 4.653

2.39 3.71 1.42 2.96 2.49 3.22 5.23

0.06968 4.43 0.08140 6.41 0.07746 0.90 0.07773 7.06 0.06613 5.38 0.07724 4.23 0.07720 15.84

1116.6 1113.8 990.7 1123.1 1078.5 1053.6 1255.0

24.6 38.0 13.0 30.6 24.8 31.3 59.8

918.8 1231.2 1133.1 1140.1 810.5 1127.4 1126.3

94.0 131.3 18.2 147.3 116.8 86.7 353.0

222 10 13 2 233 7 211

Igneous Igneous Rim Igneous Igneous Igneous Igneous

Session N1 5b.2 8.1 9.1 9.2 10.1 11.1 12.1 12.2 13.1 13.1b 13.2 14.1 15.1 15.2

318 157 566 161 326 433 562 236 441 346 148 528 2436 144

73 3 57 75 245 76 50 139 96 59 72 247 165 80

0.24 0.02 0.10 0.49 0.78 0.18 0.09 0.61 0.22 0.18 0.50 0.48 0.07 0.57

0.36 0.49 0.14 0.66 0.45 0.31 0.05 0.31 0.58 0.43 0.95 0.26 0.12 1.39

54 25 87 22 51 72 93 39 53 39 26 91 335 24

5.090 5.430 5.613 6.321 5.520 5.148 5.214 5.239 7.184 7.681 4.930 4.988 6.251 5.096

1.20 1.30 1.30 1.40 1.20 1.10 1.20 1.50 1.20 1.20 1.30 1.10 1.20 1.30

0.07914 0.07904 0.07745 0.07775 0.08079 0.07862 0.07768 0.08179 0.07790 0.07613 0.08669 0.08076 0.07632 0.08880

0.76 1.10 0.55 1.20 0.72 0.64 0.82 0.86 0.75 1.10 1.00 0.53 0.31 1.70

5.109 5.457 5.621 6.363 5.545 5.163 5.216 5.256 7.226 7.714 4.978 5.001 6.258 5.168

1.20 1.30 1.30 1.40 1.20 1.10 1.20 1.50 1.20 1.20 1.30 1.10 1.20 1.40

0.07610 0.07490 0.07627 0.07220 0.07700 0.07603 0.07728 0.07920 0.07310 0.07260 0.07860 0.07857 0.07529 0.07710

1152.3 1084.6 1055.5 941.0 1068.9 1141.2 1130.6 1122.8 835.5 785.8 1180.1 1175.2 955.6 1140.3

12.9 13.1 12.6 12.1 11.5 11.9 12.2 15.6 9.2 9.1 14.3 12.0 10.6 14.1

1096.5 1065.2 1102.1 992.5 1122.1 1095.8 1128.4 1177.2 1017.5 1002.7 1163.0 1161.4 1076.4 1123.1

40.2 46.1 16.2 57.8 26.3 18.9 19.8 30.1 30.4 46.1 44.0 15.3 8.5 73.8

25 22 4 5 5 24 0 5 22 28 21 21 13 22

Ig. rim Rim Rim Ig. core Ig. rim Rim Rim Ig. core Rim Rim Ig. core Ig. rim Rim Ig. Core

2.00 2.30 0.81 2.80 1.30 0.95 0.99 1.50 1.50 2.30 2.20 0.77 0.42 3.70


K. SHIRAISHI ET AL.

Session A4 1.1 2.1 3.2 4.1 4.2 5.1 6.1

29 56 5 14 4 6 9 4 19 5 Session A7 1.1 191 2.1 71 3.1 10 4.1 22 5.1 17 6.1 10 7.1 13 8.1 9 9.1 116 10.1 12

All errors are quoted at 1s; Pbc and Pb* indicate the common and radiogenic portions, respectively. Common Pb corrected assuming 206Pb/238U – 207Pb/235U concordance. External spot-to-spot error from scatter in standard is 0.93% (included in the calculation of sample-spot errors). Error in standard calibration is 0.38% (not included in errors but required when comparing data from different sessions).

Crack

Crack

6.4 8.0 19.1 14.0 13.9 17.6 15.4 18.5 6.7 15.6 511.8 520.8 443.7 538.2 523.3 461.8 511.0 485.7 520.5 495.0 1.1 1.6 3.8 5.4 2.2 2.7 2.3 3.0 1.4 2.5 0.0658 0.0883 0.2033 0.1680 0.1535 0.1903 0.1726 0.1898 0.0727 0.1774 1.3 1.6 4.2 2.3 2.7 3.8 3.0 3.8 1.3 3.1 11.98 11.44 11.48 9.94 10.43 11.24 10.41 10.69 11.67 10.67 13.7 5.3 0.7 1.9 1.4 0.7 1.1 0.7 8.5 1.0 1.01 3.75 18.21 13.48 11.76 16.54 14.16 16.38 1.83 14.81 0.16 0.81 0.52 0.69 0.24 0.65 0.69 0.43 0.17 0.39

Pb* (ppm)

206

% Pbc 206

Th/238U 232

Th (ppm) U (ppm) Spot

Table 6. SHRIMP U– Pb data for titanite from sample 90102801A

238

Total U/206Pb

+%

207

Total Pb/206Pb

+%

206

Pb*/238U age

1s error

Notes


33

NE Terrane Sample 85020401C. This is a migmatitic sillimanite – garnet – biotite paragneiss from the northernmost peak of Perlebandet. The gneiss is intercalated with marble and biotite gneiss showing strong deformation with shear fractures. Garnet porphyroblasts are cracked and filled with secondary muscovite. Monazite, apatite and opaque minerals are also found. Asami et al. (2005) dated monazites from this sample by CHIME (see Discussion). Zircon grains have diverse morphologies and complex internal structures (Fig. 4a). The majority are elongate and rounded, with prismatic and oscillatory-zoned cores enclosed in thin rims of low-CL zircon. Fewer grains are ovoid or equant, with broader low-CL rims that are unzoned or have round, concentric growth zones. Several grains have additional thin outer rims of slightly higher CL, with irregular growth and sector zones, that can be distinguished from inner rims of low-CL zircon (Fig. 4a). The structural complexity required many analyses, and a total of 37 analyses from 30 grains (Table 1) were obtained. From these, seven analyses of spots that covered a mixture of zircon cores and rims (open ellipses, Fig. 5a), and analyses 13.2 and 21.1 on cracked zircon (crosses, Fig. 5a), were excluded from interpretation. Prismatically zoned cores have U contents between 230 and 880 ppm and Th/U ratios between 0.1 and 0.5. Excluding discordant analysis 26.2, seven data from seven cores define a concordia age of 1009 +13 Ma (MSWD ¼ 1.4, probability of concordance ¼ 0.2). The age is of detrital zircon from an igneous source, probably of local derivation (see Discussion). Zircon rims are characterized by high U (850 – 2320 ppm) and very low Th/U (,0.05). Concordant rim ages (from multiple sessions) fall into three discrete populations (Fig. 5a), with 206Pb/238U mean ages of 736 +13 Ma (five data, MSWD ¼ 0.6), 609 +11 Ma (six data, MSWD ¼ 1.3) and 565 + 7 Ma (five data, MSWD ¼ 0.7). Zircon textures or compositions in each age group cannot be distinguished, but are characteristic of zircon growth during granulite-grade metamorphism. In this sample alone, it is uncertain if each age group represents a discrete metamorphic event. It is also not obvious whether the c. 750 Ma age group represents in situ metamorphic growth on detrital (c. 1009 Ma) zircon grains in the metasedimentary host, or if the c. 750 Ma rims are also detrital. Results from other samples (below) help clarify the interpretation of these rim ages. Sample 84022004. This is a foliated medium-grained garnet–biotite paragneiss from Utnibba Nunatak.

34


232

Th/ U

238

206 % Pb* Pbc (ppm)

206

1920 212 177 1176 179 258 390 1277 852 230 343 357 1224 1063 176 385

21 17 11 50 129 5 5 18 36 9 8 6 13 11 0 7

0.01 0.08 0.07 0.04 0.75 0.02 0.01 0.01 0.04 0.04 0.02 0.02 0.01 0.01 0.00 0.02

0.03 0.30 1.18 3.33 0.72 0.09 0.40 2.48 0.03 0.01 0.08 0.60 0.03 0.26 0.01

159 18 15 104 21 21 33 108 70 20 29 30 137 90 15 31

1057 373 451 636 769 286 496 1309

9 8 8 12 18 6 7 28

0.01 0.02 0.02 0.02 0.02 0.02 0.02 0.02

0.06 0.08 0.85 0.06 0.00 0.05 0.02 1.27

85 28 36 53 65 22 39 131

Total U/ 206 Pb

+%

10.350 10.170 10.090 9.700 7.450 10.370 10.110 10.170 10.530 10.120 10.100 10.410 7.700 10.180 9.890 10.830

1.10 1.20 1.20 1.20 2.90 1.20 1.10 1.30 1.10 1.20 1.20 1.10 1.40 1.10 1.30 1.20

10.660 11.570 10.790 10.340 10.240 11.340 11.040 8.590

2.90 3.00 4.90 1.30 4.20 1.40 4.20 3.00

238

+%

238

U/ Pb*

+%

0.05942 0.40 0.06116 1.20 0.06040 3.50 0.07190 2.70 0.09430 3.00 0.06668 1.10 0.06076 1.50 0.06260 0.59 0.07990 4.30 0.06144 1.10 0.06065 0.97 0.05954 1.00 0.07320 1.70 0.05970 0.56 0.06050 3.00 0.06154 10.00

10.360 10.210 10.090 9.820 7.710 10.440 10.120 10.210 10.800 10.130 10.100 10.420 7.750 10.190 9.910 10.830

0.06130 0.05871 0.06753 0.05975 0.05783 0.05858 0.05868 0.07660

10.670 11.580 10.880 10.340 10.240 11.340 11.040 8.700

Total Pb/ 206 Pb

207

206

1.70 1.30 0.77 1.10 1.30 1.10 1.40 3.50

207

206

207

Pb*/ Pb*

+%

1.10 1.20 1.20 1.20 3.00 1.20 1.10 1.30 1.20 1.20 1.20 1.10 1.40 1.10 1.30 1.20

0.05915 0.05850 0.06050 0.06130 0.06450 0.06020 0.05998 0.05896 0.05750 0.06120 0.06059 0.05882 0.06780 0.05941 0.05820 0.06147

0.47 1.80 3.50 3.80 8.20 2.40 1.60 0.99 8.10 1.90 1.00 1.20 2.10 0.63 3.80 1.20

594.2 602.6 609.2 625.4 786.2 589.6 607.5 602.2 571.0 607.1 608.5 591.0 782.4 603.7 619.5 569.5

6.1 6.9 7.1 7.2 22.0 6.7 6.6 7.2 6.3 6.9 6.7 6.5 10.2 6.3 7.7 6.3

572.8 10.2 546.7 38.5 621.7 74.6 651.5 82.2 756.5 173.6 611.2 52.7 602.8 34.2 565.6 21.6 511.5 177.9 645.7 39.8 624.8 21.5 560.6 27.0 862.3 43.9 582.2 13.7 536.4 84.0 655.8 24.8

24 29 2 4 24 4 21 26 210 6 3 25 10 24 213 15

Rim Sector zoned Sector zoned Mixed, crack Inherited Sector zoned Sector zoned Rim Rim, crack Sector zoned Zoned rim Sector zoned Mixed, crack Rim Sector zoned Zoned rim

2.90 3.00 4.90 1.30 4.20 1.40 4.20 3.00

0.06080 0.05803 0.06060 0.05922 0.05781 0.05820 0.05856 0.06620

1.80 1.50 1.70 1.20 1.30 1.30 1.50 4.30

578.0 534.0 567.0 594.9 601.0 544.7 559.0 702.0

16.0 15.0 27.0 7.2 24.0 7.5 22.0 20.0

631.0 531.0 626.0 575.0 523.0 537.0 551.0 812.0

9 21 10 23 213 21 21 16

Rim Sector z, crack Sector zoned Sector zoned Zoned core Zoned rim Sector zoned Mixed, crack

206

Pb/ 1s U error age

238

Pb/ Pb age

206

1s error

38.0 32.0 37.0 25.0 29.0 28.0 32.0 89.0

% Discordant

Notes

K. SHIRAISHI ET AL.

Session A2 1.1 2.1 3.1 4.1 4.2 5.1 6.1 7.1 8.1 9.1 10.1 11.1 12.1 13.1 14.1 15.1 Session N5 1.2 3.2 5.2 6.2 10.2 10.3 11.2 12.2

U Th (ppm) (ppm)

436 3904 213 357 611 704 289 436 635 132 587 180 343 671 715 320 187 843 963

80 99 72 7 10 11 16 6 9 4 6 19 9 8 9 12 155 99 8

0.19 0.03 0.35 0.02 0.02 0.02 0.06 0.01 0.02 0.03 0.01 0.11 0.03 0.01 0.01 0.04 0.85 0.12 0.01

5.92 0.01 0.36 0.64 0.00 0.32 0.11 0.02 1.59 0.01 0.06 0.06 0.80 1.15 0.16 0.07

32 316 27 29 47 55 24 36 54 8 46 15 29 55 56 24 32 78 83

11.920 10.610 6.880 10.750 11.080 11.050 10.570 10.330 10.110 13.330 10.930 10.530 10.250 10.510 10.930 11.360 4.960 9.240 10.010

1.30 5.20 4.20 1.30 1.30 1.30 1.50 1.50 3.90 2.40 1.30 1.30 3.80 2.20 5.40 1.30 8.50 1.50 2.30

0.11060 0.06052 0.07180 0.06506 0.06020 0.06206 0.06150 0.06220 0.06064 0.07710 0.05972 0.06058 0.05895 0.06037 0.06117 0.06525 0.08270 0.06539 0.06025

1.80 1.30 3.00 1.40 1.80 0.64 2.00 3.40 0.69 1.60 0.67 1.40 1.60 0.62 1.40 1.40 1.60 1.10 1.00

12.670 10.610 6.910 10.820 11.080 11.090 10.570 10.340 10.110 13.540 10.930 10.540 10.240 10.520 10.920 11.450 5.010 9.260 10.020

1.40 5.20 4.20 1.30 1.30 1.30 1.50 1.50 3.90 2.40 1.30 1.40 3.80 2.20 5.40 1.30 8.50 1.50 2.30

0.06290 0.06046 0.06880 0.05990 0.06020 0.05946 0.06210 0.06130 0.06047 0.06430 0.05967 0.06010 0.05910 0.05988 0.06128 0.05880 0.07290 0.06412 0.05965

4.70 489.8 1.30 581.0 3.60 872.0 3.10 569.8 1.80 556.9 0.99 556.7 2.00 583.0 3.50 595.1 0.74 608.0 4.10 459.0 0.80 564.2 2.40 584.4 1.60 600.0 0.71 586.0 1.40 565.0 2.10 539.7 4.80 1172.0 1.40 661.4 1.20 613.0

6.4 706.0 29.0 620.0 34.0 893.0 7.3 599.0 6.9 612.0 6.8 584.0 8.4 676.0 8.4 650.0 23.0 620.0 11.0 753.0 6.9 592.0 7.6 608.0 22.0 571.0 12.0 599.0 29.0 649.0 6.7 558.0 91.0 1011.0 9.2 746.0 13.0 591.0

99.0 28.0 75.0 68.0 39.0 21.0 43.0 76.0 16.0 86.0 17.0 51.0 34.0 15.0 30.0 46.0 97.0 29.0 27.0

44 7 2 5 10 5 16 9 2 64 5 4 25 2 15 3 214 13 24

Mixed, crack Rim Inherited Mixed, crack Sector zoned Inherited Sector zoned Sector zoned Zoned core Mixed, crack Sector zoned Sector zoned Sector zoned Sector zoned Sector zoned Sector z, crack Inh, crack Inherited Zoned core



16.1 17.1 18.1 18.2 19.1 20.1 21.1 22.1 23.1 24.1 25.1 26.1 26.2 27.1 28.1 29.1 30.1 31.1 32.1

35

36

Table 8. SHRIMP U-Pb data for zircon from sample 90112302A Spot

Th (ppm)

1034 156 317 176 2696 356 163 191 232 94 482 362 222 306 147 180 157

56 19 95 14 188 70 18 15 16 13 222 85 20 66 22 25 32

232

Th/ U

238

0.06 0.13 0.31 0.08 0.07 0.20 0.11 0.08 0.07 0.14 0.48 0.24 0.09 0.22 0.15 0.14 0.21

% Pbc

206

0.27 2.42 0.84 6.16 0.32 1.10 3.87 6.12 2.42 7.62 1.56 2.01 1.39 1.72 2.39 1.72 2.81

206

Pb* (ppm)

95 12 29 13 210 37 12 16 18 8 38 28 17 23 12 13 12

Total U/ 206 Pb

+%

9.390 10.860 9.270 11.630 11.050 8.320 11.650 10.160 11.110 10.660 10.800 10.940 11.070 11.600 10.700 11.630 11.150

1.50 2.10 1.70 2.10 1.40 1.70 2.20 2.10 2.00 3.00 1.70 1.90 2.30 2.00 2.70 2.40 2.80

238

Total Pb/ 206 Pb

+%

0.06640 0.07550 0.06660 0.07490 0.06119 0.07220 0.08110 0.08540 0.07750 0.07860 0.06880 0.06670 0.07510 0.07090 0.08390 0.07640 0.08390

2.00 2.70 1.90 2.80 0.79 1.80 5.00 2.70 4.60 4.90 1.90 2.30 5.70 2.90 3.90 3.70 3.80

207

238

U/ Pb*

+%

9.410 11.130 9.350 12.390 11.090 8.410 12.110 10.830 11.390 11.540 10.970 11.160 11.230 11.800 10.960 11.840 11.470

1.50 2.20 1.80 2.50 1.40 1.80 2.50 2.50 2.20 4.40 1.80 2.00 2.60 2.10 3.10 2.90 3.60

206

207

Pb*/ Pb*

+%

0.06420 0.05570 0.05970 0.02300 0.05855 0.06320 0.04900 0.03400 0.05780

2.50 14.00 6.00 51.00 1.50 5.90 20.00 33.00 12.00

0.05600 0.05020 0.06400 0.05690 0.06500 0.06200 0.06100

7.30 9.20 17.00 9.50 21.00 21.00 30.00

206

206

Pb/ U age

1s error

650.9 554.5 654.8 500.4 556.7 724.2 511.3 569.5 542.6 535.7 562.2 553.0 550.0 524.2 562.9 522.8 538.9

9.1 11.9 11.0 12.1 7.6 12.2 12.3 13.7 11.3 22.7 9.6 10.7 13.6 10.8 16.7 14.5 18.7

238

207

Pb/ Pb age

1s error

% Discordant

Notes

206

748.4 441.9 594.0

53.3 305.6 129.2

15 220 29

550.5 713.5 165.5

33.3 125.4 474.6

21 21 268

520.9

269.8

24

453.4 203.6 738.6 489.2 759.9 690.2 643.6

161.0 212.8 350.3 210.2 438.2 450.4 649.3

219 263 34 27 35 32 19

Inh core Igneous Inh core Igneous Rim Inh core Igneous Igneous Igneous Igneous Inh core Inh core Igneous Igneous Igneous Igneous Igneous


K. SHIRAISHI ET AL.

Session A5 1.1 1.2 2.1 2.2 2.3 3.1 3.2 4.1 4.2 5.1 5.2 6.1 6.2 7.1 7.2 8.1 8.2

U (ppm)

Table 9. SHRIMP U–Pb data for zircon from sample 90112302B Spot

131 619 109 175 630 129 85 436 335 658 590 162 775 1238 476 611 507 116 394 882 149 443 291 207 158 358 323 96 508 715 126 67 4021

56 68 25 77 32 42 25 132 146 264 92 52 276 107 242 311 216 39 102 103 50 189 144 71 60 117 135 25 201 313 52 20 101

232

Th/ U

238

0.44 0.11 0.24 0.46 0.05 0.34 0.31 0.31 0.45 0.41 0.16 0.33 0.37 0.09 0.53 0.52 0.44 0.35 0.27 0.12 0.35 0.44 0.51 0.35 0.40 0.34 0.43 0.27 0.41 0.45 0.43 0.30 0.03

% Pbc

206

0.55 0.20 1.43 0.30 1.36 0.36 1.16 0.11 0.09 0.18 0.09 0.83 0.03 0.07 0.15 0.05 0.16 0.63 0.25 0.13 0.42 0.02 0.36 0.29 0.23 0.51 0.41 0.24 0.18 0.33 2.26 0.04

206

Pb* (ppm)

19 65 7 15 64 17 6 65 45 70 52 23 67 112 65 80 43 10 32 96 16 38 38 17 23 30 25 7 43 60 14 8 450

Total U/ 206 Pb

+%

6.050 8.216 13.100 9.951 8.452 6.406 12.000 5.803 6.347 8.120 9.841 6.089 9.982 9.520 6.322 6.563 10.173 9.710 10.683 7.937 8.070 10.146 6.558 10.620 5.960 10.260 11.280 11.170 10.205 10.286 7.550 7.580 7.679

2.00 0.75 1.40 0.93 0.74 0.93 1.10 0.76 1.10 0.73 0.75 0.88 0.73 0.78 1.60 0.75 0.76 1.80 0.78 0.82 7.30 0.88 0.83 2.00 0.91 0.83 1.10 1.10 0.76 0.76 1.30 1.60 0.75

238

Total Pb/ 206 Pb

+%

0.07519 0.06647 0.07120 0.06112 0.07543 0.07442 0.06820 0.07344 0.07222 0.06837 0.06169 0.07778 0.06148 0.06225 0.07374 0.07088 0.06177 0.06961 0.06085 0.06520 0.06905 0.06127 0.07030 0.06233 0.07390 0.06251 0.06396 0.06450 0.06124 0.06048 0.06941 0.08620 0.06427

1.00 0.53 1.50 1.50 1.20 1.10 1.70 0.52 0.63 0.54 0.67 0.85 0.56 0.73 0.54 0.97 0.71 1.40 0.82 0.49 1.10 0.76 0.73 1.20 0.93 0.87 1.00 2.20 0.81 0.90 1.40 2.60 0.35

207

238

U/ Pb*

+%

6.090 8.232 13.290 9.980 8.569 6.428 12.140 5.809 6.353 8.135 9.850 6.141 9.985 9.527 6.331 6.566 10.190 9.770 10.709 7.948 8.100 10.148 6.554 10.660 5.977 10.284 11.340 11.220 10.229 10.305 7.570 7.760 7.683

2.00 0.75 1.50 0.94 0.76 0.95 1.10 0.76 1.10 0.74 0.75 0.90 0.73 0.78 1.60 0.76 0.76 1.90 0.79 0.82 7.30 0.88 0.83 2.00 0.92 0.84 1.10 1.20 0.77 0.76 1.30 1.70 0.75

206

207

Pb*/ Pb*

+%

0.07060 0.06485 0.05970 0.05870 0.06420 0.07150 0.05880 0.07251 0.07147 0.06686 0.06092 0.07080 0.06123 0.06167 0.07246 0.07045 0.06045 0.06450 0.05884 0.06415 0.06560 0.06112 0.07073 0.05940 0.07150 0.06061 0.05980 0.06120 0.05931 0.05898 0.06670 0.06760 0.06392

2.40 0.79 7.00 2.90 2.50 2.30 5.70 0.75 0.86 0.84 0.92 2.20 0.67 0.80 0.76 1.00 1.00 3.40 1.40 0.71 2.20 0.83 0.93 2.90 1.60 1.60 2.10 5.60 1.30 1.20 2.40 6.50 0.36

206

206

Pb/ 1s U error age

238

980.7 739.0 467.6 615.6 711.5 932.0 510.1 1023.9 942.3 747.3 623.3 972.6 615.3 643.4 945.3 913.8 603.5 628.1 575.5 764.0 750.4 605.8 915.4 578.0 997.2 598.2 544.9 550.4 601.3 597.0 799.7 781.6 788.8

17.8 5.3 6.8 5.5 5.1 8.2 5.6 7.2 10.1 5.2 4.5 8.1 4.3 4.8 13.7 6.4 4.4 11.1 4.3 5.9 51.7 5.1 7.1 11.1 8.5 4.8 5.8 6.1 4.4 4.3 10.1 12.1 5.5

207

Pb/ Pb age

1s error

% Discordant

Notes

944.8 769.4 592.9 555.3 749.0 970.7 558.7 1000.4 971.0 833.5 636.4 952.2 647.4 662.6 998.9 941.5 619.6 756.5 561.3 746.5 793.9 643.3 949.6 580.6 971.2 625.3 595.8 644.5 578.7 566.4 828.8 854.9 738.9

49.1 16.6 150.9 62.9 52.1 47.1 125.0 15.3 17.5 17.4 19.7 44.3 14.4 17.2 15.5 20.8 22.4 70.7 31.2 14.9 45.8 17.8 19.0 63.4 33.0 34.0 46.1 120.7 28.5 25.3 51.1 135.1 7.7

24 4 27 210 5 4 10 22 3 12 2 22 5 3 6 3 3 20 22 22 6 6 4 0 23 5 9 17 24 25 4 9 26

Inh core ig Inner rm Leach rim Sector Inner rim Inh core ig Leach rim Inh core ig Inh core ig Mixed Inh core ig crack Out rim / mix? Outer rim Outer rim Inh core ig Inh core ig mix? Outer rim Inh core ig Outer rim Inner rim Inh core frag Outer rim Inh core sector Sector Inh core ig Outer rim Mixed Leach rim Out? unzoned Outer rim INH core ig crack Inh core ig crack Inner rim

206

37



Session A6 1.1 1.2 2.1 2.2 3.1 3.3 4.1 4.2 4.3 4.4 4.5 5.1 5.2 6.2 6.3 7.1 7.2 8.1 8.2 9.1 9.2 10.1 10.2 11.1 12.1 12.2 12.3 13.1 13.2 14.1 14.2 15.1 15.2

U Th (ppm) (ppm)

38


232

Th/ U

% Pbc

Th (ppm)

Session A4 2.a 2.1 2.2 3.1 4.1 4.2 5.1 5.2 6.1 8.2 9.1 10.1 10.2 11.1 11.2

302 162 471 412 441 1239 278 3485 271 1550 278 1512 427 10361 3880

202 48 229 172 229 78 107 207 72 90 92 96 142 1005 1232

0.69 0.30 0.50 0.43 0.54 0.06 0.40 0.06 0.28 0.06 0.34 0.07 0.34 0.10 0.33

0.17 1.42 0.72 1.26 0.46 1.86 0.79 0.46 1.16 3.59 1.24 0.45 0.77 0.02 0.04

Session N1 8.3 13.1 14.1 15.1 16.1 17.1

394 589 806 319 787 1029

131 184 68 95 208 39

0.34 0.32 0.09 0.31 0.27 0.04

0.55 0.29 0.24 1.03 3.69 0.25

238

206

206

Total U/ 206 Pb

+%

50 13 38 33 35 100 40 298 20 124 21 119 33 955 318

5.194 10.700 10.710 10.750 10.910 10.611 6.050 10.059 11.470 10.710 11.310 10.920 10.990 9.325 10.496

31 46 63 25 63 81

10.940 10.930 11.080 10.970 10.810 10.910

Pb* (ppm)

Total Pb/ 206 Pb

+%

10.00 1.20 1.10 1.00 1.00 0.92 1.10 0.88 1.20 0.94 1.30 0.93 1.10 0.86 0.88

0.07787 0.07340 0.06580 0.06673 0.06264 0.07403 0.07758 0.06282 0.06450 0.08880 0.06580 0.06240 0.06570 0.05891 0.05971

1.20 1.10 1.10 1.50 1.10 1.10

0.06420 0.06157 0.06056 0.06172 0.09590 0.06110

238

238

U/ Pb*

+%

0.71 1.50 0.93 1.50 1.10 0.65 1.00 0.77 1.70 1.50 1.70 0.72 2.10 0.26 0.44

5.203 10.850 10.790 10.880 10.960 10.810 6.098 10.105 11.610 11.110 11.450 10.970 11.070 9.326 10.499

1.10 0.87 0.77 1.30 6.20 0.83

11.000 10.960 11.110 11.090 11.230 10.940

207

206

207

Pb/ U age

1s error

207

Pb/ Pb age

1s error

% Discordant

Notes

0.94 5.80 2.50 3.70 3.00 2.50 2.20 1.10 5.90 5.20 6.20 1.60 4.00 0.28 0.47

1133.3 568.3 571.3 566.7 562.9 570.2 978.9 608.3 532.8 555.5 539.6 562.4 557.3 656.6 586.5

10.4 7.0 6.3 5.7 5.6 5.2 10.3 5.1 6.6 5.4 6.9 5.1 6.1 5.4 4.9

1107.6 668.1 601.6 470.1 564.4 560.4 956.6 569.1 412.1 587.7 439.9 558.0 582.0 559.0 582.3

18.7 123.7 54.8 81.4 66.1 54.3 45.9 24.1 132.8 113.0 137.0 33.9 87.8 6.1 10.2

22 18 5 217 0 22 22 26 223 6 218 21 4 215 21

Inherited Igneous? Ig. core Ig. core Ig. core Rim Inherited Mixed Ig. core Rim Ig. core Rim Ig. core Inherited Igneous?

2.40 2.20 1.40 4.10 9.70 1.10

560.8 562.7 555.7 556.7 549.9 563.8

6.5 6.2 6.0 8.3 6.1 6.0

594.7 575.1 552.5 341.2 808.5 569.7

51.8 48.6 30.8 93.2 203.8 24.2

6 2 21 239 47 1

Ig. core Ig. core Rim Ig. core Ig. core Rim

Pb*/ Pb*

+%

10.00 1.30 1.20 1.10 1.00 0.95 1.10 0.88 1.30 1.00 1.30 0.94 1.10 0.86 0.88

0.07648 0.06180 0.05990 0.05640 0.05890 0.05880 0.07100 0.05906 0.05500 0.05960 0.05570 0.05876 0.05940 0.05878 0.05942

1.20 1.10 1.10 1.60 1.20 1.10

0.05980 0.05920 0.05861 0.05330 0.06610 0.05907

206

206

238

206


K. SHIRAISHI ET AL.

U (ppm)

99 182 216 246 172 101 147 41 181 315 280 315 304 244 256 161 Session N6 1.1 1.2 2.1 2.2 3.1 3.2 5.1 6.1

All errors are quoted at 1s; Pbc and Pb* indicate the common and radiogenic portions, respectively. Common Pb corrected assuming 206Pb/238U – 207Pb/235U concordance. External spot-to-spot error from scatter in standard is 0.85% (included in the calculation of sample-spot errors). Error in standard calibration is 0.41% (not included in errors but required when comparing data from different sessions).

22.3 9.0 5.9 6.0 12.1 8.9 7.6 7.7 561.1 594.0 571.1 570.4 552.1 585.3 560.6 559.5 9.74 6.38 0.51 0.50 7.87 2.58 4.03 0.99 0.2492 0.1286 0.0788 0.0955 0.1653 0.2089 0.1499 0.2860 1.13 1.10 1.06 1.07 1.27 1.28 1.10 1.13 8.430 9.488 10.534 10.327 9.719 8.598 9.778 7.957 18.4 28.5 22.8 26.2 26.9 24.4 22.5 17.4 0.57 0.60 0.80 0.81 0.58 0.43 0.60 0.27

23.34 8.43 2.42 4.47 13.09 18.29 11.16 27.86

206

Pbc 206

% Th/238U 232

Th (ppm) U (ppm) Spot

Table 11. SHRIMP U– Pb data for titanite from sample 9091405A

Pb* (ppm)

Total 238U/206Pb

+%

Total

207

Pb/206Pb

+%

206

Pb*/238U age

1s error


39

Other minerals are abundant plagioclase, quartz and minor K-feldspar, occasionally with myrmekites on grain boundaries. Monazite, apatite and opaque minerals are also found. Monazites from this sample were dated by CHIME and yield 542 +12 Ma for rims of zoned grains and 517 +14 Ma for chronologically homogeneous grains (Asami et al. 2005). Zircon grains have a variety of rounded and anhedral morphologies. Both elongate and squat grain types are mostly oscillatory zoned, with elongate grains having lower CL. Cores with oscillatory zoning are truncated at rounded boundaries by oscillatory-zoned rims. A few grains have minor overgrowths of unzoned zircon, on the ends of prismatic cores and rims. U contents of oscillatory-zoned zircon vary between 50 and 530 ppm, with Th/U ratios between 0.2 and 0.9. Analysis 2.1, which has a U content of 508 ppm and a Th/U ratio of 0.05, was performed on a low-CL band between prismatic zircon and an overgrowth. Rounded cores have consistently higher Pb –Pb spot ages than rims and core-free grains (Fig. 4b, Table 2). Most of the core age data are discordant, but four concordant analyses define a concordia age of 1014 + 15 Ma (MSWD ¼ 1.4, probability of concordance ¼ 0.4, Fig. 5b). From oscillatory-zoned grains and rims, five concordant analyses define a concordia age of 790.8 + 9.8 Ma (MSWD ¼ 1.2, probability of concordance ¼ 0.09). The concordia ages are interpreted as timing the growth of igneous zircon at c. 791 Ma, from a magma that inherited zircon from source or country rocks, which grew magmatically at c. 1014 Ma. As the host rock represents metamorphosed sediment, the igneous zircon is detrital, and it is worth noting that all c. 1014 Ma zircon is inherited in c. 791 Ma zircon, so that analysed detrital grains may have derived from the weathering of a single igneous source. Overgrowths of unzoned zircon were not dated, but may be metamorphic. Sample 85011503D. This is an enderbitic orthogneiss from the northwestern part of Brattnepene and consists of plagioclase, quartz, orthopyroxene, brown hornblende, biotite, clinopyroxene and a trace amount of garnet. On the basis of various geothermobarometries, the peak metamorphic conditions were estimated to be around 800 8C and 7– 8.5 kbar (Shiraishi & Kojima 1987). The sample was used for Sm–Nd and Rb–Sr isotope analyses as one of four whole-rock samples (Shiraishi & Kagami 1992). The isochrons yielded 978 + 52 Ma (initial ratio 0.70426) and 961 +101 Ma (initial ratio 0.51163) for the Rb– Sr and Sm –Nd systems, respectively. A Sm –Nd internal mineral isochron yields an age of 624 +18 Ma, with an initial ratio of 0.51193.

Spot

232

Th/ U

% Pbc

Th (ppm)

Session A3 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 16.2 17.1

1398 1744 1550 2269 1917 1099 1774 2265 1054 1988 2796 407 1157 2722 1584 2336

40 177 272 29 108 111 99 94 232 129 156 35 158 74 71 189

0.03 0.10 0.18 0.01 0.06 0.10 0.06 0.04 0.22 0.07 0.06 0.09 0.14 0.03 0.05 0.08

0.24 0.05 0.18 0.03 0.04 0.07 0.16 0.02 0.12 0.02 0.06 0.88 0.85 0.03 0.33 0.24

Session N5 1.3 2.3 5.3 6.3 7.2 8.3 10.3 14.4 14.5 16.3 18.1 18.2 19.1 19.2 20.1 21.1 22.1 23.1 23.2 24.1

102 201 107 720 1907 88 2479 180 2744 89 662 164 2538 52 5419 44 78 99 886 75

47 191 72 88 93 45 180 101 39 65 103 79 33 30 77 23 43 59 47 35

0.48 0.98 0.69 0.13 0.05 0.53 0.08 0.58 0.01 0.75 0.16 0.50 0.01 0.60 0.01 0.54 0.56 0.62 0.06 0.48

0.43 0.29 0.58 0.46 0.06 0.08 0.01 0.03 0.00 0.07 20.01 0.33 0.03 20.24 0.00 20.07 20.15 0.07 0.00 0.04

238

206

206

Total U/ 206 Pb

+%

153 200 176 264 216 125 195 258 121 224 316 44 128 308 107 161

10.944 10.682 11.038 10.286 10.725 10.846 11.027 10.576 11.037 10.777 10.705 11.228 11.174 10.605 11.102 10.946

9 18 9 53 154 8 207 16 225 8 52 14 213 5 508 4 6 9 74 6

9.800 9.390 10.450 11.710 10.620 9.670 10.300 9.940 10.470 9.240 10.840 9.780 10.240 9.290 9.160 10.690 10.470 9.170 10.340 10.610

Pb* (ppm)

Total Pb/ 206 Pb

+%

1.31 1.05 1.06 1.07 1.05 1.15 1.43 1.03 1.13 1.22 1.02 1.23 1.33 1.03 1.17 1.09

0.06069 0.05910 0.06030 0.05934 0.05893 0.05898 0.06005 0.05872 0.05926 0.05975 0.05952 0.06505 0.06505 0.05898 0.06358 0.06164

2.90 3.00 1.50 4.50 2.80 1.80 1.20 5.60 4.60 2.70 1.30 1.30 2.40 1.60 4.50 6.00 2.80 1.70 5.20 2.60

0.06370 0.06400 0.06120 0.06512 0.06247 0.06080 0.05963 0.06383 0.05935 0.06340 0.05858 0.06079 0.06032 0.05980 0.06109 0.06030 0.06020 0.06087 0.05916 0.06060

238

238

U/ Pb*

+%

0.63 0.54 0.33 0.39 0.37 0.70 0.37 0.37 0.40 0.44 0.25 0.72 0.48 0.32 0.66 0.67

10.970 10.688 11.058 10.290 10.730 10.853 11.045 10.578 11.051 10.779 10.711 11.328 11.270 10.608 11.139 10.972

1.90 1.90 1.90 1.10 1.30 1.60 0.30 1.30 0.75 2.20 1.40 1.10 0.68 2.20 1.40 2.90 2.20 1.40 1.30 2.40

9.840 9.420 10.520 11.770 10.630 9.680 10.300 9.940 10.470 9.250 10.840 9.810 10.240 9.270 9.160 10.680 10.460 9.180 10.340 10.610

207

207

Pb*/ Pb*

+%

1.31 1.05 1.06 1.07 1.05 1.15 1.43 1.03 1.13 1.22 1.02 1.24 1.33 1.03 1.17 1.09

0.05871 0.05867 0.05882 0.05907 0.05857 0.05842 0.05874 0.05854 0.05828 0.05957 0.05906 0.05785 0.05810 0.05873 0.06090 0.05968

2.90 3.00 1.50 4.50 2.80 1.80 1.20 5.60 4.60 2.70 1.30 1.40 2.40 1.60 4.50 6.00 2.80 1.70 5.20 2.60

0.06020 0.06160 0.05650 0.06141 0.06198 0.06020 0.05956 0.06360 0.05938 0.06280 0.05865 0.05800 0.06009 0.06180 0.06108 0.06090 0.06140 0.06030 0.05920 0.06020

206

206

Pb/ U age

1s error

0.75 0.58 0.63 0.42 0.41 0.74 0.44 0.38 0.48 0.44 0.27 1.97 1.02 0.32 0.99 0.85

562.4 576.6 558.1 597.9 574.4 568.2 558.7 582.3 558.5 571.9 575.4 545.4 548.0 580.7 554.2 562.3

3.00 2.50 4.30 1.40 1.30 1.90 0.31 1.70 0.75 3.20 1.40 2.40 0.69 2.80 1.40 3.10 3.60 1.70 1.30 2.70

624.0 651.0 585.6 526.0 580.0 634.0 597.3 618.0 588.0 662.0 568.7 625.7 601.0 660.3 668.0 577.0 589.0 667.0 595.0 580.0

206

238

207

Pb/ Pb age

1s error

% Discordant

Notes

206

7.0 5.8 5.7 6.1 5.8 6.3 7.7 5.7 6.0 6.7 5.6 6.5 7.0 5.7 6.2 5.9

556.3 554.7 560.5 569.8 551.1 545.7 557.5 550.1 540.4 588.0 569.5 524.1 533.6 556.9 635.6 592.1

16.3 12.5 13.7 9.1 9.0 16.0 9.6 8.3 10.7 9.6 5.8 43.7 22.3 7.1 21.5 18.5

21 24 0 25 24 24 0 26 23 3 21 24 23 24 13 5

Metm rim Metm rim Rim, crack Rim, crack Metm rim Metm rim Rim, crack Metm rim Mixed, crack Metm rim Metm rim Rim, crack Mixed, crack Metm rim Rim, crack Metm rim

17.0 19.0 8.5 23.0 15.0 11.0 7.1 33.0 26.0 17.0 6.9 8.1 14.0 9.9 28.0 33.0 16.0 11.0 29.0 14.0

610.0 661.0 471.0 654.0 673.0 610.0 587.7 729.0 581.0 702.0 554.0 532.0 607.0 666.0 642.0 635.0 652.0 614.0 574.0 611.0

65.0 54.0 95.0 31.0 28.0 41.0 6.7 36.0 16.0 67.0 31.0 52.0 15.0 59.0 31.0 66.0 78.0 38.0 28.0 59.0

22 2 220 24 16 24 22 18 21 6 23 215 1 1 24 10 11 28 23 5

Ig. core Ig. core Ig. core, crack Rim, crack Rim, crack Ig. core Metm. rim Ig. core Metm. rim Ig. core Metm. rim Ig. core, crack Metm. rim Ig. core Metm. rim Ig. core Ig. core Ig. core Metm. rim Metm. rim


K. SHIRAISHI ET AL.

U (ppm)

40

Table 12. SHRIMP U–Pb data for zircon from sample 85012817

Table 13. SHRIMP U–Pb data for zircon from sample 9031507 Spot

232

Th (ppm)

Session A6 1.1 1.2 2.1 2.2 3.1 4.1 5.1 5.2 6.1 7.1 7.2 8.1

70 61 68 271 72 54 97 107 10 257 68 9

32 19 28 95 24 20 55 70 3 55 26 2

0.47 0.33 0.42 0.36 0.34 0.37 0.58 0.68 0.29 0.22 0.40 0.24

0.27 0.91 0.56 0.22 0.20 0.16 0.76 0.91 8.53 0.35 0.84 1.05

9.1

315

101

0.33

0.09

238

206

206

Total U/ 206 Pb

+%

9 8 9 36 10 7 13 14 1 34 9 1

6.443 6.623 6.539 6.521 6.346 6.626 6.459 6.373 7.030 6.586 6.392 6.800

40

6.765

Pb* (ppm)

Total Pb/ 206 Pb

+%

1.20 1.30 1.20 0.96 1.20 1.30 1.40 1.20 2.50 0.97 1.20 2.50

0.07290 0.07620 0.07460 0.06996 0.07270 0.07410 0.07460 0.07381 0.09600 0.07013 0.07290 0.09800

0.95

0.07004

238

238

U/ Pb*

+%

1.90 1.50 1.40 0.73 1.40 1.60 1.90 1.20 4.70 0.76 1.40 3.40

6.461 6.684 6.576 6.536 6.359 6.637 6.508 6.431 7.690 6.609 6.446 6.870

1.30 1.30 1.20 0.96 1.20 1.30 1.50 1.20 3.60 0.98 1.30 2.60

0.69

6.770

0.95

207

206

207

Pb/ U age

1s error

207

Pb/ Pb age

1s error

% Discordant

10.8 11.0 10.5 8.2 10.7 11.3 12.9 10.4 26.7 8.3 11.1 21.7

946.0 887.4 924.4 873.4 958.4 1006.1 875.0 813.4

85.8 86.9 63.1 24.9 42.0 61.0 106.0 74.6

2 21 1 25 2 11 25 213

1.40 4.30 9.10

927.7 898.8 912.6 917.8 941.5 904.7 921.4 931.7 788.2 908.2 929.7 875.8

845.0 803.4 1415.7

30.0 90.5 173.1

27 214 62

0.86

888.1

7.9

908.5

17.8

2

Pb*/ Pb*

+%

0.07060 0.06860 0.06990 0.06816 0.07100 0.07270 0.06820 0.06620

4.20 4.20 3.10 1.20 2.10 3.00 5.10 3.60

0.06723 0.06590 0.08950 0.06933

206

206

238

206

Notes

Igneous Igneous Igneous Igneous Igneous Igneous Igneous Igneous Leached rim Rim Igneous Leached patch Rim


Th/ U

% Pbc

U (ppm)


41

42

Table 14. SHRIMP U–Pb data for titanite from sample 9031507 Spot

Th (ppm)

96 74 208 47 113 113 93 87 46 68 100 131 77 45

40 35 62 22 29 41 40 46 24 41 49 29 46 22

232

Th/238U

0.43 0.49 0.31 0.48 0.26 0.37 0.44 0.54 0.54 0.63 0.51 0.23 0.62 0.49

%

206

Pbc

1.78 1.64 1.49 1.78 1.64 1.70 1.62 1.76 2.02 1.78 1.41 1.44 1.85 2.17

206

Pb* (ppm)

7.1 5.1 14.0 3.5 8.2 8.1 6.9 6.3 3.3 5.0 7.2 9.7 5.4 3.3

238

Total U/206Pb

+%

11.63 12.47 12.75 11.71 11.88 12.00 11.62 11.78 11.85 11.68 11.87 11.65 12.25 11.82

1.23 1.13 0.99 1.14 1.02 1.02 1.04 1.04 1.13 1.07 1.03 1.01 1.05 1.44

207

Total Pb/206Pb

0.07230 0.07024 0.06875 0.07221 0.07091 0.07121 0.07105 0.07197 0.07396 0.07226 0.06908 0.06957 0.07218 0.07520

+%

1.02 1.22 0.73 1.37 0.89 0.87 0.96 0.99 1.33 1.10 0.93 0.98 1.04 1.34

206

Pb*/238U age

522.8 489.4 479.8 519.4 512.9 507.7 524.0 516.3 512.2 520.4 514.2 523.4 496.8 512.5

1s error

6.3 5.4 4.7 5.8 5.1 5.1 5.3 5.3 5.7 5.5 5.2 5.2 5.1 7.3

Notes

Young Inclusions

Inclusions

All errors are quoted at 1s; Pbc and Pb* indicate the common and radiogenic portions, respectively. Common Pb corrected assuming 206Pb/238U-207Pb/235U concordance. External spot-to-spot error from scatter in standard is 0.93% (included in the calculation of sample-spot errors). Error in standard calibration is 0.38% (not included in errors but required when comparing data from different sessions).

K. SHIRAISHI ET AL.

Session A7 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

U (ppm)


43

Fig. 4. Representative images of zircon grains from samples of the NE Terrane. All images are of cathodoluminescence unless otherwise marked (BSE ¼ backscattered electron). Each SHRIMP analysis spot is labelled with grain analysis number, U content (ppm)/Th/U ratio and 204Pb-corrected 206Pb/238U age with 1s error. Ages that are discordant with 207Pb/235U ages outside 68.3% confidence limits are labelled (d).

Zircon grains are squat to elongate, 50–300 mm long, with rounded, anhedral surfaces. Most grains have oscillatory or graduated prismatic growth zones that decrease in CL brightness from core to rim. All prismatically zoned zircon grains have truncated, rounded edges, which are commonly surrounded by thin rims of high-CL, unzoned zircon. These overgrowths are thickest on the terminations of elongate grains, and are similar to discrete, squat grains of zircon that are unzoned or have gradational or sector zoning (Fig. 4c). Prismatically zoned cores and rims contain 90– 830 ppm U and have Th/U ratios between 0.14 and 0.40 (Table 3). High-CL, simple zoned rims and grains have uniform U contents (90 –120 ppm) and Th/U ratios (0.19–0.23). Six concordant data from prismatic cores define a concordia age of 951 + 17 Ma (MSWD ¼ 1.6, probability of concordance ¼ 0.6, Fig. 5c). In contrast, all eight data from high-CL rims and grains define a concordia age of 602 + 15 Ma (MSWD ¼ 1.3, probability of concordance ¼ 0.01). On the basis of zircon texture and composition, the c. 951 Ma age is interpreted as the time of magmatic zircon growth in the igneous protolith, and the c. 602 Ma age as the time of granulite-grade metamorphism.

Sample 9032401A. This is a strongly sheared sillimanite–garnet–biotite paragneiss from the northern part of Austkampane. Other constituent minerals are plagioclase, K-feldspar, quartz, zircon, apatite and opaque minerals. Secondary muscovite is common. Zircon grains are a mixture of squat to equant, clear grains with some pyramidal faceting preserved, and anhedral grains that contain irregular cores with truncated simple and oscillatory zoning, and mid- to high-CL rims of graduated and sector-zoned zircon (Fig. 4d). The latter rims are texturally identical to the clear, equant zircon grains. Irregular cores have U contents between 80 and 720 ppm U and Th/U values between 0.3 and 0.8 (except analysis 8.1, Table 4). Analysis 8.1 is of an irregular core with no CL, and has an unusually high U content (1113 ppm) and low Th/U ratio (0.01). Clear rims and grains have U contents between 230 and 810 ppm and Th/U ratios with both moderate (0.3–0.5) and low (less than 0.1) values. Ages from seven concordant analyses are grouped at c. 3260 Ma (one grain), c. 1200 Ma (two grains) and c. 1100 Ma (four grains), with the latter four analyses defining a 206Pb/238U mean age of 1125 + 18 Ma

44

K. SHIRAISHI ET AL.

Fig. 5. Tera –Wasserburg concordia diagrams of U– Pb SHRIMP data. (a) Zircon from 85020401C; (b) zircon from 84022004; (c) zircon from 85011503D; (d) zircon from 9032401A.

(MSWD ¼ 0.7) (Fig. 5d). All ages are interpreted as those of detrital zircon with an igneous origin, with the exception of analysis 8.1, which may be metamorphic. With the exception of analysis 12.2, which has a concordant age of c. 517 Ma, data from clear rims and grains define a concordia age of 637 + 6 Ma (MSWD ¼ 1.2, probability of concordance ¼ 0.5). The concordia age is interpreted as timing the growth of zircon during granulite-grade metamorphism. Sample 90102801A. This is a strongly deformed and layered hornblende orthogneiss from Isachsenfjella. Deformed syntectonic granite intrudes the gneiss (Fig. 6b). It consists of plagioclase, quartz, hornblende, biotite, titanite and zircon. Zircon grains are squat to elongate, 50–200 mm long, and mostly anhedral with a few faceted surfaces preserved. Internally, all grains have prismatic and oscillatory zoning subparallel to grain edges. In several grains, minor patches and discontinuous overgrowths of mid- to low-CL zircon are present, although in most cases there are no clear boundaries between cores and rims (Fig. 4e). Prismatic, oscillatory-

zoned zircon has U contents between 140 and 530 ppm U and Th/U ratios between 0.2 and 0.8 (Table 5). Overgrowths have similar U contents (150 –570 ppm U) but lower Th/U ratios (0.02– 0.22), except for analysis 15.1 of a low-CL overgrowth with 2436 ppm U and a Th/U ratio of 0.07. Age data are scattered, with no clear distinction between cores and rims (Fig. 7a). Together, five concordant data from cores and rims define a concordia age of 1133 +12 Ma (MSWD ¼ 1.0, probability of concordance ¼ 0.2). This is interpreted as the time of zircon growth in the magmatic protolith. The significance of rim ages is ambiguous. Textures and U/Th contents could be either late magmatic or metamorphic, with the exception of discordant analysis 15.1, which is more typical of metamorphic zircon growth. There is no evidence of zircon growth after c. 1050 Ma. Titanite is present in the orthogneiss as metamorphic grains intergrown with microcline and hornblende, and as thin rims and overgrowths on ilmenite. Grains analysed (Table 6) show little compositional zoning, and contain up to 200 ppm U and 60 ppm Th. Excluding analyses 3.1 and 6.1


45

Fig. 6. Photographs of the selected outcrops. (a) Layered hornblende– biotite gneiss (90112302B) intruded by biotite granite dyke (90112302A), Balchenfjella (eastern part of NE Terrane). (b) Foliated hornblende–biotite gneiss (90102801A), Isachsenfjella (eastern part of NE Terrane). (c) Mylonitic granite (9091405A) at Main Shear Zone, Wiederoefjellet (SW Terrane). (d) Massive tonalitic hornblende gneiss (9031507), Mefjell (SW Terrane).

on cracked titanite, data uncorrected for common Pb define a Model 1 linear array data with a lower intercept age of 516 + 9 Ma (MSWD ¼ 1.4) (Fig. 7b). The same eight analyses define an identical 207Pb-corrected 206Pb/238U mean age of 517 + 8 Ma (MSWD ¼ 1.2). Because of the potential for Pb loss from titanite by diffusion at temperatures over 660 8C (Cherniak 1993), the age may have been reset after growth, and therefore represents only a minimum estimate for the timing of the formation of the metamorphic assemblage. Sample 90112102A. This is a 2 m wide layer of garnet –biotite paragneiss in hornblende gneiss from southern Balchenfjella. It consists of garnet, biotite, plagioclase, K-feldspar, antiperthite, zircon and opaque minerals. Myrmekite replacing K-feldspar and secondary muscovite are present. Zircon grains are rounded and mostly equant or squat, with fewer elongate grains up to 150 mm long. The latter mostly contain irregular or rounded cores with oscillatory zoning, surrounded by rims of mid- to high-CL zircon with simple or

sector zoning. The latter rims are similar to the main population of equant grains (Fig. 8a). To compare analytical results between the ANU and NIPR SHRIMP facilities, data were processed separately from two sessions (Table 7, Fig. 7c and d). Only three analyses (4.2, 18.1 and 30.1) are available from irregular cores, which have U contents of about 200 ppm and Th/U ratios above 0.3. Spot 206Pb/238U ages are scattered between c. 1150 and 760 Ma. The remaining analyses have U contents between 130 and 3400 ppm and Th/U ratios below 0.1, and ages that cluster around 600 Ma. Concordant analyses from session A2 (ANU) define a concordia age of 601 + 6 Ma (MSWD ¼ 1.1, probability of concordance ¼ 0.1, Fig. 7c). Analyses from session N5 (NIPR) are more scattered; however, all younger analyses are from spots on cracked zircon, and when excluded the remaining analyses define a concordia age of 593 + 8 Ma (MSWD ¼ 1.3, probability of concordance ¼ 0.5, Fig. 7d). The concordia ages from the two sessions are identical within error, and are interpreted as timing zircon growth during granulite-grade metamorphism. Ages from irregular

46

K. SHIRAISHI ET AL.

Fig. 7. Tera– Wasserburg concordia diagrams of U– Pb SHRIMP data. (a) Zircon from 90102801A; (b) titanite from 90102801A; (c) zircon from 90112102A (session A2); (d) zircon from 90112102A (session N5).

cores are older but scattered, and these cores are interpreted as detrital zircon. Sample 90112302A. This is a biotite granite dyke intruded across intercalated layers of mafic gneiss and migmatitic biotite felsic gneiss (sample 90112302B), in north Balchenfjella (Fig. 6a). Boundaries with the host gneiss are soft and nebulitic, with entrainment of deformed fragments of host gneiss and no evidence of chill margins. These textures suggest emplacement under high-temperature, ductile conditions. The granitoid is composed of biotite, hornblende, plagioclase, quartz, K-feldspar and titanite, with secondary muscovite and carbonates. Zircon grains are squat and euhedral, with prismatic forms only slightly dominant over pyramidal forms and length to width ratios of 2:1 or less. All grains are oscillatory zoned, with several grains showing a decrease in CL brightness from core to rim (Fig. 8b). Many grains have rounded cores with irregular and oscillatory growth zones (Fig. 8c).

Oscillatory-zoned zircon has U contents mostly between 90 and 310 ppm and Th/U ratios between 0.07 and 0.22, excepting analysis 2.3 of a low-CL rim with 2696 ppm U (Table 8, Fig. 8b). Compositions from rounded cores are variable, with U contents between 310 and 1040 ppm and Th/U ratios between 0.06 and 0.5. 206Pb/238U ages from rounded cores are scattered around c. 720 Ma (analysis 3.1), c. 650 Ma (analyses 1.1 and 2.1), and c. 560 Ma (analyses 5.2 and 6.1). Ten concordant data from oscillatory-zoned zircon define a concordia age of 549 +13 Ma (MSWD ¼ 0.9, probability of concordance ¼ 0.7, Fig. 9a). The age is interpreted as timing magmatic zircon growth in the granite, with rounded zircon cores representing xenocrystic zircon incorporated into the magma. Sample 90112302B. This sample is of hornblende– biotite felsic gneiss from north Balchenfjella, and is host to granitic dykes (sample 90112302A, Fig. 6a). The gneiss is intercalated with mafic


47

Fig. 8. Representative images of zircon grains from samples of the NE and SW Terranes. All images are of cathodoluminescence unless otherwise marked (BSE ¼ backscattered electron). Each SHRIMP analysis spot is labelled with grain analysis number, U content (ppm)/Th/U ratio and 204Pb-corrected 206Pb/238U age with 1s error. Ages that are discordant with 207Pb/235U ages outside 68.3% confidence limits are labelled.

layers and is strongly deformed. It is composed of hornblende, biotite, plagioclase, quartz and K-feldspar, with minor epidote, titanite and opaque minerals. Zircon grains include rounded, equant to elongate forms 50– 400 mm long. Most contain oscillatory- or sector-zoned cores truncated by rims of low-CL zircon with simple or no zoning; grains of low-CL sector zoned zircon are also present (Fig. 8c). Most grains also have thin discontinuous rims of high-CL, unzoned zircon. Oscillatory-zoned cores have U contents between 110 and 620 ppm and Th/U ratios between 0.1 and 0.6 (Table 9). Low-CL rims and sector-zoned grains have highly variable compositions, with 200 –4000 ppm U and Th/U ratios between 0.03 and 0.5. High-CL rims (three analyses) have U contents between 80 and 110 ppm and Th/U ratios from 0.2 to 0.3. 206 Pb/238U ages are scattered between 460 and 1030 Ma (Fig. 9b). Data from oscillatory-zoned cores fall into groups at c. 1000 Ma and c. 800 Ma. Most of these analyses fall into the older group, with four out of seven analyses defining a concordia age of 983 + 17 Ma (MSWD ¼ 1.6, probability of concordance ¼ 0.8). Another three

analyses from oscillatory-zoned cores define a concordia age of 795 + 17 Ma (MSWD ¼ 0.5, probability of concordance ¼ 0.6). Analyses from low-CL rims and sector-zoned grains mostly cluster around 600 Ma, except for four analyses with 206Pb/238U ages scattered between 710 and 790 Ma. These older analyses come from the inner sides of low-CL rims (Fig. 8c), and have lower Th/U ratios (,0.15) than the remaining rim analyses (0.3–0.5). Although the older low-CL rims cannot be distinguished texturally from the rest, the differences in ages and Th/U ratios suggest that they represent a distinct stage of zircon growth at c. 750 Ma. Of the remaining analyses of low-CL rims and grains, seven out of nine data define a concordia age of 605 + 7 Ma (MSWD ¼ 1.5, probability of concordance ¼ 0.7). Analyses from thin high-CL rims are scattered between 460 and 550 Ma. An age of growth cannot be defined, but it postdates the growth of low-CL zircon at c. 600 Ma, which grew during high-grade metamorphism. Cores with oscillatory zircon are detrital, derived from igneous sources with ages of c. 1000 and 800 Ma. The significance of high-U, low Th/U rims with ages of c. 750 Ma is unclear, but may indicate that metamorphic

48

K. SHIRAISHI ET AL.

Fig. 9. Tera– Wasserburg concordia diagrams of U–Pb SHRIMP data. (a) Zircon from 90112302A; (b) zircon from 90112302B; (c) zircon from 9091405A (session A4); (d) titanite from 9091405A.

lithologies of this age were also eroded and incorporated into the sedimentary protolith.

SW Terrane Sample 9091405A. This mylonitized granite from Vengen has been called Vengen granite by Shiraishi et al. (1992) or Vikinghoegda granite by Li et al. (2003), and was intruded into a large shear zone (Main Shear Zone) between meta-tonalite and gneisses in the SW Terrane (Fig. 6c). This is a fine- to medium-grained granite, consisting of plagioclase, quartz, K-feldspar, biotite, muscovite, titanite, zircon, apatite, epidote and Fe–Ti oxides. The bulk composition falls in the alkali granite field on the Na2O þ K2O v. SiO2 diagram (Li et al. 2003). Zircon grains are euhedral to subhedral, prismatic and 50–200 mm in length. Grains have concentric and oscillatory zoning, with decreasing CL from core to rim (Fig. 8d). There is little evidence of modification that could be attributed to mylonitization. Few grains contain irregular cores with truncated growth zoning, which are considered to be

xenocrystic. Excluding these cores, U contents increase outwards, with cores having 160 and 800 ppm U and Th/U ratios between 0.3 and 0.7, and low-CL rims having 800–4000 ppm U and Th/U ratios below 0.1. To compare analytical results between the ANU and NIPR SHRIMP facilities, data were processed separately from two sessions. For session A4 (ANU), five core and four rim analyses from six grains together define a concordia age of 564 + 5 Ma (MSWD ¼ 0.9, probability of concordance ¼ 0.2, Fig. 9c). From session N1 (NIPR), all six concordant data from six grains define a concordia age of 559 + 7 Ma (MSWD ¼ 1.2, probability of concordance ¼ 0.9). Incorporating errors in standard calibration from each session, an average age of 562 + 7 Ma is suggested as timing zircon growth in the granitic magma. The high-U rims are not a distinct stage of growth, but represent U enrichment in the late stage of magmatic crystallization. Concordant ages of c. 1130 Ma (analysis 2.a) and c. 980 Ma (analysis 5.1) from xenocrystic cores give some indication of crustal material incorporated into the granitic melt.


Titanite can be observed in thin-section as pale yellow subhedral to anhedral grains. Grains are dispersed along the mylonitic foliation in aggregates that appear to represent fractured remnants of pre-existing grains. Larger, less fractured grains up to 0.5 mm long preserve rhombic crystal faces, and in contrast to other titanite-bearing samples (90102801A and 9031507) intergrowths of titanite with metamorphic mineral phases are absent. Grains contain 160–320 ppm U and 40–250 ppm Th (Table 11). Data uncorrected for common lead form a roughly linear array above the concordia on the Tera– Wasserburg plot (Fig. 9d). Excluding analysis 1.2, seven data from five grains fall along a Model 1 discordia line with a lower intercept age of 570 + 8 Ma (MSWD ¼ 1.6). The same analyses define a 207Pb-corrected 206Pb/238U mean age of 568 + 11 Ma (MSWD ¼ 1.4). The titanite age overlaps with that obtained from magmatic zircon, and represents magmatic titanite growth, prior to the development of the mylonitic fabric.

49

Sample 85012817. This is a biotite orthogneiss from Vengen. It occurs close to the Main Shear Zone and shows a strong ductile shear fabric. This rock is interpreted as metamorphosed and deformed granite. Constituent minerals are biotite, microcline, quartz and plagioclase, with minor allanite, titanite, secondary chlorite and muscovite. Zircon grains are mostly elongate, 100–400 mm long and subhedral to anhedral. Cores have oscillatory zoning characteristic of magmatic zircon, truncated by unzoned U-rich rims up to 100 mm wide (Fig. 8e). Zircon rims have U contents between 400 and 5500 ppm and Th/U ratios below 0.2, whereas core analyses have U contents mostly below 200 ppm and Th/U ratios between 1.0 and 0.5. (Table 12). Seven concordant data from cores define a concordia age of 653 + 11 Ma (MSWD ¼ 1.4, probability of concordance ¼ 0.9). Concordia ages for zircon rims were calculated separately for the sessions at ANU (Fig. 10a) and NIPR (Fig. 10b). For session A3 (ANU), nine

Fig. 10. Tera –Wasserburg concordia diagrams of U– Pb SHRIMP data. (a) Zircon from 85012817 (session A3); (b) zircon from 85012817 (session N5); (c) zircon from 9031507; (d) titanite from 9031507.

50

K. SHIRAISHI ET AL.

concordant data from rims define a concordia age of 571 + 5 Ma (MSWD ¼ 1.8, probability of concordance ¼ 0.01), whereas for session N5 (NIPR 5), five data loosely define a concordia age of 587 + 10 Ma (MSWD ¼ 1.6, probability of concordance ¼ 0.9). The c. 653 Ma age for zircon cores is interpreted as timing the crystallization of the igneous protolith, with c. 571 Ma rims grown during high-grade metamorphism. Sample 9031507. This is a tonalitic hornblende orthogneiss from Mefjell. The gneiss is considered to be the eastern extension of meta-tonalite in the southwestern area, but less deformed and preserves a massive plutonic fabric (Fig. 6d). It is composed of plagioclase, quartz, hornblende, biotite, epidote, titanite, apatite and zircon. Zircon occurs as a mixture of elongate prismatic grains up to 300 mm long with rounded edges, and irregular fragments. Internal zoning is euhedral, prismatic and oscillatory, with no inherited cores. Annealed fractures occur in several grains, which merge with very thin, discontinuous rims of unzoned zircon (Fig. 8f). Most grains also have extremely thin (less than 5 mm) rims and (rarely) patches of very high-CL zircon. Oscillatory-zoned zircon has ,320 ppm U and Th/U ratios between 0.3 and 0.7, whereas high-CL rims and patches have U and Th contents of 10 ppm or less (Table 13). Nine concordant age data from six grains define a concordia age of 920 + 8 Ma (MSWD ¼ 1.3, probability of concordance ¼ 0.2), interpreted as the time of crystallization of the tonalitic protolith. Concordant age data from three out of four rim analyses are not significantly different from the magmatic age, and are interpreted as minor annealing and recrystallization of the igneous zircon, rather than growth of new zircon during metamorphism. Titanite is an abundant accessory mineral in the orthogneiss, occurring as grains up to 300 mm long intergrown with hornblende and biotite in a metamorphic assemblage. Grains separated show little compositional zoning, and contain up to 200 ppm U and 60 ppm Th (Table 14). Data uncorrected for common lead cluster above the concordia on the Tera–Wasserburg plot (Fig. 10d). Excluding three analyses, carried out on areas of titanite with micro-inclusions of unknown silicates, 11 data from 11 grains define a 207Pb-corrected 206Pb/238U mean age of 517 + 5 Ma (MSWD ¼ 1.0). As with titanite in sample 90102801A, the result provides a minimum age for the growth of metamorphic titanite.

Summary of SHRIMP dating The SHRIMP results are complex, and the interpretation requires an approach that integrates information from textures and compositions of zircon

with structural and compositional relationships in the host lithologies. Zircon and titanite age populations identified in all 11 samples, and the magmatic or metamorphic origin of zircon as determined from grain textures and compositions, are summarized in Figure 11. In the NE Terrane, the majority of samples are from localities of gneisses that are highly deformed and that do not preserve predeformational lithological contacts. Gneiss samples 84022004, 85020401C, 9032401A and 90112102A are garnet-bearing and aluminous, characteristic of a sedimentary origin, and sample 90112302B comes from heterogeneous laminated gneiss interlayered with aluminous and calcareous metasedimentary rocks, and is likely to have a volcaniclastic protolith. Sample 90102801A of hornblende – biotite gneiss is also compositionally layered, contains lenses of calcareous composition, and occurs in a sequence of aluminous gneisses, and therefore is possibly volcaniclastic. Sample 85011503D is interpreted as a metamorphosed igneous enderbite, consistent with the compositionally uniform nature of the lithology in outcrop, which was emplaced at c. 951 Ma. Pre-700 Ma zircon cores and grains are mostly of magmatic origin, and provide a detrital age signature for the terrane, with magmatic activity in three discrete time windows at c. 1130 Ma, c. 1000 Ma and c. 800 Ma (Fig. 11). This simple detrital signature may indicate a local derivation of sedimentary detritus, from synsedimentary volcanism or basement lithologies. Excluding one detrital magmatic zircon with an age of c. 3200 Ma, the lack of pre-1200 Ma ages in the detrital signature indicates a relatively juvenile provenance for the sediments, with no significant contribution from early Proterozoic or Archaean continental crust. Meso-Neoproterozoic sedimentation is also proposed by various researchers from other East Gondwana crustal fragments, such as southern India (Santosh et al. 2006; Collins et al. 2007). Magmatic zircon in the metasedimentary rocks must be detrital in nature, deposited prior to the formation of the gneisses through granulite-grade metamorphism. The presence of c. 800 Ma magmatic zircon in three samples of metasedimentary rocks therefore constrains subsequent metamorphism to a younger age. There is a high productivity of metamorphic zircon (in five samples) within a window between 640 and 600 Ma, with most age populations at c. 600 Ma. This is considered to be the time of zircon growth at or near peak metamorphism and deformation. Evidence for previous metamorphic events from zircon analysis is more ambiguous. Two samples of metasedimentary rocks have high-U, low Th/U overgrowths on


51

Fig. 11. Summary of SHRIMP zircon and titanite U– Pb ages from the Sør Rondane Mountains.

magmatic zircon cores with ages of c. 750 Ma, and may be attributed to metamorphism. However, the textural relationships within zircon grains are not clear enough to say whether this zircon represents metamorphism of the host metasedimentary rocks, or of source rocks that provided detritus to sediments deposited after 750 Ma. Because of the rarity of this generation of zircon in the analysed samples in general, and the lack of field evidence for deformation and metamorphism prior to the formation of gneisses at c. 600 Ma, the hypothesis that 750 Ma zircon is detrital is tentatively proposed. However, there is no textural and morphological evidence. Analyses of high-U, low-Th/U zircon that might be attributed to metamorphic growth at c. 1000 Ma or c. 1130 Ma are scant, being restricted to isotopically discordant analyses of zircon rims from sample 90102801A and a xenocrystic zircon core from sample 9032401A. The formation of high-U, low-Th/U rims can also occur in magmatic zircon, where fractional crystallization concentrates U in late-stage magmatic fluids, and this mechanism is invoked to explain euhedral high-U rims in magmatic zircon from samples 90112302A and 9091405A. Consequently, the few U-rich analyses with c. 1000 Ma ages do not represent substantial evidence of high-grade metamorphism at this time.

Zircon growth after c. 600 Ma is also very limited, being restricted to a minor population of c. 560 Ma metamorphic zircon from sample 85020401C, and a few isotopically discordant analyses from thin high-CL rims that are present in many samples. The latter may be attributed to marginal recrystallization and U leaching by hydrothermal fluids after the metamorphic event (Geisler et al. 2003). Magmatic zircon from sample 90112302A dates the intrusion of nebulitic granitic dykes at c. 550 Ma. The relatively undeformed nature of these dykes demonstrates that high-strain deformation had waned by this time. Zircon dating results from the SW Terrane have some important differences from those from the NE Terrane. All three samples derive from metamorphosed and deformed igneous lithologies, with each timing a separate stage of magmatism at c. 920 Ma, c. 650 Ma and c. 560 Ma. The latter age derives from granite that intrudes across a gneissic fabric that is also present in the c. 650 Ma orthogneiss, and therefore constrains the timing of high-strain deformation between 650 and 560 Ma. Metamorphic zircon growth is recognized in sample 85012817 only, and the c. 570 Ma age may coincide with near-peak (amphibolite-grade) metamorphism and deformation in the orthogneiss. If this is the case, it suggests different metamorphic histories for the

52

K. SHIRAISHI ET AL.

SW and NE Terranes, prior to their current juxtaposition by later faults and/or shear zones. The age of mylonitization in association with the Main Shear Zone, which juxtaposes the c. 920 Ma meta-tonalites with amphibolite-grade gneisses in the SW Terrane, is constrained by the c. 560 Ma intrusive age of the sheared granite at Vengen (sample 85012817), and preservation of the intrusive age in titanite from the same sample suggests that this area was not subject to thermal metamorphism after this time. In contrast, titanite ages of c. 517 Ma were obtained from both the metatonalite in the SW Terrane and hornblende– biotite gneiss from the NE Terrane. Both of these localities (Mefjell and Isachsenfjella) lie in the vicinity of large bodies of post-tectonic granite (Li et al. 2003). The c. 517 Ma titanite ages represent either new titanite growth or isotopic resetting through Pb loss, under elevated thermal conditions that may be a result of cooling from an earlier metamorphic peak or of contact metamorphism caused by the emplacement of granitic plutons, as suggested by Asami et al. (1992). In either case, the ages indicate that the NE and SW Terranes share a common history by c. 517 Ma.

Nd model ages Samples and procedures for Nd isotopic analysis Nd model ages have been used by many workers for revealing histories of crustal genesis (e.g. DePaolo 1988; Dickin 1995; Stern 2002; Kagami et al. 2006). In this study we have compiled 180 Nd isotope data from previous studies of central to eastern Dronning Maud Land and Enderby Land, along with 31 new data, mainly from the LHC (Table 15). The new data are derived from samples of paragneiss and basic to intermediate orthogneiss. Quartzofeldspathic gneisses and granites are also included. Some paragneisses have migmatitic textures. Nd isotope analytical procedures follow those of Kagami et al. (1987, 2006). Isotope analyses were performed on a MAT261-type mass spectrometer equipped with five Faraday cups at Niigata University. 143 Nd/144Nd ratios were normalized to 146 Nd/144Nd ¼ 0.7219. 143Nd/144Nd ratios are reported relative to 143Nd/144Nd ¼ 0.511858 for La Jolla or 0.512640 for BCR-1. Sm and Nd concentrations were measured by the isotope dilution method using a 149Sm – 150Nd mixed spike. We estimate an error of 0.1% for the Sm/Nd ratio of each sample based on reproducibility of the data. Depleted mantle Nd model ages (TDM) in this study were calculated using an Excel

spreadsheet provided by Stern (2002), in which the model in Nelson & DePaolo (1984) and DePaolo (1988) was applied. Most of the previous Nd isotope studies (Shiraishi & Kagami 1992; Owada et al. 1994, 2001; Shiraishi et al. 1995, 1997b; Yoshida et al. 1999; Suzuki et al. 2001, 2006; Nishi et al. 2002; Ajishi et al. 2004; Kawano et al. 2005) were performed at the same laboratory as the present study, and 143 Nd/144Nd ratios are reported relative to 143 Nd/144Nd ¼ 0.511858 (La Jolla), with the same isotopic parameters being used. The 147Sm/144Nd ratios in Table 15 vary widely across the region. Higher 147Sm/144Nd ratios are regarded to yield unreliable Nd model ages (e.g. Stern 2002). In this study, data interpretation was restricted to samples with147Sm/144Nd ratios of 0.13 or less, representing 96 analyses out of a total of 180 (Table 15, Fig. 12). Following Stern (2002), analyses with147Sm/144Nd ratios of 0.165 or less represent more than 87% (156 data) of the total dataset.

Regional distribution of Nd model ages A compilation of depleted mantle Nd model ages (TDM), including new analyses and those from previous works recalculated following Stern (2002), provides new insights into crustal residence time in each terrane (Fig. 13). TDM values for the Napier Complex were reported from four areas: Mt. Riiser-Larsen (3.03–3.43 Ga), Tonagh Island (3.10– 3.71 Ga), Mt. Sones (3.43–3.95 Ga) and Fyfe Hills (3.27–3.52 Ga). Tonalitic orthogneiss from Mt. Sones, which has yielded the oldest age in the Napier Complex, also provides the oldest TDM (3.9 Ga) (Black et al. 1986). TDM values for 12 para- and orthogneisses from the Rayner Complex vary widely from 1.27 to 2.28 Ga, with no indications of an Archaean crustal component. In particular, orthogneisses from Sandercock Nunataks, located inland to the south of the Napier Complex, have consistent 1.61 –1.71 Ga ages, postdating the latest pervasive metamorphic and magmatic event at 2.5 Ga in Napier Complex, and suggesting the presence of late Palaeoproterozoic juvenile crust. The only indication of an Archaean component in the RC comes from a single c. 2.6 Ga xenocrystic zircon core in pelitic gneiss, from Mt. Underwood in the Nye Mountains (Shiraishi et al. 1997b). This is significantly older than the c. 1.7 Ga TDM of the same gneiss. TDM values for 37 samples from the LHC and YBC also present a wide range of 0.87 –2.70 Ga, with a major mode at 1.0–1.25 Ga and a small mode at 2.29–2.70 Ga. The latter group derives from both ortho- and paragneisses from the

Table 15. Nd model ages for the rocks from eastern Dronning Maud Land and western Enderby Land, East Antarctica. No.

Sample

Enderbitic gneiss Enderbitic gneiss Enderbitic gneiss Garnet granite Enderbitic gneiss Enderbitic gneiss Enderbitic gneiss Retrograde gneiss Retrograde gneiss Retrograde gneiss Retrograde gneiss Granite Granite Granite Granite Granite Granite Granite Bt qtzfelds gneiss Bt qtzfelds gneiss Bt qtzfelds gneiss Bt qtzfelds gneiss Px–pl gneiss

Locality

Brattnipene Brattnipene Brattnipene Brattnipene Brattnipene Brattnipene Brattnipene Brattnipene Brattnipene Brattnipene Brattnipene Dufek Dufek Dufek Dufek Mejell Mejell Pingvinane Balchen Balchen Balchen Balchen Balchen

Nd Sm (ppm) (ppm)

147

Sm/144Nd

143

Nd/144Nd T(DM) Ga

4.93 3.14 2.73 14.50 3.44 2.37 3.31 1.22 3.77 6.56 4.86 6.91 10.71 9.80 7.58 3.02 18.77 15.65 1.55 2.90 4.81 1.80 3.67

20.03 12.19 12.00 96.40 13.64 12.50 14.60 4.28 14.48 27.69 20.42 44.70 69.11 71.04 61.32 18.54 197.38 84.31 7.48 18.70 14.20 10.90 12.80

0.1488 0.1557 0.1375 0.0909 0.1525 0.1146 0.1371 0.1723 0.1574 0.1432 0.1439 0.0934 0.0937 0.0834 0.0747 0.0985 0.0575 0.1122 0.1253 0.0937 0.2048 0.0998 0.1733

0.512560 0.512620 0.512490 0.512273 0.512540 0.512360 0.512470 0.512720 0.512620 0.512550 0.512530 0.512302 0.512297 0.512278 0.512230 0.512306 0.512101 0.512381 0.512130 0.512180 0.512750 0.512110 0.512770

1.13 1.11 1.11 0.96 1.25 1.05 1.14 1.14 1.07 1.12 0.94 0.95 0.90 0.90 0.98 0.92 1.00 1.57 1.10 1.25

References

Shiraishi & Kagami 1992 Shiraishi & Kagami 1992 Shiraishi & Kagami 1992 This study Shiraishi & Kagami 1992 Shiraishi & Kagami 1992 Shiraishi & Kagami 1992 Shiraishi & Kagami 1992 Shiraishi & Kagami 1992 Shiraishi & Kagami 1992 Shiraishi & Kagami 1992 Arakawa et al. 1994 Arakawa et al. 1994 Arakawa et al. 1994 Arakawa et al. 1994 Arakawa et al. 1994 Arakawa et al. 1994 Arakawa et al. 1994 Grew et al. 1992 Grew et al. 1992 Grew et al. 1992 Grew et al. 1992 Grew et al. 1992

Yamato–Belgica Complex 24 Y80A529 Opx–bt gneiss 25 A79121511 Bt gneiss

Yamato Mts. Belgica Mts.

4.10 1.14

23.36 8.81

0.1061 0.0782

0.512101 0.512214

1.33 0.94

This study This study

Lu¨tzow-Holm Complex 26 80S74 27 80S15 28 80S52 29 80S57 30 80S59

Sinnan Rock Sinnan Rock Sinnan Rock Sinnan Rock Sinnan Rock

2.70 5.40 3.00 2.30 5.20

11.35 25.70 15.70 9.80 26.10

0.1438 0.1270 0.1155 0.1419 0.1204

0.512341 0.512289 0.512229 0.512335 0.512239

1.52 1.32 1.26 1.49 1.31

This study This study This study This study This study

Sil–bt –grt gneiss Bt amphibolite Bt granite Grt–bt gneiss Bt gneiss

53

(Continued )


Sør Rondane Mountains 1 85011503B 2 85011503C 3 85011503D 4 85011504B 5 85011602D 6 9022502A 7 9022502B 8 85011601A 9 85011601C 10 85011602A 11 85011602B 12 B9001-2301A 13 2302A 14 2303A 15 2305B 16 2502 17 2405C 18 1406 19 EG88011109 20 EG88011212 21 EG88011319 22 EG88012804 23 EG88012105

Rock type

No.

80S78 73123106 73123103 73123116 74010606 74010701 73123106K 74010105 74010115 74010107 74010113 74010304 No.1 No.2 No.3 No.4 No.5 No.6 K95010804 K950110m2 K950110m3 K950110m4 K950110m6 K950110m7 K950110m8 K950110m9 K950110m10 0101-2 0102A 0102B 0107 1802 95020203 95020205 o3020512 95020301 95020303 o3020616

Rock type Bt gneiss Meta-trondjemite Meta-trondjemite Meta-trondjemite Meta-trondjemite Meta-trondjemite Meta-trondjemite Meta-trondjemite Meta-trondjemite Meta-trondjemite Meta-trondjemite Sil–bt –grt gneiss Bt gneiss Granite Granite Granite Granite Granite Granite Hbl–bt gneiss Hbl–bt gneiss Hbl–bt gneiss Hbl–bt gneiss Hbl–bt gneiss Hbl–bt gneiss Hbl–bt gneiss Hbl–bt gneiss Bt–hbl gneiss Bt–hbl gneiss Qtzfeld gneiss Bt gneiss Grt–bt gneiss Granite Granite Granite Granite Granite Granite

Locality Sinnan Rock Cape Hinode Cape Hinode Cape Hinode Cape Hinode Cape Hinode Cape Hinode Cape Hinode Cape Hinode Cape Hinode Cape Hinode Cape Hinode Kasumi Rock Kasumi Rock Kasumi Rock Kasumi Rock Kasumi Rock Kasumi Rock Oku-iwa Rock Oku-iwa Rock Oku-iwa Rock Oku-iwa Rock Oku-iwa Rock Oku-iwa Rock Oku-iwa Rock Oku-iwa Rock Oku-iwa Rock Nesoya Nesoya Nesoya Nesoya Nesoya East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul

Nd Sm (ppm) (ppm) 3.80 0.72 1.67 1.18 3.60 0.90 1.00 0.23 0.52 2.39 2.51 5.65 3.55 0.27 0.25 0.12 0.23 0.41 22.80 7.37 6.48 4.23 3.94 5.19 8.79 4.33 1.50 10.92 10.05 1.53 3.34 6.24 14.77 12.30 0.95 20.25 11.54 8.36

19.10 4.80 7.86 5.70 14.54 5.84 5.72 2.06 3.30 11.40 11.62 27.98 23.90 1.00 0.95 0.55 0.82 1.18 163.00 26.00 22.90 18.40 23.00 29.30 33.30 18.10 9.39 50.61 45.57 7.01 22.31 35.75 68.83 57.10 5.74 89.61 75.80 36.33

147

Sm/144Nd 0.1203 0.0907 0.1284 0.1251 0.1497 0.0932 0.1057 0.0675 0.0953 0.1267 0.1306 0.1220 0.0897 0.1651 0.1615 0.1332 0.1713 0.2095 0.0845 0.1714 0.1711 0.1390 0.1036 0.1071 0.1596 0.1446 0.0966 0.1304 0.1333 0.1319 0.0905 0.1055 0.1297 0.1302 0.1000 0.1366 0.092 0.1391

143

Nd/144Nd T(DM) Ga

0.512249 0.512173 0.512383 0.512384 0.512510 0.512141 0.512220 0.511985 0.512159 0.512390 0.512427 0.511943 0.512139 0.512303 0.512307 0.512230 0.512258 0.512475 0.511928 0.512680 0.512679 0.512533 0.512378 0.512400 0.512820 0.512553 0.512263 0.512372 0.512376 0.512407 0.512202 0.512315 0.512402 0.512474 0.512286 0.512435 0.512231 0.512438

1.29 1.08 1.18 1.13 1.26 1.14 1.16 1.10 1.14 1.14 1.13 1.82 1.11 2.15 1.53 1.31 1.04 0.92 0.92 0.65 1.08 1.02 1.23 1.26 1.18 1.04 1.03 1.16 1.04 1.02 1.20 1.02 1.23

References This study Shiraishi et al. 1995 Shiraishi et al. 1995 Shiraishi et al. 1995 Shiraishi et al. 1995 Shiraishi et al. 1995 Shiraishi et al. 1995 Shiraishi et al. 1995 Shiraishi et al. 1995 Shiraishi et al. 1995 Shiraishi et al. 1995 Shiraishi et al. 1995 Ajishi et al.2004 Ajishi et al.2004 Ajishi et al.2004 Ajishi et al.2004 Ajishi et al.2004 Ajishi et al.2004 Nishi et al. 2002 Nishi et al. 2002 Nishi et al. 2002 Nishi et al. 2002 Nishi et al. 2002 Nishi et al. 2002 Nishi et al. 2002 Nishi et al. 2002 Nishi et al. 2002 This study This study This study This study This study Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005

K. SHIRAISHI ET AL.

31 32 33 34 35 36 37 38 39 40 41 42 43 44 45 46 47 48 49 50 51 52 53 54 55 56 57 58 59 60 61 62 63 64 65 66 67 68

Sample

54

Table 15. Continued

Rayner 105 106 107

o3020608 95020308 o3020504 o3020506 95020305 95020402 o3020308 95020405 o3020601 o3020517 79022008 79022016 79100201 79100202 950205M6 81012802 Y70012905 68031601 Y69101014 Y69101023 A2 A3 A4 A5 2601 2602 2702 C1 C4 C7 20111A Y69020613 Y69020614 Y70020515 RH19B 84011105 Complex MA88021608 78285009 78285010

Granite Granite Granite Granite Granite Granite Granite Granite Leucosome in Hbl gn Leucosome in Hbl gn Bt–opx amphibolite Bt–hbl–opx gneiss Opx granulite Opx–hbl granulite Metamorphic rock Sil–bt –grt gneiss Grt gneiss Noritic charnockite Charnockitic band Granitic gneiss Bt–grt gneiss Bt–grt gneiss Grt–opx –bt granulite Grt–opx –bt gneiss Grt–two –px–bt granulite Grt–qtzfels gneiss Grt–qtzfels gneiss Grt–bt gneiss Px granulite Grt–bt gneiss Sil–bt –grt gneiss Metabasite Grt dioritic gneiss Hbl charnockite Sil–bt –grt gneiss Hbl–bt gneiss

East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul East Ongul West Ongul Ongul Strait Utholmen Fleynoya Fleynoya Skarvsnes Skarvsnes Skarvsnes Skarvsnes Skarvsnes Skarvsnes Skarvsnes Skarvsnes Skarvsnes Skarvsnes Telen Skallen Skallen Skallevikhalsen Rundvagshetta Innhovde

10.31 16.83 15.51 12.81 15.62 10.40 17.45 2.70 9.40 2.61 0.78 5.33 1.65 3.57 5.77 3.15 5.15 7.69 7.91 10.20 5.57 6.39 14.15 5.56 9.40 4.25 10.20 1.09 14.23 5.24 1.87 7.76 1.43 9.69 3.46 6.13

45.41 79.30 76.95 58.06 62.28 57.20 86.29 9.74 43.94 28.58 3.81 22.16 6.40 13.90 41.85 12.09 21.10 35.70 46.40 41.70 21.22 26.72 56.60 20.59 37.21 16.57 52.42 4.68 65.20 25.80 7.11 29.90 9.38 50.80 21.48 30.37

0.1372 0.1283 0.1218 0.1334 0.1516 0.1099 0.1222 0.1676 0.1293 0.0552 0.1238 0.1454 0.1559 0.1553 0.0833 0.1575 0.1475 0.1302 0.1030 0.1479 0.1587 0.1445 0.1511 0.1631 0.1528 0.1552 0.1176 0.1412 0.1319 0.1228 0.1588 0.1569 0.0921 0.1153 0.0973 0.1220

0.512375 0.512439 0.512383 0.512396 0.512525 0.512233 0.512361 0.512410 0.512359 0.512142 0.512383 0.512548 0.512500 0.512555 0.512295 0.512420 0.512519 0.512377 0.512211 0.512387 0.512481 0.512363 0.512499 0.512502 0.512509 0.512405 0.512246 0.512400 0.512340 0.512290 0.511605 0.512000 0.511220 0.511310 0.511121 0.512224

1.33 1.08 1.10 1.23 1.27 1.19 1.14 1.23 0.87 1.12 1.11 1.42 1.27 0.88 1.68 1.20 1.21 1.15 1.52 1.55 1.49 1.32 1.62 1.33 1.65 1.26 1.33 1.29 1.25 4.22 2.84 2.29 2.70 2.53 1.23

Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 Kawano et al. 2005 This study Yoshida & Kagami 1995 Yoshida & Kagami 1995 Yoshida & Kagami 1995 Yoshida & Kagami 1995 This study This study This study This study This study This study This study Tanaka et al. 1985 Tanaka et al. 1985 Tanaka et al. 1985 This study Yoshida et al. 1999 Yoshida et al. 1999 Yoshida et al. 1999 This study This study

Pelitic gneiss Paragneiss Pegmatite

Mt. Vechernaya Mt. Underwood Mt. Underwood

5.66 6.43 36.20

17.10 33.40 215.00

0.2001 0.1163 0.1017

0.511996 0.511175 0.510791

1.67 1.98

Shiraishi et al. 1997b Black et al. 1987 Black et al. 1987 55

(Continued )


69 70 71 72 73 74 75 76 77 78 79 80 81 82 83 84 85 86 87 88 89 90 91 92 93 94 95 96 97 98 99 100 101 102 103 104

Sample

108 109 110 111 112 113 114 115 116 117 118 119 120

78285023 78285024 78285027 80285043B 80285043M 77283498 77283554 45121408 45121504 45121505 45121507 451216-0 45121506

Napier 121 122 123 124 125 126 127 128 129 130 131 132 133 134 135 136 137 138 139 140 141 142

Complex SS97021307 SS97021208-1 SS97021303B SS96122803B-1 A90021603B A90021604G 21601G 21602A 21602AB 21602AW 21602B 21602C 21603C 21603E 21603H 21603I 21603N 21603G 76283267 77283464 77283465 77283466

Rock type

Locality

Nd Sm (ppm) (ppm)

147

Sm/144Nd

143

Nd/144Nd T(DM) Ga

References

Granitic orthogneiss Granite Gr orthogneiss Anorthosite layer Gabbroic layer Tonalitic orthogneiss Granitic orthogneiss Opx–grt gneiss Grt–biotite gneiss Opx–grt gneiss Fine grt –biotite gneiss Opx–grt gneiss Opx–grt gneiss

Mt. Fletta Condon Hills Thala Hills Amphitheatre Lakes Amphitheatre Lakes Ward Nunataks Mt. Underwood Sandercock Nunataks Sandercock Nunataks Sandercock Nunataks Sandercock Nunataks Sandercock Nunataks Sandercock Nunataks

6.26 6.61 3.82 1.73 1.43 2.92 3.64 7.66 10.20 8.86 7.52 10.20 15.11

34.70 48.20 20.90 10.70 8.00 12.30 21.10 33.50 54.60 47.40 38.40 54.30 63.52

0.1090 0.0829 0.1104 0.0977 0.1080 0.1435 0.1043 0.1382 0.1129 0.1130 0.1184 0.1135 0.1438

0.510682 0.510462 0.510710 0.510692 0.510716 0.511311 0.511326 0.512039 0.511943 0.511977 0.511991 0.511916 0.512080

2.28 2.08 2.27 2.04 2.26 2.02 1.27 2.03 1.66 1.61 1.68 1.71 2.10

Black et al. Black et al. Black et al. Black et al. Black et al. Black et al. Black et al. This study This study This study This study This study This study

Granitic gneiss Psammitic gneiss Mafic granulite Sapphirine –quartz gneiss Grt–felsic gniess Fe–rich grt–px gneiss Mafic gneiss Mafic gneiss Mafic gneiss Mafic gneiss Mafic gneiss Mafic gneiss Mafic gneiss Felsic gneiss Felsic gneiss Felsic gneiss Ultramafic gneiss Ultramafic gneiss Paragneiss Leucogneiss Grt–bg. gneiss Metapelite

Mt. Riiser-Larsen Mt. Riiser-Larsen Mt. Riiser-Larsen Mt. Riiser-Larsen Tonagh Island Tonagh Island Tonagh Island Tonagh Island Tonagh Island Tonagh Island Tonagh Island Tonagh Island Tonagh Island Tonagh Island Tonagh Island Tonagh Island Tonagh Island Tonagh Island Mount Sones Mount Sones Mount Sones Mount Sones

32.00 3.68 2.45 6.90 8.89 1.61 3.12 11.10 12.70 3.96 5.16 5.14 1.71 1.66 1.18 3.34 1.50 1.61 5.84 1.10 1.96 61.61

162.00 22.60 7.01 50.40 46.20 7.45 12.40 54.70 59.90 21.80 21.10 21.10 7.99 10.00 11.60 27.90 5.63 7.45 31.89 6.27 15.29 96.39

0.1194 0.0984 0.2113 0.0827 0.1163 0.1306 0.1521 0.1226 0.1281 0.1098 0.1478 0.1472 0.1293 0.1003 0.0615 0.0723 0.1610 0.1306 0.1106 0.1060 0.7740 0.3866

0.510961 0.510697 0.512756 0.510464 0.510914 0.510944 0.511498 0.510977 0.511015 0.510660 0.511805 0.511585 0.511109 0.510609 0.509972 0.510128 0.511576 0.510944 0.50996 0.50986 0.50893 0.51433

3.43 3.14

Suzuki et al. 2006 Suzuki et al. 2006 Suzuki et al. 2006 Suzuki et al. 2001 Owada et al. 2001 Owada et al. 2001 Owada et al. 1994 Owada et al. 1994 Owada et al. 1994 Owada et al. 1994 Owada et al. 1994 Owada et al. 1994 Owada et al. 1994 Owada et al. 1994 Owada et al. 1994 Owada et al. 1994 Owada et al. 1994 Owada et al. 1994 Black & McCulloch Black & McCulloch Black & McCulloch Black & McCulloch

3.03 3.39 3.99 3.53 3.71 3.56 2.90 3.43 3.58 3.32 3.10 3.18 3.97 3.44 3.43 3.77

1987 1987 1987 1987 1987 1987 1987

K. SHIRAISHI ET AL.

No.

56

Table 15. Continued

1987 1987 1987 1987

77283467 78285007-A 78285001-F 78285007-J 78285008-9 M N J1 J5 50 53 54 51 56 57

Mafic granulite Tonalitic orthogneiss Tonalitic orthogneiss Tonalitic orthogneiss Paragneiss Charnockite Charnockite Charnockite Charnockite Leuconorite Leuconorite Leuconorite Gabbro Gabbro Gabbro

Central Dronning Maud Land 158 J1704 Felsic gneiss 159 J1838 Felsic gneiss 160 J1671 Felsic gneiss 161 J1795 Felsic gneiss 162 J1736 Augen gneiss 163 J1797 Augen gneiss 164 J1698 Metagranodiorite 165 J1695 Metagranodiorite 166 SR39B/16 167 SR7W/23 Metagabbro 168 SR4/23 S-type granite 169 MS4/23 S-type granite 170 MS2/23 Metaquartzite 171 MS5/23 Calc-silicate rock 172 SR 17/14 Metanoritic dyke 173 SR 28D/16 Metanoritic dyke 174 13A-2/16 Enderbitic gneiss Enderbitic gneiss 175 13A-4/16 Enderbitic gneiss 176 13A-6/16 177 13A-8/16 Enderbitic gneiss 178 13A-9/16 Enderbitic gneiss 179 13A-10/16 Enderbitic gneiss 180 13A-0/16 Enderbitic gneiss

Mount Sones Mount Sones Mount Sones Mount Sones Mount Sones Fyfe Hills Fyfe Hills Fyfe Hills Fyfe Hills Fyfe Hills Fyfe Hills Fyfe Hills Fyfe Hills Fyfe Hills Fyfe Hills

6.66 3.49 4.36 3.97 5.81 1.81 0.82 1.70 0.53 1.73 1.34 1.15 10.05 4.99 5.49

26.87 24.42 29.24 30.64 20.42 15.72 6.11 10.37 4.65 13.37 7.49 5.50 46.92 20.61 29.84

0.1498 0.0863 0.0901 0.0783 0.1719 0.0695 0.0811 0.0990 0.0688 0.0782 0.1081 0.1263 0.1294 0.1463 0.1112

0.51057 0.50901 0.50912 0.50886 0.51084 0.50913 0.50936 0.50979 0.50917 0.50939 0.50997 0.51032 0.51034 0.51067 0.51007

Orvinfjella Wohlthatmassiv Orvinfjella Wohlthatmassiv Orvinfjella Wohlthatmassiv Orvinfjella Wohlthatmassiv Orvinfjella Wohlthatmassiv Orvinfjella Wohlthatmassiv Orvinfjella Wohlthatmassiv Orvinfjella Wohlthatmassiv Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills Schirmacher Hills

10.51 46.8 9.67 31.3 15.68 77.5 8.355 34.32 8.991 41.03 10.14 58.42 4.759 17.04 8.355 34.32 0.698 1.96 12.3 40.2 8.93 38.2 12.8 62.4 5.58 27.2 15.8 105 11.50 178.00 6.16 39.40 13.4 46.7 5.81 28 11.5 44.7 10.1 58.3 6.77 27.3 7.92 31.3 9.97 36.34

0.1352 0.1860 0.1218 0.1568 0.132 0.1046 0.1681 0.1466 0.2157 0.1850 0.1412 0.1239 0.1237 0.0907 0.0594 0.1435 0.1735 0.1254 0.1555 0.1047 0.1499 0.153 0.1659

0.512145 0.512708 0.512294 0.512314 0.512165 0.512348 0.512468 0.512278 0.512884 0.512765 0.512391 0.512491 0.512202 0.512141 0.511884 0.512791 0.512538 0.51233 0.512487 0.51228 0.512426 0.512462 0.512461

3.33 3.36 3.31 3.27 3.25 3.33 3.42 3.52 3.67 3.28

1.72 1.22 1.62 1.61 0.95 1.71 1.33 0.92 1.40 1.10

1.21 1.42 1.05 1.44 1.42

Black & McCulloch 1987 Black & McCulloch 1987 Black & McCulloch 1987 Black & McCulloch 1987 Black & McCulloch 1987 McCulloch & Black 1984 McCulloch & Black 1984 McCulloch & Black 1984 McCulloch & Black 1984 McCulloch & Black 1984 McCulloch & Black 1984 McCulloch & Black 1984 McCulloch & Black 1984 McCulloch & Black 1984 McCulloch & Black 1984

Jacobs et al. 1998 Jacobs et al. 1998 Jacobs et al. 1998 Jacobs et al. 1998 Jacobs et al. 1998 Jacobs et al. 1998 Jacobs et al. 1998 Jacobs et al. 1998 Ravikant, 2007 Ravikant, 2007 Ravikant, 2007 Ravikant, 2007 Ravikant, 2007 Ravikant, 2007 Ravikant, 2006 Ravikant, 2006 Ravikant, 2004 Ravikant, 2004 Ravikant, 2004 Ravikant, 2004 Ravikant, 2004 Ravikant, 2004 Ravikant, 2004

147

Sm/144Nd , 0.13; remaining ages are 0.13

57

Rock types follow the original references. Depleted mantle Nd model ages (TDM), recalculated following Stern (2002). Bold numerals indicate TDM with 147 Sm/144Nd , 0.165. Mineral abbreviations are after Kretz (1983). Place names are shown in in Figures 2, 3 and 13.

4.18 3.95 3.94 3.88


143 144 145 146 147 148 149 150 151 152 153 154 155 156 157

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K. SHIRAISHI ET AL.

Fig. 12. Histogram of depleted mantle model ages for the six terranes of eastern Dronning Maud Land and Enderby Land. Only samples yielding 147Sm/144Nd , 0.13 are plotted. Data are provided in Table 15.

southern Soˆya Coast region (Skallen, Skallevikshalsen and Rundva˚gshetta; see Fig. 3). Yoshida et al. (1999) and Suda et al. (2008) have suggested the presence of late Archaean to early Palaeoproterozoic crust in the southern part of the LHC. Such crust may represent an ancient basement to supracrustal lithologies, and/or source lithologies of detritus from the hinterland. TDM values of 12 samples from the SRM concentrate in a narrow range of 0.90–1.25 Ga, except one quartzofeldspathic gneiss (1.57 Ga) from the eastern part of the mountains. TDM values (1.0–1.1 Ma) for enderbitic gneiss in the

NE Terrane are only slightly older than the time of magmatism, estimated from SHRIMP zircon (951 + 17 Ma, this study), Sm –Nd whole-rock isochron (961 + 101 Ma) and Rb–Sr whole-rock isochron (978 + 52 Ma) ages (Shiraishi & Kagami 1992). TDM values of less than 1.0 Ga were obtained from younger granites in the SRM. The significance of 1.57 Ga gneiss from the eastern area (Balchenfjella) is not clear (Grew et al. 1992). Pan-African post-orogenic granites have only slightly younger TDM values(0.9–1.0 Ga) than orthogneisses (as recalculated from Arakawa et al. 1994). Although the number and variety of samples are limited, the

Fig. 13. Distribution of TDM in east Dronning Maud Land. Only samples with 147Sm/144Nd , 0.13 are plotted. Data sources are given in Table 15.


data indicate that the majority of the SRM is built on a basement of c. 1100–1000 Ma juvenile crust. This is in a good agreement with c. 1130 and c. 1000 Ma signatures from magmatic zircon in orthogneisses and detrital zircon in paragneisses, and supports the observation that there is no significant contribution from Palaeoproterozoic or Archaean continental crust. To the west of the SRM, TDM values reported by Jacobs et al. (1998), Ravikant et al. (2004, 2007) and Ravikant (2006) are consistently Mesoproterozoic. Recalculation in this study yields slightly younger results, with TDM values ranging from 1.4 to 0.9 Ga that are comparable with those obtained in the SRM. Thus the central Dronning Maud Land crust formed at similar times to that of the SRM.

Discussion and conclusions Tectonothermal events in the Sør Rondane Mountains Integrating the results of the SHRIMP dating and depleted mantle Nd model ages (TDM) with previous geochronological studies, tectonothermal events in the Sør Rondane Mountains and other terranes in central –eastern Dronning Maud Land are summarized in Table 16. Coincident ages of extensive magmatism and TDM values suggest that the basement crust of the SRM formed during the late Mesoproterozoic, as pointed out by Grew et al. (1992) and Shiraishi & Kagami (1992). Petrochemical studies of mafic to felsic orthogneiss indicate formation in oceanic and island-arc to continental margin environments (Osanai et al. 1992; Ikeda & Shiraishi 1998). There is little indication of high-grade metamorphism during c. 1130 and 1000 Ma magmatism, and, significantly, the data provide no clear evidence of a Grenvillian orogenic event produced by continent–continent collision. Mesoproterozoic juvenile crust has also been reported from the Lurio Belt of northern Mozambique (e.g. Grantham et al. 2008), the Vijayan Complex of Sri Lanka (e.g. Kro¨ner et al. 2003) and other Gondwana fragments. Some workers have discussed the unity of these terranes during Gondwana formation (e.g. Ravikant et al. 2007; Grantham et al. 2008). Although it is far beyond of the scope of this paper, it is essential to compare these terranes with precise age data. It is characteristic that c. 750–800 Ma magmatism (and possible metamorphism) is recorded in detrital zircon from paragneisses in the NE Terrane. These detrital ages demonstrate that sedimentation occurred after c. 750 Ma in the NE Terrane. Because some of

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the sediments have volcaniclastic characteristics, it is also possible that sedimentation on a c. 1000 Ma basement occurred before and at 800–750 Ma. In any case, the formation of high-grade metamorphism occurred after 750 Ma. The abundant and widespread growth of metamorphic zircon at c. 600 Ma is interpreted as timing the peak of granulite-facies metamorphism and ductile deformation in the NE Terrane. Grew et al. (1989) and Asami et al. (2007) suggested two stages of metamorphism; around 760–800 8C and 7–8 kbar for peak conditions of granulite-facies metamorphism, and around 500–600 8C for subsequent amphibolite-facies metamorphism. We suggest that the later metamorphism took place at c. 550–570 Ma, as indicated by minor zircon growth in paragneiss (sample 85020401C) and the emplacement of relatively undeformed granite dykes (sample 90112302A) at c. 550 Ma, which may have derived from the mobilization of anatectic melts originally produced during the earlier granulite-grade metamorphism. It is worth noting that the c. 600 Ma episode of zircon growth is not present in all samples, even those from adjacent localities, and that the timing of granulite-grade metamorphism could be easily missed with a smaller set of samples. A similar phenomenon was also discussed for the southern Indian ultrahigh-temperature granulite-facies rocks (Santosh et al. 2008). Only one published Sm–Nd mineral isochron age of 624 + 18 Ma, from enderbitic gneiss in the NE Terrane, corresponds to the metamorphic zircon age (Shiraishi & Kagami 1992). Although the zircon results broadly support the assertion of Asami et al. (2005) that the granulitegrade metamorphic terrane was produced by a Pan-African event, it is difficult to reconcile the dominance of c. 600 Ma ages from metamorphic zircon with the dominance of c. 540– 515 Ma ages from metamorphic monazite. The discrepancy is present even in single samples of paragneiss: sample 85020401C, with metamorphic zircon ages of 609 + 11 Ma and 565 + 7 Ma, yielded a monazite age of 541 + 15 Ma (five unzoned grains); and sample 8402004, with no metamorphic zircon age identified, yielded monazite ages of 517 + 14 Ma (cores of 11 grains) and 542 + 12 Ma (rims of 11 grains; Asami et al. 2005). In another sample, two zoned monazite grains in garnet porphyroblasts yielded rim ages of 553 + 11 and 544 + 17 Ma and core ages of 753 + 37 and 697 + 47 Ma (Asami et al. 2005). Isotopic resetting of monazite through Pb loss at temperatures below granulite facies is now widely regarded as unlikely (Cherniak et al. 2004), so the younger monazite ages should be attributed to either new metamorphic growth or recrystallization. Because the emplacement of c. 550 Ma granitic dykes places a lower age

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Table 16. SHRIMP and TDM ages from central–eastern Dronning Maud Land Central DML

Schirmacher Hills

Sør Rondane Mountains SW

c. 510 Ma low-temperature hydrothermal processes

Western Rayner Complex

c. 520 –550 Ma Peak metamorphism (UHT)

c. 520– 540 Ma Metamorpshim

515 Ma metamorphism (Ttn), post-orogenic granite c. 530 Ma Magmatism c. 560 Ma Metamorphism, Magmatism

c. 605 Ma late-kinematic grt –mus pegmatite c. 625 Ma c. 650 Ma Granulite-facies Magmatism metamorphism (UHT?)

c. 540 Ma Metamorphism c. 600– 650 Ma Peak granulite facies (UHT?)

c. 620 Ma Magmatism? c. 660 Ma Metamorphism?

c. 800 Ma Metamorphism/ magmatism?

c. 800 Ma Magmatism (det), metamorphism? (det) c. 1100 Magmatism and granulite-facies metamorphism

c. 1000 Ma Metamorphism

c. 1000 Ma Magmatism (prot þ det) c. 1130 Ma Magmatism

c.1000 Ma Magmatism(prot)

c. 1200, 3260 Ma (det) 0.9 –1.0, 1.6 Ga Crustal inheritance (TDM)

Rayner Complex

NE

c. 550 Ma Lamprophyre c. 575 –590 Ma Cooling

Lu¨tzow– Holm Complex

0.9– 1.0, 1.6 Ga Crustal inheritance (TDM)

K. SHIRAISHI ET AL.

c. 530 –515 Ma Granulite-facies metamorphism c. 570 Ma Amphibolite-facies metamorphism c. 600 Ma Magmatism

Yamato– Belgica Complex

c. 2470 (inh) 1.0 –1.8, 2.3 –2.7 Ga Crustal inheritance (TDM)

c. 2900 – 1500 Ma (det)

c. 910 –980 Ma Metamorphism, magmatism c. 1040 –2440 Ma (det) 2.3 Ga Crustal inheritance (TDM)

c. 1810 –2580 (det) 1.6 –2.3 Ga Crustal inheritance (TDM)

Bold numbers are SHRIMP zircon or titanite (Ttn) ages (inh, inherited; prot, protolith; det, detrital). TDM, depleted mantle model age (Ga); UHT, ultrahigh-temperature metamorphism. Ages from central Dronning Maud Land are after Jacobs et al. (1998) and Henjes-Kunst (2004). Those from Schirmacher Hills are based on Sm –Nd data after Ravikant (2006). Other data sources are Shiraishi et al. (1997b, 2003, and this study).


constraint on high-strain deformation and the formation of granulite-grade gneissic fabrics, the ages may be an indication that amphibolite-facies conditions (with monazite and garnet growth) persisted after 550 Ma. Monazite growth and/or recrystallization may also have occurred during contact metamorphism in the vicinity of late to post-tectonic intrusions of granitic and syenitic plutons. This magmatic activity can also be invoked to explain the c. 517 Ma age of titanite in sample 90102801A, through new titanite growth or closure of Pb diffusion at 700 –600 8C (Cherniak 1993). Further monazite dating in a detailed petrographic context is required to resolve these issues. The geochronological results from the SW Terrane are different, and shed some light on the polyphase nature of Pan-African metamorphism in the NE Terrane. The extensive meta-tonalites, produced by subduction-related magmatism (Ikeda & Shiraishi 1998), were emplaced at 960– 920 Ma (Takahashi et al. 1990, and this study). The relationship of meta-tonalites to adjacent lithologies is unclear, as they are bounded by the mylonitic Main Shear Zone, the timing of which is constrained at or later than c. 560 Ma by the intrusion of mylonitized granite at Vengen. Gneissic fabrics in the SW Terrane, associated with amphibolitefacies metamorphic assemblages, are temporally constrained between the ages of magmatic zircon (c. 650 Ma) and metamorphic zircon (c. 570 Ma) in a granitic orthogneiss. There is an absence of c. 600 Ma zircon in the SW Terrane, so that the c. 570 Ma age provides the best estimate of amphibolite-facies metamorphism and ductile deformation. This age coincides with our estimate of the timing of amphibolite-facies metamorphism in the NE Terrane, and unifies the metamorphic history of the terranes at this time. There is little evidence that the SW Terrane experienced high-grade metamorphism before 570 Ma, suggesting that the terranes have different origins. Scant c. 1130 Ma and c. 980 Ma age data from inherited zircon may tentatively indicate common crustal sources for both terranes, but this is far from proven. A unified geological history after 570 Ma is also supported by the presence of post-tectonic granitoids in both terranes, along with c. 517 Ma ages from metamorphic titanite and 500 –420 Ma K –Ar and Rb –Sr ages from various localities across the SRM (e.g. Shiraishi et al. 1997a). Lamprophyric and doleritic dykes, which have been metamorphosed to amphibolite-facies grade, intrude throughout the NE and SW Terranes in the western and central SRM, probably during the waning stages of c. 570 –560 Ma metamorphism. The geochemistry of the mafic dykes indicates a continental within-plate tectonic setting, with source magmas generated from a mixture of

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subduction-related materials and metasomatically enriched mantle (Ikeda & Shiraishi 1995). A similar scenario was suggested for syenite magmatism in the Yamato Mountains (Zhao et al. 1995). Subsequent A-type granitic magmatism is the latest thermal event in the SRM (Li et al. 2003). Similar late to post-tectonic magmatism is also common in central to east Dronning Maud Land (e.g. Bauer et al. 2003; Roland 2004; Jacobs et al. 2008). In summary, supracrustal protoliths of the NE Terrane were deposited on a 1130–1000 Ma juvenile basement, at least partially after 750 Ma, and were metamorphosed under granulite-facies conditions during orogenesis at c. 600 Ma. Subsequent retrograde metamorphism took place at c. 560 Ma, synchronous with the juxtaposition of the NE and SW Terranes by large-scale shear zones that evolved into the Sør Rondane Suture and the mylonitic Main Shear Zone. The 560 Ma event may relate to the extensional collapse of the orogen, whereas the 600 Ma event relates to the peak collisional event. There is no evidence indicating whether the two events are in the same orogenic cycle or not. The 560 Ma event was followed by alkaline magmatism, which induced hightemperature contact metamorphism between 550 and 510 Ma. Cooling and sporadic magmatism continued over a protracted interval, possibly until as late as 420 Ma (Shiraishi et al. 1997a).

Comparing the SRM with neighbouring terranes Lu¨tzow-Holm, Rayner and Yamato – Belgica Complexes. TDM values from the Lu¨tzow-Holm Complex (LHC) suggest that para- and orthogneisses were derived from a variety of Mesoproterozoic and older basement lithologies. The southern Soˆya Coast differs from other parts of the LHC, with indications of Archaean and/or earlyPalaeoproterozoic basement. Areas with similarly ancient crustal residence ages are found in the Rayner and Napier Complexes in East Antarctica, as well as the Highland Complex in Sri Lanka. It is likely that these components relate to an unknown extent of ancient crust, hidden under the ice sheets of East Antarctica. The tectonic significance of Cape Hinode of LHC is controversial. TDM values (1.0–1.1 Ga) of meta-trondhjemite from Cape Hinode are almost contemporaneous with the crystallization age of magmatic zircon (c. 1017 Ma) and Sm –Nd wholerock isochron ages (c. 1030 Ma; Shiraishi et al. 1994, 1995). The meta-trondhjemite is characterized by relatively high Al and Sr, and low Y and Yb contents, similar to Archaean trondhjemites

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and adakites. Trace element modelling suggests that parental magmas were derived by partial melting of a mid-ocean ridge basalt (MORB) source under garnet-stable P–T conditions (Ikeda et al. 1997; Hiroi et al. 2008). Thus both geochemical and geochronological evidence indicate a non-continental source for the c. 1000 Ma magmatism at Cape Hinode. In contrast, a paragneiss sample from Cape Hinode with TDM of 1.8 Ga, comparable with ages from Sandercock Nunataks in the Rayner Complex, indicates the involvement of older crustal sources in the sedimentary provenance. These various lines of evidence from the LHC are consistent with the gradual development of a continental margin throughout, and possibly after, the Mesoproterozoic. Hokada & Motoyoshi (2006) obtained electron microprobe CHIME ages and REE signatures for monazite from metapelitic granulites of Skallen in the LHC. They discovered a two-stage growth of monazite in two out of four samples, at 560– 500 Ma and 650 –580 Ma. They interpreted the older monazite growth as predating the peak metamorphism, on a prograde stage of the P –T path. How this two-stage metamorphism relates to orogenic events in the LHC has not been established, and requires further work. In the Rayner Complex, it has been known that gneisses represent granulite-grade reworking of the Archaean Napier Complex. However, inherited zircon ages and TDM values indicate that there is only minor contribution of Archaean crust to the Rayner Complex, in accordance with Black et al. (1987). The age of crustal protoliths to granulites in the Rayner Complex is still poorly understood. From a combination of SHRIMP U –Pb zircon and Sm– Nd mineral isochron ages, Shiraishi et al. (1997b) revealed a Pan-African aged overprint on rocks from the coastal regions of the western Rayner Complex. In contrast, inland outcrops of the Nye Mountains and Sandercock Nunataks yield zircon ages of 900– 1000 Ma. The Pan-African overprint of the Rayner Complex was also recognized by Motoyoshi et al. (2006), in CHIME monazite ages from metapelitic granulite of Forefinger Point. These monazites yield 750– 1000 Ma core ages with 517 –528 Ma rims. The rim ages correspond to SHRIMP U – Pb zircon ages of 530 –537 Ma (Shiraishi et al. 1997b). Thus the western Rayner Complex is geochronologically heterogeneous, with the coastal region adjacent to the LHC overprinted by a Cambrian orogeny. The Yamato –Belgica Complex (YBC) lies between the SRM and the LHC, and plays a critical role in understanding the relationship between the two regions. The Yamato Mountains are composed of large volumes of c. 500 Ma syenite (Zhao

et al. 1995). It has been suggested that syenite plutons formed within the hinterland of a c. 530–550 Ma continental collision zone, with the generation of syenite parental magma by partial melting in a mantle wedge above a subduction zone (Shiraishi et al. 1994; Zhao et al. 1995). Considering the TDM and lithological types present in the LHC, the subducting plate included Neoproterozoic continental margin sediments associated with a wide range of protolith and provenance ages. Shiraishi et al. (1994, 2003) showed for the first time that East Antarctica was not a united continent before the amalgamation of East and West Gondwana during the Pan-African tectonic event. Central and western Dronning Maud Land. In recent years, multiple stages of thermal events during the Neoproterozoic to Cambrian PanAfrican orogeny have been well documented in central and western Dronning Maud Land (e.g. Jacobs et al. 1998; Paulsson & Austrheim 2003; Bisnath et al. 2006). It has been suggested that the tectonothermal history of the SRM is similar to that of central Dronning Maud Land (DML) (e.g. Jacobs & Thomas 2002). Felsic magmatism took place at c. 1130 Ma in central to western Dronning Maud Land, (as described by Jacobs et al. (1998) and Bisnath et al. (2006), and has been related by these workers to the formation of an extensive volcanic arc. TDM values for meta-igneous rocks from Wohlthatmassiv and Orvinfjella in central DML vary from 1.02 to 1.74 Ga, assuming a two-stage evolution of the Sm–Nd isotope system. This age range is similar to results recalculated by the method of Stern (2002), with a range of 0.9–1.7 Ga that is in good agreement with TDM values for the SRM (0.9–1.6 Ga). Jacobs et al. (1998) concluded that the Grenville-age crust in DML is basically juvenile, and did not involve significant amount of previous crust. This also appears to be the case in the SRM. However, the present study reveals differences between the two regions. Mesoproterozoic (c. 1000 Ma) metamorphism in the SRM is not yet confirmed by zircon growth, although this may be an artefact of insufficient SHRIMP data. Two stages of late Neoproterozoic–Cambrian metamorphism are recognized in both central DML (Jacobs et al. 1998) and the SRM, but the timing is not identical: 570– 550 Ma and 530–515 Ma in Wohlthatmassiv and Orvinfjella, v. c. 650– 600 Ma and c. 570–550 Ma in the SRM. In contrast, Henjes-Kunst (2004) reported granulite-facies metamorphism at c. 625 Ma and cooling below c. 5008C after amphibolite-facies metamorphism at c. 575–590 Ma in the coastal Schirmacher Hills. This corresponds closely to the inferred timing of


similar events in the SRM. Ultrahigh-temperature metamorphism was reported from the Schirmacher Hills (Baba et al. 2006). Baba et al. (2008) have also reported contrasting P –T–t paths between the Schirmacher Hills and the inland mountains of central DML. They suggested that an isobaric cooling path in Schirmacher Hills is comparable with the granulite- to amphibolite-grade retrograde path for the SRM proposed by Asami & Shiraishi (1987) and Asami et al. (2007). Although c. 1000 Ma metamorphism has not been demonstrated in the SRM, the metamorphic conditions, P–T paths and ultrahigh-temperature metamorphism are comparable between the NE Terrane of the SRM and Schirmacher Hills. In contrast, the lack of c. 600 Ma granulite-grade metamorphism in the SW Terrane is comparable with inland mountains in central Dronning Maud Land, where c. 580 –550 Ma granulite-grade metamorphism occurs. However, further petrological studies tied to geochronology are required. In this context, future zircon chronology to constrain the age of sedimentation of supracrustal rocks in Schirmacher Hills is necessary to establish these regional relationships. In the present discussion we have focused on the crustal development and subsequent metamorphism of terranes in eastern Dronning Maud Land. The age signatures of each terrane in eastern DML are clearly distinguishable (Table 16). Basement rocks of the YBC may represent the easternmost part of the SRM terrane, and the tectonic setting of the SRM is comparable with that in central DML. In contrast to the SRM, protoliths of the LHC have a more complicated nature, and may represent a collage of terranes involving Archaean protoliths in the southern part of Lu¨tzow-Holm Bay and Proterozoic volcanic arc and oceanic components along the Prince Olav Coast. The late Neoproterozoic to Cambrian reworking of the western margin of the Mesoproterozoic Rayner Complex and the presence of Mesoproterozoic crustal signatures in the LHC suggests that the two terranes were associated prior to orogenesis during the Pan-African event. This event is polyphase, involving c. 600 Ma orogenesis in Dronning Maud Land and a final collision at c. 550 –530 Ma in the LHC. From the extensive database of geochronological data obtained in this region of East Antarctica, a picture of the complex formation of this critical section of Gondwana is beginning to emerge. We are grateful to R. Armstrong, K. Misawa, H. Kaiden and O. Tachikawa for helping with SHRIMP analysis, and S. Ohno for making thin sections. The depleted mantle model age by DePaolo’s model was calculated with the Excel spreadsheet given by courtesy of R. J. Stern. Discussions with Y. Hiroi, G. H. Grantham,

63

Y. Motoyoshi, S. Baba, Y. Osanai and M. Owada are highly appreciated. We thank M. Santosh, R. Fuck and J. Jacobs for their constructive reviews and for improving the manuscript. This research was financially supported by a Grant-in-Aid for Scientific Research from the Japan Society for the Promotion of Science No. 13440151 to K.S.

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Early Palaeozoic orogenic collapse and voluminous late-tectonic magmatism in Dronning Maud Land and Mozambique: insights into the partially delaminated orogenic root of the East African – Antarctic Orogen? JOACHIM JACOBS1, BERNARD BINGEN2, ROBERT J. THOMAS3, WILFRIED BAUER3, MICHAEL T. D. WINGATE4 & PAULINO FEITIO5 1

Department of Earth Science, University of Bergen, Alle´gaten 41, 5007 Bergen, Norway (e-mail: [email protected]) 2

Geological Survey of Norway, 7491 Trondheim, Norway

3

British Geological Survey, Kingsley Dunham Centre, Keyworth, Nottingham NG12 5GG, UK

4

Geological Survey of Western Australia, 100 Plain Street, East Perth, W.A. 6004, Australia 5

Direcaõ Nacional Geologia, Maputo, Mozambique

Abstract: The late tectonic history of the southern part of the Late Neoproterozoic– Early Palaeozoic East African–Antarctic Orogen (EAAO) is characterized by lateral extrusion, extensional collapse and large volumes of high-temperature A2-type granitoids. This late-tectonic igneous province covers an area more than 15 000 km2 of the EAAO in Dronning Maud Land (East Antarctica) and its northerly continuation as the Nampula Complex of NE Mozambique. The magmatic province is bounded in the north by the Lurio Belt. New secondary ionization mass spectrometry (SIMS) U–Pb analyses of zircons from two major late-tectonic granitoid intrusions from Dronning Maud Land indicate crystallization ages of 501 + 7 and 499 + 4 Ma, whereas a major extensional shear zone was dated at 507 + 9 Ma. New SIMS zircon U–Pb analyses of late-tectonic granitoid sheets and plutons from the Nampula Province indicate ages of 512 + 4, 508 + 4, 508 + 2 and 507 + 3 Ma. Consequently, the late-tectonic magmatism can be bracketed between c. 530 and 485 Ma. It started with small gabbro bodies emplaced at c. 530– 520 Ma, culminated with the intrusion of major granite– charnockite plutons at c. 510–500 Ma and terminated with the introduction of small volumes of sheet-like granite at c. 485 Ma. The new dates demonstrate that extensional shearing and granitoid intrusion are synchronous, and that orogenic collapse and the magmatism are related. We ascribe the distribution, structural style, geochemical composition and age of the late magmatic province to a process of partial delamination of the orogenic root in the southern third of the EAAO. It remains to be tested whether there is a relationship between orogenic collapse–granitoid magmatism and south-directed escape tectonics in the southernmost EAAO.

Northern Mozambique and Dronning Maud Land (East Antarctica) are interpreted to together represent the southern end of the Late Neoproterozoic– Early Palaeozoic East African–Antarctic Orogen (EAAO) (e.g. Stern 1994; Jacobs & Thomas 2004). This orogen stretches for more than 8000 km from Egypt–Arabia in the north, southwards through East Africa (including Madagascar) into northern Mozambique and thence into Dronning Maud Land in East Antarctica (Fig. 1). The EAAO resulted from a multi-plate collision of various parts of East and West Gondwana and shows a strong lateral variation in orogen style, probably as a result of this complex collision. It is characterized by accretion in its northern third (the ‘Arabian–Nubian Shield’) and

by continent–continent collision in its central and southern part. The deep erosion level exposed allows unique insights into an orogen that fundamentally changes in character along strike. The southernmost segment of the orogen, and its eventual termination, is recognized within Dronning Maud Land. This area is characterized by orogenic collapse and lateral extrusion tectonics, similar to the present situation in southeastern Asia, resulting from the collision of India and Asia (Jacobs & Thomas 2004). The southern third of the orogen from Antarctica into northern Mozambique is characterized by the intrusion of large volumes of A2-type granite (Roland 2004a, b), the volume of which decreases dramatically at the Lurio Belt, a conspicuous shear belt in


70 J. JACOBS ET AL. Fig. 1. (a) Geological setting of the East African– Antarctic Orogen (EAAO) after Jacobs & Thomas (2004). (b) The southern part of the EAAO affected by lateral extrusion and widespread intrusion of late-tectonic granitoids. The latter are largely confined to the EAAO and effectively terminate along the Lurio Belt in northern Mozambique. ANS, Arabian– Nubian Shield; C, Coats Land; DML, Dronning Maud Land; EH, Ellsworth– Haag; F, Filchner block; FI, Falkland Islands; G, Grunehogna; Ga, Gariep Belt; H, Heimefrontfjella; K, Kirwanveggen; L, Lurio Belt; Na---Na, Namaqua –Natal; SR, Shackleton Range; Da, Damara Belt; LH, Lu¨tzow-Holm Bay; M, Madagascar; Sa, Saldania Belt; Z, Zambezi Belt.

GRANITES IN EAAO

NE Mozambique, that trends oblique to the overall north–south strike of the EAAO. In this paper we review and highlight the significance of extensional shearing within the EAAO of northern Mozambique and Dronning Maud Land, and provide new age constraints for the extensional shear zones and associated late-tectonic granitoid intrusions in both areas.

Geological setting In Dronning Maud Land the Late Neoproterozoic – Early Palaeozoic collision along the EAAO to a large extent overprints Mesoproterozoic basement (Figs 1 and 2). It has been shown that this older crust is predominantly juvenile and was generated in island arcs along the margin of the ProtoKalahari Craton at c. 1.1 Ga (Jacobs et al. 1998, 2008; Bauer et al. 2003). The rocks underwent a first high-grade metamorphic event, associated with abundant syntectonic granitoids, between c. 1090 and 1070 Ma (Jacobs et al. 1998). This regional metamorphism was related to a collision event, which led to this area being incorporated into the supercontinent of Rodinia. After the

71

Mesoproterozoic orogenesis, there is little evidence for tectonic activity between c. 1050 and 650 Ma, with the exception of the Schirmacher Oasis area, where there is limited evidence for granitoid intrusion at c. 760 Ma (Jacobs et al. unpubl. data). The Late Neoproterozoic –Early Palaeozoic collision history can be separated into three major phases, as follows. (1) An earliest granulite facies stage is recorded in the Schirmacher Nappe at c. 620 Ma (HenjesKunst 2004), followed by anorthosite magmatism in the main mountain range at c. 600 Ma (Jacobs et al. 1998). This event is associated with shallowly inclined structures that are probably related to nappe emplacement, the age of which is unknown, but that might be related to the granulite- facies metamorphism. (2) The main deformation and medium- to highgrade metamorphism in the main mountain range is bracketed in age by metamorphic zircon rims between c. 590 and 550 Ma, and is interpreted to represent the collision phase (Jacobs et al. 1998, 2003b). It produced tight to isoclinal, upright east –west- to ESE–WNW-trending folds, which are post-dated by a major sinistral shear zone along the southern margin of the mountain range

Fig. 2. Overview map of western and central Dronning Maud Land, Antarctica, depicting main structural trends of the EAAO, extent of late-tectonic intrusions and sample localities. Western orogenic front of the EAAO is a major shear zone. A, Annandagstoppane; AH, Alexander von Humboldt Gebirge; B, Borgmassivet; F, Filchnerfjella; G, Gjelsvikfjella; H, Heimefrontfjella; MG, Mu¨hlig-Hofmann Gebirge; N, Novolazarevskaya Station; O, Orvinfjella; OG, Otto von Gruber Gebirge; S, Schirmacher Nappe; W, Wohlthatmassiv.

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J. JACOBS ET AL.

in Orvinfjella and transpressive structures in the Wohlthatmassiv (Bauer et al. 2004). These structures are cut by discrete extensional shear zones and are intruded by undeformed pegmatites and granite veins, constraining a Cambrian age for this tectonic event (Jacobs et al. 2003b). (3) A late-tectonic stage is associated with extension, tectonic exhumation and south-directed crustal extrusion between c. 530 and 485 Ma, exposing mid- to lower crustal levels (e.g. Jacobs et al. 2003a; Engvik & Elvevold 2004; Jacobs & Thomas 2004). This period is accompanied by syntectonic extensional shearing, late- to post-tectonic intrusions and isothermal decompression (e.g. Jacobs et al. 2003b; Colombo & Talarico 2004). The volume of igneous rocks generated increases towards the end of the extensional period, culminating in voluminous and extensive granitoid magmatism, which is in part charnockitic. Thermobarometry studies show that the charnockites were emplaced as relatively dry melts at temperatures exceeding 900 8C at pressures of c. 5 kbar (Frost & Bucher 1993; Bucher & Frost 1995). Many of the charnockites are retrogressed in part to granite, especially at the contacts and adjacent to late hydrous granite pockets. The geochemistry of the charnockites and associated granitoids is relatively heterogeneous, but they are typically peraluminous to metaluminous and subalkaline with a weak trend to alkaline A-type granites (e.g. Klimov et al. 1964; Ravich & Kamenev 1975; Joshi et al. 1991; Roland 2002, 2004a, b; Li et al. 2003). However, they are not typical A-type granites, being relatively low in Ca, Rb, Nb and Ga, such that they plot as A2-type according to the classificaton of Eby (1992), and unlike common A-type granite associations they form a voluminous and extensive magmatic suite, covering an area of at least 15 000 km2 (Roland 2004a, b). The available geochronology of this late-tectonic stage is summarized in Table 1. As in Dronning Maud Land, the EAAO of northern Mozambique reworks rocks with predominantly Mesoproterozoic protolith ages. However here, the EAAO is divided into two different crustal segments by the ENE –WSW trending Lurio Belt (Figs 1 and 3). The crust south of the Lurio Belt has close similarities to the EAAO of Dronning Maud Land, whereas the EAAO to the north of the Lurio Belt, composed of a collage of terranes, is structurally and lithologically different. The crust south of the Lurio Belt, referred to as the Nampula Complex, is made up of Mesoproterozoic gneisses and migmatites of upper amphibolite facies (e.g. Pinna et al. 1993). This Mesoproterozoic ‘basement complex’ is overlain by: (1) a sequence of Neoproterozoic synorogenic immature clastic sediments, the Mecuburi and Alto Benfica

Groups (e.g. Thomas et al. 2006); (2) tectonic slices (thrust sheets?) of granulite-facies rocks, the Mocuba and Monapo klippen (e.g. Pinna et al. 1993), which might be similar to comparable structures such as the Schirmacher nappe in Dronning Maud Land. The timing of Neoproterozoic – Cambrian amphibolite-facies metamorphism in the Nampula Complex and overlying sedimentary rocks is estimated at 520–490 Ma (e.g. Bingen et al. 2006a, b), whereas the timing of granulitefacies metamorphism in the structurally overlying klippen is dated at 615 + 8 Ma (Kro¨ner et al. 1997). Available data suggest that the Mocuba and Monapo klippen correlate with the Neoproterozoic upper nappes observed north of the Lurio Belt (Pinna et al. 1993). The nature of the observed juxtaposition of the Mesoproterozoic Nampula Complex basement and the granulite klippen is a matter of debate, but previous studies interpreted the contact as a thrust (e.g. Pinna et al. 1993; Kro¨ner et al. 1997). As in Dronning Maud Land, the Nampula Complex is characterized by the presence of large volumes of post-Mesoproterozoic late-tectonic granitoid, with a concentration of such intrusions along the Lurio Belt. Very limited geochronological or geochemical data are available from these rocks. The late-tectonic granite magmatism essentially terminates along the Lurio Belt, although a diminishing number of scattered plutons are found to the north. The Lurio Belt is a northerly inclined high-strain zone, interpreted as a repeatedly reactivated shear zone with an apparent intense late-tectonic pure shear deformation history (Viola et al. 2006). It is marked by isoclinal folding, strongly attenuated lenses of granulites and a characteristic suite of highly sheared leucogneisses. The Lurio Belt shows an intense late-tectonic deformation history as young as c. 500 Ma (Bingen et al. 2006a, b). The belt is an extremely well-defined and coherent structure in the east, prominent in the field and on remote-sensing images, which becomes progressively diffuse and ill-defined towards the west (e.g. Thomas et al. 2006). The EAAO crust north of the Lurio Belt is made up of a collage of generally north–south- to NE – SW-trending nappes. The lower nappes, referred to as the Unango and Marrupa Complexes, are made up of gneisses with Mesoproterozoic protolith ages. These were reworked by high-grade metamorphism dated between c. 560 and 520 Ma (Bingen et al. 2006a, b; Engvik et al. 2007; Norconsult Consortium 2007). The upper nappes, referred to as the Xixano, M’Sawize and Lalamo Complexes, are dominated by rocks with Neoproterozoic protolith ages, and carry evidence for Late Neoproterozoic high-pressure metamorphism dated between 740 and 600 Ma. The upper nappes can

Table 1. Published U–Pb zircon and titanite dates for late- to post-tectonic rocks of Dronning Maud Land, East Antarctica Date (Ma)

Method

Interpretation

Reference

Highly sheared gneiss, Conrad Mts., extensional? Zwiesel gabbro Zwiesel gabbro Late tectonic lamprophyre, Risemedet Charnockitized orthogneiss, Hochlinfjellet Mesoproterozoic metavolcanic rock, Dallmannberge Leucocratic segregation in boudin neck, Conradgebirge Post-tectonic syenite, Humboldtgebirge Hornblende leucosome, Festninga Late-tectonic lamprophyre, Risemedet Migmatite, Jutulsessen Granite, Stabben Aplite dyke in granite, Stabben Post-tectonic granite sheet, Gygra Gabbro, Stabben

530 + 8 527 + 6 521 + 6 523 + 5 521 + 3 522 + 10 516 + 5 512 + 2 510 + 14 508 + 7 504 + 6 500 + 8 495 + 14 487 + 4 483 + 14

U –Pb SHRIMP U –Pb SHRIMP U –Pb SHRIMP U –Pb SHRIMP U –Pb SHRIMP U –Pb SHRIMP U –Pb SHRIMP U –Pb TIMS U –Pb SHRIMP U –Pb SHRIMP U –Pb SIMS U –Pb SIMS Evaporation U –Pb SHRIMP U –Pb SHRIMP*

Metamorphic zircon rim Crystallization Crystallization Crystallization Late charnockitization Metamorphic zircon rim Crystallization of leucosome Crystallization Youngest metamorphic zircon rim Hydrothermal or metasomatic overprint Migmatisation Crystallization Crystallization Crystallization Cooling through 600 8C

Jacobs et al. (1998) Jacobs et al. (2003b) Jacobs et al. (2003b) Jacobs et al. (2003a) Jacobs et al. (2003a) Jacobs et al. (1998) Jacobs et al. (1998) Mikhalsky et al. (1997) Jacobs et al. (2003a) Jacobs et al. (2003a) Paulsson & Austrheim (2003) Paulsson & Austrheim (2003) Paulsson & Austrheim (2003) Jacobs et al. (2003a) Jacobs et al. (2003a)

GRANITES IN EAAO

Lithology and locality

*

Titanite.

73

74

J. JACOBS ET AL.

Fig. 3. Geological overview map of northern Mozambique with sample localities.

possibly be correlated with similar structures in Tanzania and southern Madagascar. The nappe pile north of the Lurio Belt shows a complex and prolonged tectonometamorphic history as a result of an overall NW-directed collisional history along the western margin of the Congo–Tanzania craton between c. 590 and 520 Ma (e.g. Pinna et al. 1993; Jacobs & Thomas 2002; Bingen et al. 2006a, b; Viola et al.

2006). The EAAO north of the Lurio Belt shows few late-tectonic extensional structures or latetectonic igneous rocks compared with the Nampula Complex and Dronning Maud Land. Both in northern Mozambique and Dronning Maud Land, the late-tectonic intrusions have impressive morphological expressions (Fig. 4). In Dronning Maud Land many of the plutons form

GRANITES IN EAAO

Fig. 4. The late-tectonic granitoids form impressive mountains, both in Dronning Maud Land (a) and northern Mozambique (b).

high, steep-sided nunataks, which are not dissimilar in form from the inselberg equivalents in Mozambique.

Analytical details, samples and geochronological results Seven samples were selected for U – Pb secondary ionization mass spectrometry (SIMS) zircon dating, three from Dronning Maud Land and four from NE Mozambique, to constrain the Early Palaeozoic extensional shearing and the late magmatic history of the two areas. From Dronning Maud Land, one sample (J3012/1), containing complex zircons, were collected from a high-strain zone at Armlenet. Two samples (J1670, J1870) from the Mu¨hligHofmann-Gebirge are from different late- to post-tectonic granitoid intrusions at Schneide and Oddesteinen (J1670, J1870) (Fig. 2). In northeastern Mozambique, a sample (JJ238) was selected from a highly foliated granite gneiss sheet, interpreted as syntectonic, another

75

sample (WB295) was selected from a mangerite intrusion within the Lurio Belt (WB295), and two samples (GV01, JJ259) were obtained from two late-tectonic porphyritic granite intrusions with largely undeformed centres and weakly deformed margins (Fig. 3). Zircons were recovered from crushed samples using conventional separation techniques (Wilfley table, magnetic separation, high-density liquids, and hand picking). Zircons were cast, together with a zircon reference standard, in epoxy discs, which were then polished to expose the interiors of the crystals. All zircons were examined and documented in transmitted and reflected light, and cathodoluminescence (CL) images were used to reveal the internal structures of the zircons, such as core –rim relationships in the metamorphic sample (J3012/1). After cleaning, the sample mounts were vacuum-coated with c. 500 nm Au. Zircon U – Pb analysis of five samples were conducted by sensitive high-resolution ion microprobe (SHRIMP) using the SHRIMP II ion microprobe at the John de Laeter Centre of Mass Spectrometry, Curtin University of Technology, Perth, Australia. Two samples were analysed on a Cameca IMS 1270 ion microprobe at the NORDSIM facility, Swedish Museum of Natural History, Stockholm. SHRIMP analytical methods follow those of Williams (1998) and references therein. Analyses consist of six scans through the mass range using a spot size of c. 20 mm diameter. Absolute 238U and 232Th concentrations were determined by comparison with the CZ3 zircon standard (551 ppm 238U) and 238U/206Pb ratios were determined relative to the Temora zircon (417 Ma); standard (206Pb/238U ¼ 0.0668 Black et al. 2003), analyses of which were interspersed with those of unknown zircons. Analyses with the Cameca IMS1270 followed the methods outlined by Whitehouse et al. (1999) and Whitehouse & Kamber (2005), with a spot size of c. 15 mm. The Cameca analyses were calibrated to the Geostandard 91500 reference zircon with an age of 1065 Ma (Wiedenbeck et al. 1995). Uncertainties on U/Pb ratios include propagation of the uncertainty of the day-to-day calibration curve determined from repeated analysis of the reference zircon. A common Pb correction was performed using non-radiogenic 204Pb with an average modern crustal Pb composition (Stacey & Kramers 1975). Data were analysed using software programs SQUID and ISOPLOT (Ludwig 2001, 2003). Uncertainties in ratios and ages in Table 2 are listed at the 1s level. Weighted mean ages are quoted below with 95% uncertainties.

76

Table 2. Ion microprobe data on zircon from Dronning Maud Land and Mozambique Identifier 1

U (ppm) 2

3

Th (ppm)

Pb (ppm)

f (%) 4

238

U/ Pb 5

206

1s (%)

207

Pb/ Pb 5

206

1s (%)

238

U/ Pb 6

206

1s (%)

207

Pb/ Pb 6

1s (%)

206

206 238

Pb/ U (Ma) 7

1s

207

Pb/ Pb (Ma) 8

206

1s

Disc (%) 9

2.2 2.3 2.1 2.2 2.2 2.1 2.4 2.2 2.4 2.1 5.8

0.0575 0.0516 0.0568 0.0574 0.0563 0.0570 0.0757 0.0751 0.0770 0.0761 0.0769

1.7 4.2 0.8 2.1 1.3 0.7 0.9 0.6 1.6 0.8 2.4

498 507 508 514 518 522 972 1087 1128 1147 1166

10 11 10 11 11 11 22 23 26 24 65

513 268 483 507 464 490 1086 1070 1121 1098 1146

38 96 18 45 28 16 17 12 32 17 47

3 283 25 21 211 26 10 21 21 24 22

J1670, post-tectonic granite, Dronning Maud Land, Oddesteinen, SHRIMP Perth 4.1 106 83 0.05 12.82 1.3 0.0577 2.0 5.2 59 40 0.38 12.67 1.4 0.0613 2.6 6.1 90 102 0.11 12.61 2.3 0.0574 2.2 12.1 236 179 0.17 12.53 1.3 0.0571 1.8 10.1 212 67 0.15 12.52 1.2 0.0580 1.4 5.1 94 51 0.00 12.46 1.6 0.0597 2.1 11.1 183 140 0.00 12.42 1.2 0.0557 1.6 3.1 117 76 0.29 12.32 1.3 0.0575 1.9 1.1 287 146 0.00 12.33 1.2 0.0570 1.3 9.1 259 146 0.18 12.30 1.3 0.0571 1.3 2.1 196 109 0.07 12.30 1.5 0.0577 1.5 7.1 173 71 0.00 12.29 1.4 0.0593 1.6 9.2 336 130 0.15 12.14 1.8 0.0572 1.6 8.1 157 53 0.16 12.09 1.6 0.0576 1.9

12.83 12.72 12.62 12.55 12.53 12.46 12.42 12.36 12.33 12.32 12.31 12.29 12.16 12.11

1.3 1.5 2.3 1.3 1.2 1.6 1.2 1.3 1.2 1.3 1.5 1.4 1.8 1.6

0.0573 0.0583 0.0566 0.0557 0.0567 0.0597 0.0557 0.0552 0.0570 0.0556 0.0571 0.0593 0.0560 0.0563

3.7 5.4 4.3 2.7 2.4 2.1 1.6 3.7 1.3 1.7 1.8 1.6 2.6 2.7

484 487 492 495 495 496 500 503 503 504 504 503 510 512

6 7 11 6 6 8 6 6 6 6 7 7 9 8

503 539 474 440 480 657 486 419 515 438 496 604 453 463

82 119 96 61 53 55 40 83 32 38 40 37 57 60

4 10 23 212 23 24 23 219 2 215 21 16 212 210

J1870, post-tectonic granite, Dronning Maud Land, Schneide, 1.1 69 42 0.36 12.56 2.1 170 105 0.00 12.53 4.1 166 119 0.00 12.50 10.1 99 65 0.81 12.35 9.1 76 52 0.66 12.28 7.1 205 118 0.16 12.28

12.60 12.53 12.50 12.45 12.36 12.30

2.2 2.0 2.1 2.1 2.2 1.9

0.0545 0.0567 0.0586 0.0525 0.0536 0.0572

7.8 1.9 1.9 4.8 6.5 1.9

494 495 495 501 504 504

11 10 10 10 11 10

390 612 580 307 356 500

176 76 44 109 146 43

226 19 14 262 241 21

SHRIMP Perth 2.2 0.0574 2.0 0.0567 2.1 0.0586 2.1 0.0591 2.1 0.0590 1.9 0.0586

2.9 1.9 1.9 2.3 2.7 1.8

J. JACOBS ET AL.

J3012, sheared felsic gneiss, Dronning Maud Land, Armlenet, Gjelsvikfjella, SHRIMP Perth 9.1 748 64 0.15 12.43 2.2 0.0587 1.3 12.45 10.1 x 607 34 0.91 12.19 2.2 0.0590 1.6 12.30 11.1 2107 9 0.03 12.20 2.1 0.0570 0.7 12.20 6.1 331 44 0.07 12.04 2.2 0.0579 1.6 12.05 4.1 1364 1 0.09 11.96 2.2 0.0571 0.9 11.97 5.1 1625 3 0.01 11.88 2.1 0.0570 0.7 11.88 8.1 x 700 37 0.13 6.11 2.4 0.0767 0.7 6.11 3.1 1954 48 0.09 5.44 2.2 0.0758 0.6 5.45 7.1 570 148 0.02 5.23 2.4 0.0772 1.6 5.23 1.1 452 145 0.07 5.14 2.1 0.0767 0.8 5.15 2.1 269 84 0.00 5.05 5.8 0.0769 2.4 5.05

173 244 440 69

2.3 1.9 1.9 2.2

0.0556 0.0549 0.0549 0.0577

2.9 2.3 1.7 2.9

506 508 509 513

11 10 9 11

438 410 408 541

64 51 38 65

215 223 224 5

JJ238, granite gneiss, 11.1 x 10.1 x 1.1 14.1 12.1 16.1 9.1 13.1 4.1 2.1 3.1 7.1 8.1 15.1 6.1 5.1

Mozambique, Nampula Complex, UTM: Zone37, 266082, 8299521, SHRIMP 189 240 0.12 12.97 1.1 0.0588 1.2 12.98 1.1 277 75 0.08 12.75 0.9 0.0574 0.9 12.76 0.9 144 179 0.00 12.43 1.1 0.0571 1.4 12.43 1.1 199 230 0.17 12.40 1.0 0.0577 1.2 12.42 1.0 319 104 0.00 12.29 0.9 0.0563 0.9 12.29 0.9 274 214 0.10 12.27 1.0 0.0569 1.0 12.29 1.0 275 115 0.01 12.27 1.0 0.0564 1.0 12.27 1.0 412 1293 0.00 12.25 0.9 0.0574 0.8 12.25 0.9 270 172 0.00 12.21 1.0 0.0578 1.0 12.21 1.0 298 99 0.19 12.16 1.0 0.0575 1.0 12.18 1.0 311 97 0.16 12.14 1.0 0.0577 0.9 12.16 1.0 230 84 0.12 12.14 1.0 0.0566 1.1 12.15 1.0 263 736 0.10 12.10 1.0 0.0572 1.0 12.11 1.0 243 379 0.20 12.08 1.0 0.0594 1.1 12.10 1.0 269 182 0.23 12.07 1.0 0.0577 1.0 12.09 1.0 262 674 0.00 11.98 1.0 0.0579 1.1 11.98 1.0

Perth 0.0579 0.0568 0.0571 0.0563 0.0563 0.0561 0.0563 0.0574 0.0578 0.0559 0.0564 0.0556 0.0564 0.0578 0.0558 0.0579

1.8 1.3 1.4 1.8 0.9 1.7 1.1 0.8 1.0 1.6 1.8 1.4 1.6 1.9 1.8 1.1

478 486 499 500 505 505 506 506 507 510 510 511 512 512 513 517

5 4 5 5 5 5 5 5 5 5 5 5 5 5 5 5

525 484 551 464 499 456 464 514 524 450 470 435 469 522 446 546

39 29 40 39 29 37 25 18 22 35 39 32 34 43 41 25

10 21 10 27 21 210 28 2 3 211 28 215 28 2 213 6

498 494 503 504 505 507 512 513 515 516 517 517 518 519 563

8 8 8 8 8 8 8 8 8 8 8 8 8 8 9

201 590 474 491 481 467 488 485 512 491 452 494 477 489 587

133 56 15 39 24 19 16 24 11 15 29 14 20 18 28

2151 17 26 23 25 29 25 26 21 25 215 25 29 26 4

105 192 186 36

0.17 0.33 0.22 0.00

12.25 12.20 12.18 12.08

2.3 1.9 1.9 2.2

0.0570 0.0576 0.0567 0.0577

1.8 1.6 1.1 2.9

12.27 12.24 12.21 12.08

JJ259, Phenocryst granite, Mozambique, Nampula Complex, UTM: Zone 37, 270036, 8290067, CAMECA Stockholm 01b c x 115 257 15 1.00 12.44 1.5 0.0580 3.0 12.57 1.5 0.0501 6.0 16b c x 68 90 8 0.69 12.50 1.6 0.0596 2.6 12.50 1.6 0.0596 2.6 17a r 1187 206 106 0.08 12.31 1.5 0.0571 0.7 12.32 1.5 0.0565 0.7 01a r 1553 885 153 1.07 12.16 1.5 0.0653 0.9 12.30 1.5 0.0570 1.8 02a c 1391 320 126 0.10 12.26 1.5 0.0575 1.0 12.28 1.5 0.0567 1.1 13a r 808 197 74 0.06 12.23 1.5 0.0568 0.8 12.24 1.5 0.0564 0.8 17b r 1311 206 119 0.12 12.12 1.5 0.0569 0.7 12.12 1.5 0.0569 0.7 05a c 1024 147 92 0.13 12.07 1.5 0.0578 1.0 12.08 1.5 0.0568 1.1 10b r 1174 95 106 0.02 12.02 1.6 0.0577 0.5 12.02 1.6 0.0575 0.5 10a c 725 57 65 0.07 12.00 1.6 0.0576 0.6 12.01 1.6 0.0570 0.7 06a r 1036 86 93 0.32 11.97 1.6 0.0585 1.0 12.01 1.6 0.0560 1.3 16a r 979 75 88 0.08 11.98 1.5 0.0577 0.6 11.99 1.5 0.0571 0.7 08a c 449 539 53 0.07 11.97 1.6 0.0571 0.9 11.97 1.6 0.0566 0.9 04a r 1716 261 157 0.15 11.95 1.6 0.0569 0.8 11.95 1.6 0.0569 0.8 13b c x 186 50 19 0.15 10.94 1.6 0.0596 1.3 10.94 1.6 0.0596 1.3

GRANITES IN EAAO

5.1 3.1 6.1 8.1

(Continued) 77

78

Table 2. Continued Identifier 1

U (ppm) 2

3

Th (ppm)

Pb (ppm)

f (%) 4

238

U/ Pb 5

206

1s (%)

207

Pb/ Pb 5

206

1s (%)

238

U/ Pb 6

206

1s (%)

207

Pb/ Pb 6

206

1s (%)

Pb/ U (Ma) 7

500 500 501 504 505 505 506 507 507 508 508 508 508 508 508 510 512 511 513 514 514 515 515 517 525 531 625

1s

5 5 6 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 5 6

207 206

Pb/ Pb (Ma) 8

506 514 546 501 488 500 513 488 512 503 512 488 510 511 517 508 494 528 491 497 520 499 516 526 281 486 704

1s

Disc (%) 9

17 11 30 18 21 15 15 13 13 11 19 11 15 16 13 15 10 23 12 11 18 15 13 19 20 12 11

1 3 8 21 23 21 1 24 1 21 1 24 0 0 2 21 24 3 25 23 1 23 0 2 289 29 11

J. JACOBS ET AL.

GV01, Phenocryst granite, Mozambique, Nampula Complex, UTM: Zone 37, 426538, 8351026, CAMECA Stockholm 48 c 600 440 62 0.08 12.40 1.1 0.0580 0.8 12.41 1.1 0.0574 0.8 47 c 1018 315 96 0.10 12.38 1.1 0.0583 0.5 12.39 1.1 0.0576 0.5 61 c 420 286 46 1.00 12.22 1.2 0.0662 0.8 12.35 1.3 0.0584 1.4 64 r 722 125 65 0.15 12.28 0.9 0.0584 0.7 12.30 0.9 0.0572 0.8 49 c 492 203 48 0.04 12.28 1.0 0.0572 0.9 12.28 1.0 0.0569 0.9 61 r 694 159 64 0.05 12.27 1.0 0.0576 0.7 12.28 1.0 0.0572 0.7 50 c 487 834 63 0.07 12.22 1.0 0.0581 0.6 12.23 1.0 0.0576 0.7 59 r 772 179 71 0.05 12.23 0.9 0.0573 0.6 12.23 0.9 0.0569 0.6 62 r 670 124 61 0.06 12.21 1.0 0.0580 0.6 12.22 1.0 0.0575 0.6 52 r 1018 84 90 0.06 12.19 1.1 0.0578 0.5 12.20 1.1 0.0573 0.5 51 c 347 117 33 0.16 12.18 1.0 0.0587 0.8 12.20 1.0 0.0575 0.9 53 r 944 90 84 0.02 12.20 1.0 0.0571 0.5 12.20 1.0 0.0569 0.5 57 c 530 196 51 0.05 12.19 0.9 0.0578 0.7 12.20 0.9 0.0575 0.7 56 c 440 463 50 0.05 12.18 0.9 0.0579 0.7 12.19 0.9 0.0575 0.7 45 c 578 336 59 0.02 12.18 1.0 0.0578 0.6 12.18 1.0 0.0577 0.6 52 c 535 269 54 0.10 12.13 1.0 0.0582 0.6 12.14 1.0 0.0574 0.7 57.1 c 1274 432 121 0.05 12.10 0.9 0.0575 0.4 12.11 0.9 0.0571 0.4 54 c 528 875 68 0.14 12.09 0.9 0.0591 1.0 12.10 0.9 0.0580 1.0 60 r 826 164 77 0.04 12.09 1.0 0.0573 0.5 12.09 1.0 0.0570 0.6 45 r 769 76 69 0.02 12.06 1.0 0.0573 0.5 12.06 1.0 0.0571 0.5 57.2 c 619 327 62 0.04 12.04 0.9 0.0581 0.8 12.05 0.9 0.0577 0.8 46 c 475 167 46 0.03 12.02 1.0 0.0574 0.7 12.03 1.0 0.0572 0.7 63 r 753 98 69 0.05 12.02 0.9 0.0581 0.5 12.02 0.9 0.0576 0.6 63 c 380 289 41 0.18 11.96 1.1 0.0593 0.8 11.98 1.1 0.0579 0.9 58 r x 959 171 89 0.71 11.79 0.9 0.0575 0.5 11.88 0.9 0.0519 0.9 65 c x 933 310 93 0.08 11.66 1.0 0.0575 0.5 11.67 1.0 0.0569 0.5 55 c x 1544 887 192 0.06 9.78 0.9 0.0633 0.5 9.79 0.9 0.0629 0.5

206 238

422 476 501 505 505 512 507 509 503 511 514 519 510 518 514 521 523 528

5 16 7 8 7 10 5 9 8 7 6 9 8 6 8 8 8 13

270 972 399 585 616 178 544 518 906 452 587 307 884 464 833 688 775 490

87 238 200 113 120 307 56 168 237 101 62 286 138 52 170 114 122 208

236 100 220 16 22 265 7 2 78 211 14 241 71 210 60 31 47 27

1, analysis identifier. 2, r, rim; c, core. 3, x, not selected for age calculation. 4, per cent of total 206Pb made by common Pb. 5, ratios uncorrected for common Pb. 6, ratios corrected for common Pb using 204Pb method, if correction is positive. 7, age corrected for common Pb with the 207Pb method. 8, age corrected for common Pb with the 204Pb method. 9, discordance of the analysis; positive number for an analysis above the concordia curve in a Tera– Wasserburg diagram.

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WB295, Quartz mangerite, Mozambique, Unango Complex, UTM: Zone 37, 219462, 8288506, SHRIMP Perth 8.1 x 79 26 0.20 14.83 1.2 0.0533 2.0 14.86 1.2 0.0516 3.8 6.1 x 26 22 0.00 13.02 3.5 0.0580 4.0 13.02 3.5 0.0580 4.0 4.1 45 39 0.26 12.38 1.5 0.0568 2.4 12.41 1.6 0.0547 8.9 11.1 38 34 0.17 12.23 1.5 0.0608 2.5 12.25 1.6 0.0595 5.2 1.1 28 16 0.00 12.26 1.4 0.0585 3.0 12.26 1.4 0.0585 3.0 8.1 20 18 1.14 12.09 1.9 0.0590 3.4 12.22 2.0 0.0496 13.2 13.1 101 85 0.14 12.19 1.1 0.0595 1.4 12.20 1.1 0.0584 2.6 9.1 24 23 0.03 12.16 1.8 0.0579 3.1 12.17 1.8 0.0577 7.7 17.1 26 25 0.00 12.30 1.7 0.0590 2.9 12.30 1.7 0.0590 2.9 18.1 37 17 0.28 12.11 1.5 0.0583 2.4 12.14 1.5 0.0560 4.5 14.1 56 46 0.00 12.05 1.3 0.0580 2.0 12.05 1.3 0.0580 2.0 15.1 24 10 0.64 11.92 1.7 0.0578 3.1 12.00 1.8 0.0525 12.5 7.1 28 13 0.00 12.08 1.7 0.0623 2.9 12.08 1.7 0.0623 2.9 12.1 90 67 0.17 11.96 1.1 0.0577 1.5 11.98 1.1 0.0563 2.4 3.1 33 21 0.00 12.04 1.6 0.0583 2.7 12.04 1.6 0.0583 2.7 10.1 35 20 0.00 11.89 1.5 0.0574 2.6 11.89 1.5 0.0574 2.6 16.1 30 25 0.00 11.83 1.6 0.0579 2.7 11.83 1.6 0.0579 2.7 2.1 31 30 0.46 11.67 2.5 0.0608 2.7 11.72 2.6 0.0570 9.4

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Results Sample J3012/1, mylonitic felsic gneiss, Armlenet, Dronning Maud Land Sample J3012/1 is a mylonitic felsic gneiss from Armlenet and was collected from the central part of a high-strain zone (Fig. 5), several tens of metres wide, which affected mainly granitic gneisses and augen gneisses as well as highly dismembered amphibolites. The shear zone is cut by pegmatite veins (Fig. 5d) that are in places also sheared. The central part of the shear zone shows high shear strain with no mesoscopic shear-sense indicators, although it contains oblique folds with fold axes dipping towards the NNE. Closer to the shear zone margins at lower shear strains, abundant shear-sense indicators signify extension, with NNEplunging stretching lineations. The analysed sample is equigranular and consists of microcline, partly antiperthitic plagioclase, quartz, and dark brown biotite. Both straight and bulged grain boundaries of quartz are common. This sample contains small (50 –150 mm), clear and colourless to dark brown, equant to elongate zircons, with aspect ratios up to 5:1. Many zircons

are metamict and various inclusions are common. Many show zircon overgrowth and/or resorption features, and several show necking. In CL, the complex internal structure of the zircons becomes apparent (Fig. 6), with some grains showing oscillatory concentric growth zoning in their cores. These cores have a high-U reaction zone of varying width. It is often irregular and appears to be a diffusion front, rather than a simple overgrowth. Eleven areas in five cores and six rims were analysed (Fig. 7a and b). The zircon cores contain 269– 2154 ppm 238U and have Th/U of 0.03– 0.33, whereas rims have 331– 2107 ppm 238U and Th/U from almost zero (0.0006) to 0.14 (Table 2). The proportion (f204) of common 206Pb to total 206Pb is low (,1% for all analyses). The five core analyses form a concordant to slightly discordant group (Fig. 7a) with a mean 207Pb/206Pb age of 1086 + 25 Ma (MSWD ¼ 1.2). Four of them yield a concordia age (Ludwig 1998) of 1098 + 25 Ma (MSWD ¼ 1.9). Five of the six rim analyses plot close to concordia (Fig. 7b). The rim that produced reversely discordant analysis 10.1 has relatively low U and the largest common Pb correction, and its 207Pb/206Pb ratio is probably over-corrected for common Pb. However, this has a

Fig. 5. High-strain extensional shear zone at Armlenet. (a, b) Highly sheared grey migmatic gneisses and sample locality J3012/2; (c) complex folded part of the shear zone; (d) mylonitic gneisses intruded by felsic leucosomes.

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Fig. 6. CL image of selected zircons from mylonite sample J3012/2. Many of the grains are highly complex and show strong recrystallization, probably as the result of high-temperature shearing.

negligible effect on its 238U/206Pb ratio and the weighted mean 238U/206Pb age for all six rim analysis is 507 + 9 Ma (MSWD ¼ 0.99). The age of 1098 + 25 Ma for zircon cores is interpreted as the age of crystallization of the protolith of the gneiss, whereas the age of 507 + 9 Ma for six rim analyses is interpreted to reflect the time of metamorphism during shearing along the Armlenet Shear Zone.

Sample J1670, granite, Oddesteinen, Dronning Maud Land Sample J1670 is a coarse-grained granitoid from Oddesteinen. It is composed of approximately equal amounts of plagioclase, perthitic K-feldspar, quartz, and olive –green hornblende. Accessory minerals include xenomorphic titanite, biotite, apatite and opaque minerals. The sample contains clear and colourless to pale yellow zircons, which commonly occur as inclusions in hornblende. Most are fragments of elongate crystals up to 300 mm in length, with aspect ratios up to 5:1. Some crystals have rounded terminations. Some contain large vermicular (worm-like) melt inclusions, and evidence of resorption is also common. CL images reveal oscillatory growth zoning, but also irregular zoning and recrystallization (Fig. 7c). Some zircon cores are high in U and metamict. Fourteen areas in 13 zircons were analysed, mostly targeting material showing magmatic oscillatory growth zoning (Fig. 7c). The analysed areas

have low common Pb ( f204 ,0.29), and low to moderate 238U concentrations of 59–336 ppm and Th/U of 0.33–1.17, values typical for magmatic zircon. The data form a concordant to very slightly discordant group, and all 14 analyses yield a concordia age of 499 + 4 Ma (MSWD ¼ 1.2). This result is interpreted as the crystallization age of the Oddesteinen granite.

Sample J1870, post-tectonic charnockite, Schneide, Dronning Maud Land Sample J1870 was collected from a post-tectonic charnockitic granitoid at Schneide. The sample consists of perthitic K-feldspar, plagioclase, minor quartz, and relics of altered ortho- and clinopyroxene and hornblende. K-feldspar contains large patchy mesoperthite exsolutions of plagioclase and is typically surrounded by myrmekite. Orthopyroxene has a reaction zone consisting of hornblende, plagioclase and opaque minerals. Apatite and allanite are accessory phases. The sample contains colourless to brown, and equant to acicular (aspect ratios up to 6:1) zircons and zircon fragments. Some zircons possess high-U cores and many contain abundant inclusions. Several zircons have long prismparallel cracks. CL images reveal strong oscillatory growth zoning (Fig. 7d), which was targeted for analysis; high-U cores were avoided. The analysed areas contain 69 – 440 ppm 238U and have Th/U of 0.44 – 0.81. Common Pb is

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Fig. 7. U–Pb analytical results for samples: (a, b) J3012/1; (c) J1670; (d) J1870; (e) JJ238; (f) JJ259; (g) GV01; (h) WB295.

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low ( f204 ,0.8%). Ten areas from 10 zircons were analysed (Table 2). The measured compositions are concordant to slightly discordant and yield a concordia age of 501+ 7 Ma (MSWD ¼ 1.12), which is interpreted as the crystallization age of this charnockitic granite.

Sample WB295, mangerite, Lurio Belt, northern Mozambique WB295 was sampled from a mangerite intrusion immediately to the north of the Lurio Belt within the Unango Complex (Fig. 3). The sample contains large (up to 400 mm), stubby (length –width ratios ,2), clear, colourless to light brown zircons, many of which contain large inclusions of other minerals. Although some are euhedral, most zircons are anhedral and irregular in shape. In CL, irregular and sector zoning is weakly to moderately defined in most zircons. Some exhibit weak oscillatory growth zoning, and several zircons show reaction zones. Eighteen areas were analysed in 18 zircons (Fig. 7h). All analyses are low in U (20–100 ppm) and Th (10–85 ppm) and have Th/U between 0.3 and 1, typical of magmatic zircon. Common Pb ranges up to 1.1%. Excluding two significantly younger analyses (8.1 and 6.1), 16 analyses form a mainly concordant group with a concordia age of 512 + 4 Ma (MSWD ¼ 1.2), which is interpreted as the crystallization age of the mangerite.

Sample JJ238, granite gneiss, Gurue`, northern Mozambique Sample JJ238 is a medium- to coarse-grained, highly foliated biotite granite gneiss collected close to Gurue in the Nampula Complex, not far from the boundary with the Lurio Belt. The sample contains trace amounts of titanite, allanite, apatite and opaque minerals. Zircon is abundant and occurs as a simple population of clear and elongate crystals, up to 200 mm long and with length–width ratios of up to six. Several have rounded terminations, and many contain mineral inclusions. In CL, the cores of most zircons exhibit concentric growth zoning, and are surrounded by thin reaction zones and overgrown by thin zircon rims. Some cores show evidence of resorption and alteration prior to overgrowth by zircon rims. The oscillatory zoned parts of 16 different zircons were analysed; the rims were not sufficiently wide to be analysed (Fig. 7e). Uranium concentration ranges from c. 140 to 410 ppm, and Th/U from 0.3 and 3.2, typical for igneous zircons. The common lead is uniformly low (see value of f204 in Table 2 with a mean of 0.07%). Fourteen

83

analyses form a concordant group with a concordia age of 507 + 3 Ma (MSWD ¼ 1.4), which we interpret as the crystallization age of the protolith of the biotite gneiss. Two slightly younger analyses (10.1 and 11.1) may reflect minor Pb loss from the analysed areas, possibly at the time the thin zircon rims were formed, during deformation of the gneiss.

Sample JJ259, porphyritic biotite–hornblende granite, Gurue`, northern Mozambique Sample JJ259 was collected from a small (a few kilometres long) elongate porphyritic granite pluton in the Nampula Complex. K-feldspar phenocrysts are up to 6 cm in length and locally show a well-defined igneous flow foliation. The sample contains biotite and amphibole and trace amounts of allanite and apatite. Zircons are equant to elongate and are characterized in CL images by an oscillatory-zoned core and a U-rich rim, commonly growing at the expense of the core. Fifteen analyses were made in 10 crystals. Uranium concentration ranges from 68 to 1716 ppm, and Th/U from 0.08 and 2.2. Common lead is generally low, but f204 is up to 1% for two analyses. No age difference can be detected between cores and rims. Twelve analyses define a concordia age of 508 + 4 Ma (Fig. 7f), interpreted as the intrusion age of the granite. One analysis (13b) in a core is significantly older (563 + 18 Ma), suggesting the presence of inherited material in the cores of the zircons. Two analyses in low-U cores are dispersed from the main group owing to the presence of significant common Pb.

Sample GV01, porphyritic biotite–hornblende granite, Ribaúe`, northern Mozambique Sample GV01 is a porphyritic granite, collected about 2 km north of Ribaúe` in the marginal part of a major, at least 60 km long, late-tectonic granite intrusion, roughly aligned with the regional ENE –WSW regional fabric. At the sampling locality, a planar fabric is well defined by Kfeldspar phenocrysts (1– 3 cm in length), possibly produced by magmatic flow or solid-state deformation. The sample contains biotite, and minor amounts of amphibole, muscovite, titanite, allanite, apatite and opaque minerals. Zircon forms elongate, transparent to light brown, zoned crystals up to 300 mm long, with aspect ratios up to 5:1. The zircons have irregular outlines, rounded terminations, and abundant mineral inclusions. CL images reveal core –rim structures, with oscillatory-zoned cores and rims up to 30 mm wide. The rims might have grown during late intrusive deformation.

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Twenty-seven areas were analysed in 21 grains, including 10 analyses of rims and 17 of cores (Fig. 7g). The rims contain 670 –1018 ppm 238U, have Th/U of 0.08 –0.23, and low common Pb ( f204 ,0.7%). The cores contain 347– 1544 ppm 238 U, have Th/U of 0.3–1.7, and have mainly low common Pb (f204 ,0.18% for 16 of 17 analyses). The cores and rims are indistinguishable in age and combine to yield a concordia age of 508 + 2 Ma. This age is interpreted as bracketing the crystallization age of the granite and the late-intrusive deformation around its margin.

Geochemical characteristics and tectonic interpretation of the late- to post-tectonic magmatic province in Dronning Maud Land The late-tectonic magmatism in Dronning Maud Land occurred between c. 530 and 485 Ma. It began with subordinate gabbro intrusions at c. 530– 520 Ma (e.g. Jacobs et al. 2003a) and terminated with similarly small amounts of sheet-like granite intrusions at c. 485 Ma (e.g. Jacobs et al. 2003b). Between these events, the major pulse of igneous activity was relatively short-lived between c. 510 and 500 Ma. The two samples dated in this study, the Oddesteinen granite dated at 499 + 4 Ma and the Schneide charnockite at 501+ 7 Ma, were emplaced towards the end of this interval and consequently confirm this previously recognized trend. Probably more than 90% of the latetectonic igneous rocks were intruded during this time interval. Fewer published data are available on the Nampula Complex in NE Mozambique. The four samples dated in this study represent four different occurrences of late-orogenic granitoids; a mangerite pluton (WB295), a sheared granite orthogneiss (JJ238), a porphyritic granite sheet (JJ259) and a large, weakly deformed porphyritic granite pluton (GV01). The samples gave dates of 512 + 4, 507 + 3, 508 + 4 and 508 + 2 Ma, respectively, indicating that the main pulse of late-orogenic magmatism in the Nampula Complex and along the Lurio Belt was synchronous with, or slightly older than, in Dronning Maud Land. In Dronning Maud Land, the main magmatism is associated with coarse-grained to megacrystic, generally undeformed, granitoid –charnockite bodies, that are exposed over an estimated area of more than 15 000 km2 between 28E and 288E in central and eastern Dronning Maud Land (e.g. van Autenboer & Loy 1972; Shiraishi et al. 1983, 1994; Ohta et al. 1990; Moyes et al. 1993; Paech 2001; Owada et al. 2003; Paech et al. 2004; Roland 2004a, b; Bisnath et al. 2006). A compositional

range from granite, quartz monzonite, monzonite, syenite and minor anorthosite is recognized, typically associated with reddish brown weathering charnockites, which are volumetrically the most abundant rocks exposed. The charnockites are composed of quartz, mesoperthite and plagioclase, with (especially in magnesian varieties) primary anhydrous mafic phases of pigeonite and augite. Ironrich varieties are typified by coexisting fayalite þ hedenbergite þ quartz (Frost & Bucher 1993). The original mafic phases are very often replaced or overgrown by hornblende and biotite, indicating hydration by late-magmatic fluids. Thermobarometrical studies by Frost & Bucher (1993) and Bucher & Frost (1995), and petrological studies by Markl & Henjes-Kunst (2004), suggest that the charnockites were emplaced as relatively dry melts at temperatures exceeding 900 8C and pressures of c. 4.8 kbar. When the charnockites cooled to below 800 8C, the remaining melt pockets became water saturated, causing hydration of the mineralogy. Furthermore, inclusions of gneiss and adjacent host rocks expelled fluids to flux the remaining melt and led to the crystallization of mafic hydrous silicates. These late hornblende– biotite granitic melts expelled aqueous fluids during their crystallization, which invaded adjacent parts of charnockite plutons and altered them to biotite–hornblende granodiorite. Such processes produced bleached zones in the otherwise brownish charnockites. Several papers on the bulk geochemical characteristics of the charnockites in Dronning Maud Land and their retrogressed varieties have been published (e.g. Klimov et al. 1964; Ravich & Soloviev 1966; Joshi et al. 1991; Roland 2002, 2004a, b; Jacobs et al. 2003a, b; Li et al. 2003; Paulsson & Austrheim 2003; Bauer & Jacobs 2005; Engvik et al. 2005), but the interpretation of the results and the implications for the geotectonic setting remain ambiguous. Geochemically, the rocks can be characterized as peraluminous to metaluminous, or mildly alkaline granites. In contrast to the contemporaneous, Mg-rich, charnockites of India and Sri Lanka the rocks from Dronning Maud Land are very enriched in Fe, manifested as coexisting ferrosilite and/or fayalite and quartz. Most of the Dronning Maud Land charnockites and related granitoids are enriched in Ba, Sr, Zr, Y, Zn and Fe, but they are relatively low in Cs, Rb and Ca normalized to a primitive mantle composition (e.g. Markl & Henjes-Kunst 2004; Roland 2004a, b). That is, they represent highly fractionated melts from a mainly mafic to intermediate source but are not typical A-type granitoids as originally defined by Loiselle & Wones (1979). Nevertheless, on various discrimination diagrams, the rocks broadly plot into the ‘within-plate’ tectonic setting field of the granitoids, although such

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85

diagrams were defined for A-type granitoids, which represent differentiates of magmas derived from mantle material (Eby 1992). Eby (1992) defined this subtype to differentiate intra-plate rift-related (A1-type) granitoids from post-collisional (A2type) granitoids. The geochemical signature of the charnockites and related granitoids in central Dronning Maud Land coincides with the field of A2-type granitoids (Fig. 8), which represent magmas derived from continental crust of tonalitic to granodioritic composition or underplated crust. According to Roland (2002, 2004a, b) the charnockites were derived from lower continental or underplated crust; this derivation would explain their relatively heterogeneous geochemical signature. The megacrystic, sometimes rapakivi-type textures often seen in such rocks was recently interpreted by Bonin (2007) as an indicator for an emplacement at a middle crustal level near the ductile –brittle transition zone (3–5 kbar). Partial melting of a mafic source of charnockitic magma requires supply of excess heat in a thickened lithosphere (Bonin 2007). The extensive high-temperature dry melt generation in the lower crust could have been accomplished by continental lithospheric mantle delamination after the collision between parts of East and West Gondwana, with the heat source provided by the large-scale uprise of hot asthenosphere (Fig. 9). The resulting highly Fig. 9. Delamination model of the southern part of the EAAO can explain the structurally defined extent of large volumes of high-temperature, A2-type granitoids within the southern part of the EAAO. (a) Collision of part of East and West Gondwana; (b) delamination of the orogen root and influx of asthenosphere (arrows); (c) generation of voluminous melts and orogenic collapse of the orogen. Cross-section is approximately east–west across the Nampula Province.

Fig. 8. Composition of charnockites and related granitoids from Dronning Maud Land (DML) plotted in the Y– Nb– 3Ga diagram after Eby (1992). Four analyses of samples from Mozambique are plotted as filled triangles.

fractionated charnockitic magmas intruded the high-grade metamorphic basement at c. 500 Ma and cooled to below c. 300 8C only by around 460–450 Ma giving a relatively slow cooling rate of 10 8C Ma21 (Markl & Henjes-Kunst 2004). This slow cooling rate was presumably due to the high volumes of hot melt invading the lower to middle crust. The spatial distribution of the granitoid– charnockite plutons appears to be structurally controlled (i.e. within the boundaries of the southern third of the EAAO), effectively terminating against the Lurio Belt in the north. Markl & Henjes-Kunst (2004) recorded significant variation in the Nd isotopic composition of the post-tectonic plutons of Dronning Maud Land, reflecting differing crustal components in their parental melts. For example, 1Nd values of around 23

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in the eastern Mu¨hlig-Hofmann-Gebirge (Fig. 1) indicate a significant contribution from Mesoproterozoic crust in this area, whereas values of around 210 in the Orvinfjella suggest that Paleoproterozoic crustal components were present. Further isotopic studies could thus be used to test and map the age of protoliths in Dronning Maud Land. For example, the Mesoproterozoic orogenic belt fringing the Kaapvaal–Grunehogna craton possibly terminates between the eastern Mu¨hlig-Hofmann-Gebirge and Orvinfjella (Bauer et al. 2003).

Tectonic models How can the widespread late-tectonic magmatic event described here be accommodated within the current tectonic models for the southern part of the EAAO? (1) The early French–Mozambiquan work in northern Mozambique, summarized by Pinna et al. (1993), interpreted the thrust-bound packages to the north of the Lurio Belt as the remains of a major latest Mesoproterozoic–Early Neoproterozoic (1100–950 Ma) thrust pile that was overthrust SE along the Lurio Belt over the Nampula Province. In this model Late Neoproterozoic and Early Palaeozoic isotopic ages from various rocks were assigned to a mainly thermal event with little tectonic component. Clearly, from the recent new work and the plentiful Palaeozoic metamorphic ages from the Lurio Belt, this model is now untenable. (2) A similar tectonic model but with different timing has been suggested by Grantham et al. (2003, 2007), who inferred SE-directed thrusting of granulite-facies rocks along the Lurio Belt over the Nampula Complex (their ‘Nampula Subprovince’) to the south between c. 620 and 550 Ma. In this model, the Monapo and Mugeba klippen are interpreted as the exposed remains of this once more widespread nappe complex. Grantham et al. furthermore suggested that Late Neoproterozoic nappes north of the Lurio Belt are part of a much larger nappe system, which extended into Sri Lanka and southern Madagascar to the east, and possibly to the Damara Belt of Namibia to the west. The Lurio Belt is interpreted as a major suture in this model. Grantham et al. (2007) suggested that the late granitoid magmatism in the Nampula Complex resulted from loading by the major thrust stack and it is argued that the time gap of c. 50 Ma between thrusting and granitoid intrusions is the normal delay time for the crust to heat sufficiently to generate these melts. In this model, the Schirmacher nappe in Dronning Maud Land could represent an extension of the aforementioned nappe to the south (e.g. Ravikant et al. 2004).

However, there are a number of possible problems with this model. First, the klippen overlying the Nampula and Dronning Maud Land crust are never seen to be intruded by the late-tectonic granites as might be expected. Second, the granitoids are typically charnockitic A-types and not the minimum-melt granitoids that might be expect to result from slowly heated crust. Finally, new dates indicate that there is significant young shearing along the Lurio Belt around 530–500 Ma (Bingen et al. 2006a, b), suggesting that the main movements on the structure are considerably younger than 620–550 Ma. (3) A third model, by the Norconsult Consortium (2007) infers an opposite, NW-directed tectonic transport direction, including thrusting and exhumation of the granulite-facies rocks at c. 620 – 550 Ma to the north of the Lurio Belt. This event was followed by a phase of extension at c. 530 – 490 Ma during which the Mesoproterozoic to Neoproterozoic basement terranes north of the Lurio Belt were juxtaposed with the Nampula Province along the Lurio Belt by extensional shearing, as indicated by the young metamorphic ages from the belt. Our present study sheds little light on the early exhumation history of the granulite-facies rocks but does contribute to the understanding of the later, mainly extensional history of the orogen. The late-tectonic granitoids intrude neither the granulite-facies klippen that overlie the Nampula Province nor the disconformably overlying latest Neoproterozoic metasedimentary rocks of the Alto Benfica and Mecuburi Groups. This suggests that the emplacement of the granulitefacies klippen and the Alto Benfica and Mecuburi Group metasediments occurred very late in the tectonic history, and is probably related to late orogenic collapse, which possibly continued up to c. 500 Ma. Furthermore, the late-tectonic granitoids represent high-T melts (charnockites, mangerites), which, at least in Dronning Maud Land, were immediately preceded by gabbroic intrusions. These high-T melts clearly require a major, possibly external, heat source. At the same time, the intruded country rocks show a steep isothermal decompression path, indicating rapid uplift without adequate cooling (Jacobs et al. 2003b; Colombo & Talarico 2004). This process can possibly be better explained by generation of the magmas in an elevated thermal environment brought about by partial delamination of the orogenic root, followed by the influx of hot asthenosphere and accompanied by rapid mechanical thinning. In this model the Lurio Belt could represent an accommodation zone between two thermomechanically very different parts of the orogen, rather than a suture zone.

GRANITES IN EAAO

Summary and conclusions A widespread late- to post-tectonic Cambrian magmatic province is recognized in the southern part of the EAAO, within northeastern Mozambique and central Dronning Maud Land, two areas that were thought to be contiguous within Gondwana. It covers an area of more than 15 000 km2, and would have stretched from the northern margin of the Nampula Complex (the Lurio Belt), in northeastern Mozambique, to central Dronning Maud Land, decreasing gradually westwards in volume to the eastern Sverdrupfjella, where the magmatism stops close to the frontal zone of the orogen. Extensional tectonics and late-tectonic magmatism is recorded between c. 530 and 485 Ma. The early stage (c. 530–520 Ma) is characterized by minor gabbro intrusions, followed by the main charnockite– granitoid magmatic event, which was short-lived between c. 510 and 500 Ma. The Cambrian plutons are mostly hosted within juvenile Mesoproterozoic (1.2–1.05 Ga) high-grade gneisses. The geochemistry of the granitoids is broadly ‘anorogenic’, ‘within-plate’ and ‘A2-type’ as defined from various geochemical discriminants. The intrusions are interpreted to have crystallized at mid-crustal levels during collapse and extension of the orogen, possibly accompanied by delamination of the lithosphere root. Hot asthenosphere, rising to the base of the lower crust above the subsiding orogenic root, would have provided the heat source for the magmatism, which is typically anhydrous, hightemperature charnockitic (especially in Dronning Maud Land). The large volumes of late-tectonic magmatism are confined to the southern third of the EAAO, bounded to the north by the Lurio Belt. The spatial extent of the magmatism is structurally controlled, which makes it unlikely that a simple mantle plume model could be invoked. The timing of the proposed collapse –extension event is constrained by new age data from metamorphic rims, which grew during shearing in synextensional late ductile mylonites. These rims, which surround cores with Mesoproterozoic protolith ages, were dated at 507 + 9 Ma, statistically indistinguishable from the ages of the main granitoid intrusions in Dronning Maud Land and northern Mozambique, which gave ages ranging from 512 to 499 Ma. This confirms that some of the extensional shearing is contemporaneous with the voluminous late-tectonic magmatism. Thus far, it is unclear whether the southdirected escape tectonic model of Jacobs & Thomas (2004) is related to the proposed delamination, as there are at present no robust age constraints for the escape tectonic event. The Lurio Belt is certainly an important boundary zone and might represent a mid- to lower-crustal

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thermomechanical accommodation zone separating the southern part of the EAAO with a delaminated orogen root from crust to the north, where delamination did not occur. We acknowledge constructive reviews from G. Grantham, M. Owada and R. J. Stern. Part of the research was funded through a Heisenberg Fellowship to J.J., DFG Ja 617/16. J.J. also acknowledges support through a Gledden Fellowship at the Tectonics Special Research Centre (TSRC), University of Western Australia, Perth. The work in Mozambique is part of a Mineral Resources Management Capacity Building Project, Republic of Mozambique, financed by the Nordic Development Fund and the World Bank. U – Pb analyses were conducted using the SHRIMP II ion microprobe at the John de Laeter Centre of Mass Spectrometry in Perth, Australia, which is operated by a university – government consortium, with the support of the Australian Research Council, and also at the NORDSIM laboratory, operated and financed under an agreement between the research councils of Denmark, Norway and Sweden, the Geological Survey of Finland, and the Swedish Museum of Natural History. L. Ilynsky, K. Lindeń and M. Whitehouse are thanked for operating the NORDSIM laboratory. R.J.T. acknowledges permission to publish from the Executive Director, BGS. M.T.D.W. publishes with permission of the Executive Director, Geological Survey of Western Australia. This is TSRC Publication 414 and NORDSIM Publication 196.

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Terrane correlation between Antarctica, Mozambique and Sri Lanka; comparisons of geochronology, lithology, structure and metamorphism and possible implications for the geology of southern Africa and Antarctica G. H. GRANTHAM1, P. H. MACEY2, B. A. INGRAM1, M. P. ROBERTS3,4, R. A. ARMSTRONG5, T. HOKADA6, K. SHIRAISHI6, C. JACKSON7, A. BISNATH8 & V. MANHICA9 1

Council for Geoscience, P/Bag X112, Pretoria, South Africa (e-mail: [email protected]) 2

Council for Geoscience, Bellville, South Africa

3

MSSP-Geomap Project, c/o Geological Survey of Papua New Guinea, Port Moresby, Papua New Guinea 4

Council for Geoscience, Walmer, Port Elizabeth, South Africa

5

RSES, Australian National University, Canberra, A.C.T. 0200, Australia 6

National Institute of Polar Research, Itabashi, Tokyo, Japan

7

51 Saint David’s Road, Claremont, Cape Town, South Africa

8

Caracle Creek International Consulting Inc., Johannesburg, South Africa 9

Direcaõ Nacional Geologia, Maputo, Mozambique

Abstract: Analysis of new lithological, structural, metamorphic and geochronological data from extensive mapping in Mozambique permits recognition of two distinct crustal blocks separated by the Lurio Belt shear zone. Extrapolation of the Mozambique data to adjacent areas in Sri Lanka and Dronning Maud Land, Antarctica permits the recognition of similar crustal blocks and allows the interpretation that the various blocks in Mozambique, Sri Lanka and Antarctica were once part of a mega-nappe, forming part of northern Gondwana, which was thrust-faulted c. 600 km over southern Gondwana during amalgamation of Gondwana at c. 590–550 Ma. The data suggest a deeper level of erosion in southern Africa compared with Antarctica. It is possible that this thrust domain extends, through the Zambezi Belt or Valley, as far west as the Damara Orogen in Namibia with the Naukluft nappes in Namibia, the Makuti Group, the Masoso Suite in the Rushinga area and the Urungwe klippen in northern Zimbabwe, fitting the mega-nappe pattern. Erosional products of the mountain belt are now represented by 700– 400 Ma age detrital zircons present in the various sandstone formations of the Transantarctic Mountains, their correlatives in Australia, as well as the Urfjell Group (western Dronning Maud Land) and probably the Natal Group in South Africa.

‘What’s in a name? That which we call a rose by any other name would smell as sweet’ (Romeo and Juliet (II, ii, 1–2), Shakespeare, c. 1595). This paper could equally have been titled ‘The nature and extent of the Lurio Belt inferred from the geochronology, structure, lithology and metamorphic histories of adjacent crustal blocks’ or alternatively ‘The errant hitchhiker terranes of northern Gondwana’.

In 2000 an ambitious project to map Mozambique was initiated by the Mozambique Government The project was funded by the World Bank, the Nordic Development Fund and the governments

of South Africa and Mozambique. The Norwegian Geological Survey (NGU) and British Geological Survey (BGS) consortium assumed responsibility for mapping most of northeastern Mozambique (Fig. 1). A consortium including the Finnish Geological Survey (GTK) and a private company assumed responsibility for most of northwestern, central and southern Mozambique (Fig. 1) whereas the Council for Geoscience of South Africa assumed responsibility for the mapping of eleven 11 degree sheets (Fig. 1). Nine of these


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Fig. 1. Map of Mozambique showing the areas mapped by the various groups. CGS, Council for Geoscience of South Africa; NGU/BGS, Norwegian Geological Survey– British Geological Survey; GTK, Finnish Geological Survey.

sheets were located in NE Mozambique and two in NW Mozambique. The funding for mapping of the areas of northern Mozambique mapped by the NGU–BGS and GTK teams came from the Nordic Development Fund. The project is now complete and the potential implications of some of the new data gathered in the mapping exercise are presented in summarized form and interpreted here. A review of the basement geology of Mozambique (Grantham et al. 2003) highlighted the paucity of reliable geochronological data in Mozambique itself (approximately five single zircon determinations at the time), in contrast to neighbouring areas. It also highlighted that a potential crustal boundary existed between the apparently juvenile southern end of the Mozambique Belt and an older block in northern Mozambique, with the northern end having limited juvenile rocks with extensive reworking of Palaeoproterozoic and Archaean rocks. The current project along with other studies has resulted in the number of new, reliable zircon ages now exceeding 100. Some of these data are summarized here and interpreted along with data from surrounding areas in a Gondwana context. These data have been critical for the definition of the model presented here. Recognizing the reconnaissance nature of the mapping, the application of low- to high-resolution magnetic and

radiometric aerial surveys and Landsat data has also contributed significantly to the new interpretation presented here by confirming the Lurio Belt as a prominent geophysical and geological structure. The Lurio Belt is manifested in the field as a north to NW, intermediate to steeply dipping, zone of high strain. The more exact nature of the Lurio Belt (i.e. thrust zone, strike-slip zone, zone of pure shear?) is the subject of current debate and has not been completely resolved by the mapping programme in Mozambique. The Lurio Belt (LB) was first described by Jourde & Vialette (1980), who described it as a suture of a major Lurian orogen with diverging nappes to the north and south (Pinna et al. 1993). Later interpretations by Pinna and others (Pinna 1995) considered it to be a late southerly thrust synform. Pinna et al. (1993) and Sacchi et al. (2000) proposed extensions of the Lurio structure into the Zambezi Belt, both of which involved late south-directed reverse thrusting. Jamal (2005) described a complex history for the Lurio Belt, reporting that it has been affected by four phases of deformation under granulite- to amphibolite-facies conditions. The history recorded by Jamal (2005) is summarized as follows. An earlier deformation (D1) along the belt is represented by felsic segregations that commonly trace the D2/F2 folds. Map- and outcrop-scale F2 NE– SW-oriented isoclinal folds generally have subhorizontal fold axes. Fold attitudes are described as varying from subvertical upright to NW-dipping. Associated fold asymmetry suggests that these structures represent SE-directed thrusting. Continuous NW– SE shortening, with a dextral shear component, led to the refolding of the D2 isoclines about the F3 open to tight subvertical folds and the formation of an S3 axial-plane cleavage. D3 is interpreted as involving strike-slip shear under amphibolite-facies P–T conditions. S– C0 mylonitic fabrics observed throughout the Lurio Belt suggest a transpressive regime in response to an oblique compression. Thus the NE– SW Lurio Belt probably acted as a shear zone that accommodated high strain during a collision event that affected NE Mozambique. Viola et al. (2006) have recently questioned the interpretation of the Lurio Belt as a major suture. They reported that structures along the belt vary greatly, involving intense linear structures in the NE, becoming wider and less belt-like in the SW. They described tight to isoclinal folds with NNW-dipping axial planes and roughly down-dip stretching lineations. No clear kinematic indicators were observed. Strain accommodation, involving folding and conjugate shear zones within the Lurio Belt, is more intense than in the surrounding rocks. Evidence of SSE–NNW-directed regional

ANTARCTIC–MOZAMBIQUE –SRI LANKA CORRELATION

compression is pervasive. Viola et al. concluded that the Lurio Belt represents a belt of repeated activity and reworking and that the last strain increment reflects pure shear bulk flattening of the belt, lacking significant regional belt parallel simple shear. In contrast to the south-directed transport direction, they have inferred extensional collapse toward the WNW. Various workers have proposed that the Lurio Belt extends into Sri Lanka (Kro¨ner 1991, 2001; Grantham et al. 2003) and is represented there by the shear zone separating the Highland Complex from the Vijayan Complex. This correlation is supported by the proximity of Sri Lanka to northern Mozambique suggested by various Gondwana reconstructions (Lawver et al. 1998; Reeves & de Wit 2000). This paper summarizes lithological, geochronological, structural and metamorphic data and interprets them to suggest that the Lurio Belt represents a deep-crustal terrane boundary in northern Mozambique and that its possible extensions into Sri Lanka and Dronning Maud Land, Antarctica, as a low-angle thrust nappe complex, permit the recognition of various crustal blocks or terranes separated by correlatives of the Lurio Belt in varying attitudes.

Crustal structure of Mozambique Interpretation of aeroradiometric and aeromagnetic data supported by reconnaissance ground mapping by the various mapping teams has facilitated the recognition that NE Mozambique is divided into two dominant blocks separated by the Lurio Belt (Fig. 2) and its possible extensions westwards. For the purposes of this paper the block north of the Lurio Belt is termed the Namuno Block and that south of the Lurio Belt the Nampula Block. The Namuno Block comprises an accretionary stack of thrust-faulted complexes (Fig. 1) of varying age (Bingen et al. 2006; Bjerkgard et al. 2006; Viola et al. 2006). Mesoproterozoic complexes include the Unango, Marrupo, Naroto, Meluco and Angonia Complexes whereas Neoproterozoic complexes include the Xixano, Montepuez, Lalamo, M’Sawize and Muaquia Complexes (Bingen et al. 2006; Hollick et al. 2006; Thomas et al. 2006; Viola et al. 2006; Grantham et al. 2007a; Fig. 2). The thrust-faulted accretionary stack defined by these workers can be extended further west, via Malawi, to northwestern Mozambique, with Mesoproterozoic rocks of the Southern Irumide Belt forming the footwall to similar-age high-grade gneisses thrust westwards (Grantham et al. 2007a). The Southern Irumide Belt is similarly interpreted to be underlain by late Palaeoproterozoic basement

93

intruded by continental margin arc-related magmas between 1.09 and 1.4 Ga and strongly overprinted during the Pan-African Orogen (Johnson et al. 2006). The Southern Irumide Belt is also interpreted to comprise four shear zone bounded terranes, with the bounding faults having NW–SE-oriented strikes (Johnson et al. 2006). The Namuno Block thrust stacks recognized by Bingen et al. (2006) and Viola et al. (2006) and the Angonia Complex (Grantham et al. 2007a) are reported to have involved top-to-the-west and -WNW deformation whereas those further west are inferred to involve top-to-the-SW orientation (Johnson et al. 2006). It is readily apparent from the geophysical data that the whole thrust stack comprising the Mesoproterozoic to Neoproterozoic rocks from the Mozambique coast to the Angonia Complex in the west is itself sheared and rotated, with an apparent sinistral sense of rotation, where these rocks merge with the Lurio Belt and its possible extensions in the south. This indicates that the intense ductile strain deformation recorded in the north- to NW-dipping Lurio Belt either post-dated the amalgamation of the various complexes in the Namuno Block or was part of a synchronous, larger-scale, sinistral transpressional structure, with the ENE – WSW-oriented Lurio Belt being the central main shear zone bounded on the north by a westward (sinistral) directed accretionary stack. In contrast, the radiometric and aeromagnetic data for the area south of the Lurio Belt, the Nampula Block, do not show the same accretionary stack configuration. The Nampula Block is dominated by medium-grade migmatitic tonalitic orthogneisses and paragneisses and quartzofeldspathic orthogneisses that are complexly interfolded and intruded by undeformed to locally weakly deformed granitic intrusions. At least two granulite-grade klippen, the Monapo Complex and the Mugeba Complex (Pinna et al. 1993) are recognized overlying the Nampula Block. These two klippen contain high-grade granulite ortho- and paragneisses and have been regarded as remnants of a larger thrust sheet or sheets (Pinna et al. 1993). The granulite-grade klippen suggested by Pinna et al. (1993) in the vicinity of Nampula is not a klippen, but rather an area of sporadic in situ charnockitization that probably developed through the action of late-tectonic fluids (Macey et al. 2007). The Nampula Block is also partially transgressed by the north–south-oriented mylonitic Namama sinistral strike-slip shear zone (Cadoppi et al. 1987) in the SW (Fig. 2). The Namama Shear Zone appears to curve and disappear into NE –SW-oriented layer-parallel structures at its northern end. The lack of continuity of the Namuno Block nappe complex across the Lurio Belt implies that

94


Fig. 2. Map showing the crustal structure of northern Mozambique and adjacent areas. The location of the sinistral Namama Shear Zone (Nm) is shown east of Mocuba in the south of the Nampula Block and the approximate location of the Geci Group (Gc) is shown in far northern Mozambique east of Malawi. The map is compiled from Barr & Brown (1987), Bingen et al. (2006), Bjerkgard et al. (2006), Hollick et al. (2006), Thomas et al. (2006), Grantham et al. (2007a, b) and Macey et al. (2007).

the Lurio Belt is not just a crustal shear zone that has segmented a uniform block of crust but represents a fundamental boundary along which two different blocks have been juxtaposed. The exact location of extensions of the Lurio structure westwards through southern Malawi into NW Mozambique and on into the Zambezi Valley are less clear, as a result of much of the area being overlain by Karoo-age sedimentary and volcanic rocks as well as being transected by the southern extensions of the western limb of the East

African Rift. Pinna et al. (1993) Pinna (1995), Sacchi et al. (2000) and Grantham et al. (2003) have supported an extension of the Lurio Belt into the Zambezi Belt. During the current mapping programme it was recognized by Koistinen et al. (2006) and Westerhof (2006) that the Tete Complex is allochthonous and forms the hanging wall of a large southward-directed thrust fault. These workers also record a large southward-directed c. east–west shear zone c. 10 km south of Tete. Barr & Brown


(1987) also reported a major east–west-oriented shear zone, the Sanangoe thrust zone c. 40 km north of Tete with a top-to-the-north direction of transport. It is possible that these structures represent extensions of the Lurio Belt westwards. In conclusion, the mapping programme has facilitated the definition of the Lurio Belt as a major crustal boundary, which will be used to interpret the geochronological, structural and lithological variations described below.

Rock types Namuno Block The composition of the Unango Complex varies widely and consists of granitic gneisses, some locally charnockitic, with biotite–hornblende gneisses and quartzite. The metamorphic grade varies from amphibolite to granulite grade and the rocks are extensively migmatized (Bjerkgard et al. 2006). Aeroradiometric and aeromagnetic surveys suggest that the rocks of the Unango Complex continue southwestwards through Malawi and become the supracrustal Angonia Complex in NW Mozambique west of the Malawi border. In this area interlayered, dominantly quartzofeldspathic gneisses with subordinate tonalitic and metabasic gneisses have been reported (Grantham et al. 2007a). Rare metapelitic gneisses and marbles are also seen. The metabasic gneisses have chemistry typical of enriched mid-ocean ridge basalt (E-MORB) rocks (Grantham et al. 2007a). Intruded into the supracrustal rocks are monzonites, syenites, anorthosite and pyroxenites of mostly uncertain age. The supracrustal sequences have ages of c. 1100–1050 Ma whereas an undeformed to weakly deformed monzonite has been dated at c. 560 Ma (Grantham et al. 2007a). Metamorphic overprints dated at c. 550 Ma have been reported from zircon rims and metamorphic titanite (Grantham et al. 2007a). The Angonia Group gneisses are thrust-faulted over granites of the Southern Irumide Complex to the east (Grantham et al. 2007a). The Marrupa Complex is dominated by granitic to tonalitic gneiss, with mafic amphibolitic orthogneisses, quartzite and quartz–feldspar gneiss (Bjerkgard et al. 2006). The rocks are characterized by amphibolite-facies mineralogy. The geochemistry of the orthogneisses suggests that they are mediumto high-K calc-alkaline rocks with SiO2 ranging between c. 42 and 78 wt% and K2O ranging between 0.3 and 6.1 wt% (Bjerkgard et al. 2006). The Nairoto Complex consists of variably migmatized granitic to tonalitic orthogneisses with calc-alkaline compositions (Bjerkgard et al. 2006). Mineral assemblages

95

are typical of amphibolite-facies metamorphism (Bjerkgard et al. 2006). The Meluco Complex comprises mostly granitic to granodioritic orthogneisses. Mineral assemblages are typical of amphibolite-facies metamorphism (Bjerkgard et al. 2006). The Xixano Complex includes part of the Chiure Supergroup and autochthonous supracrustal gneisses described by Pinna et al. (1993), and comprises a variety of paragneisses including marble, biotite gneiss, mica schists, meta-arenites, granitic to tonalitic gneisses and amphibolites (Bjerkgard et al. 2006). The metamorphic grade within the Xixano Complex is amphibolite facies to granulite facies (Bjerkgard et al. 2006). The Muaquia Complex comprises granitic, tonalitic and gabbroic gneisses, amphibolites, mica gneiss and calc-silicate gneisses, and is predominantly mafic to intermediate in composition (Bjerkgard et al. 2006). The rocks have mineralogy typical of dominantly amphibolite-facies metamorphism (Bjerkgard et al. 2006). The M’Sawize Complex comprises banded migmatitic gneisses, granulitic gneiss and mafic granulite with mineralogy typical of granulitefacies metamorphism (Bjerkgard et al. 2006). The M’Sawize Complex comprises part of the Msawize Group of Pinna et al. (1993), who included their unit as part of the Lurio Supergroup. The Lalamo Complex contains a variety of paragneisses including marble, biotite gneiss, mica schists, meta-arenites and granitoid gneisses with mineralogy typical of amphibolite-facies metamorphism (Bjerkgard et al. 2006). The Montepuez Complex was previously defined as part of the large Chiure Group by Pinna et al. (1993) and contains granitic to granodioritic gneiss, biotite gneiss and marbles with mineral parageneses typical of amphibolite-facies metamorphism (Bjerkgard et al. 2006). The Ocua Complex comprises rocks defined as the Lurio Supergroup by Pinna et al. (1993) and consists of a tectonic me´lange (Thomas et al. 2006). The main rock types are mostly granulitic gneisses of tonalitic, dioritic and granitic composition, amphibolitic and granulitic gneisses as well as ultramafic and metaluminous gneiss (Bjerkgard et al. 2006). The high strains characteristic of the eastern Lurio Belt become less distinct to the SW (Viola et al. 2006).

Nampula Block The description of the rock units of the Nampula Block is summarized from Macey et al. (2007), and Grantham et al. (2007b), who confirmed descriptions by earlier workers including Pinna et al. (1993) and Sacchi et al. (1984) amongst others. Six lithostratigraphic units are recognized in the Nampula Block. They comprise the Mesoproterozoic gneisses of the Mocuba Suite, the Culicui

96


Suite, the Rapale Gneiss, the Mamala Gneiss, the Molocue Group and the Cambrian granites of the Murrupula Suite. The Mocuba Suite consists dominantly of migmatitic banded tonalitic gneisses and subordinate mafic rocks with amphibolitefacies mineralogy. Compositions are dominantly calc-alkaline. Crystallization ages of c. 1125 Ma have been recorded (Macey et al. 2007). Not only does the strongly migmatized character of the Mocuba Suite distinguish it from the other Mesoproterozoic gneisses, it also indicates that a Mesoproterozoic orogenic event was experienced by these rocks of the Nampula Block. The Rapale Gneiss is of similar tonalitic composition but is clearly intrusive into the Mocuba Suite and has crystallization ages of c. 1090 Ma. The Culicui Suite is dominated by megacrystic, typically highly sheared, augen gneisses, which locally, in low-strain zones, preserve charnockitic mineralogy. In general, however, the metamorphic assemblages are typical of amphibolite-facies metamorphism. Crystallization ages from the Culicui Suite range typically between c. 1070 and 1090 Ma. The Mamala Gneiss is relatively uniform equigranular medium- to fine-grained leuco-quartzofeldspathic gneiss with uniform field and geophysical characteristics. Crystallization ages from the Mamala Gneiss are c. 1090 Ma. The Molocue Group comprises a banded interlayered sequence of dominantly quartzo-feldspathic para- and orthogneisses with subordinate amphibolites and calc-silicates. Three younger pre-Gondwana breakup lithological entities are recognized in the Nampula Block: the Murrupula Suite, the Mugeba Complex and the Monapo Complex. The Mugeba Complex is dominated by granulite-grade garnet–pyroxene intermediate orthogneisses with subordinate garnet– pyroxene metabasic granulites and garnet–sillimanite–rutile metapelitic gneiss. The Monapo Complex contains mostly granulite-grade banded supracrustal gneisses with subordinate Grt–Sil–Rtbearing metapelites. Intruded into the supracrustal gneisses are weakly deformed to apparently undeformed clinopyroxenites, syenites, granite and carbonatite (Siegfried 1999; Grantham et al. 2007b). Crystallization ages of c. 635 Ma are recognized from the Monapo Complex (Jamal 2005; Grantham et al. 2007b), whereas discordant zircons from the Mugeba Complex suggest a c. 1000 Ma protolith. Metamorphic ages from these complexes are c. 635 Ma and c. 580 Ma (Grantham et al. 2007b; Macey et al. 2007). The Murrupula Suite comprises granitoid intrusions, mostly undeformed and emplaced as kilometre-scale plutons to metre-scale pegmatitic dykes (Macey et al. 2007). The compositions vary from syenitic to granitic and include equigranular medium-grained varieties to coarse-grained

porphyritic types. The chemistry of the intrusions varies from metaluminous A-type rocks to peraluminous mica granites (Macey et al. 2007). The A-type intrusions have chemistries typical of A2 granites (Eby 1992), which are interpreted as typically being generated post-orogenically (Bonin 2007) and from continental crust that has been through a cycle of continent–continent collision or island arc magmatism (Eby 1992). Crystallization ages of the intrusions vary from c. 495 to c. 530 Ma.

Discussion On a purely lithological composition basis there is little to indicate a major crustal boundary defined by the Lurio Belt. Broadly, the Namuno Block and the Nampula Block are dominantly underlain by quartzofeldspathic gneisses, with the Lurio Belt and related Ocua Complex rocks being characterized by strong geophysical signatures and evidence of high strain in the field. The reduced geophysical signature of the Lurio Belt in the SW could possibly result from the structure attaining a shallower dip and being folded. This possibility requires additional investigation. Lithological distinguishing factors between the Namuno and Nampula Blocks include (1) a higher prevalence of supracrustal rocks containing metapelites and marbles north of the Lurio Belt as well as (2) the subordinate, but significant, presence of relatively alkaline syenitic orthogneisses north of the Lurio Belt. The only lithological exceptions to these distinguishing factors are found in the Monapo and Mugeba Complex klippen overlying the Nampula Block.

Geochronology of Mozambique and surrounding areas In comparing the geochronology of the different areas, histograms with bin sizes of 50 Ma are used in conjunction with probability density distributions. The limitations of these two methods have been described in detail by Sircombe (2000). The limitation of the use of histograms alone is that they do not take the error estimate of an age into account whereas a limitation in the use of probability density distributions alone is that they do not quantify the number of ages recorded within a particular age or bin range.

Southeastern Africa Figure 3 shows igneous crystallization and metamorphic ages from the Namuno Block, comprising the area north of the Lurio Belt in NE Mozambique


97

Fig. 3. (a, b) Histograms and probability density curves of crystallization ages from the Namuno Block (a), comprising an area north of the Lurio Block (NLB), a southern Irumide Belt Block (SIB) and Malawi (Mal), and (b) Nampula Block. (c, d) Histograms and probability density curves of metamorphic ages from the Namuno Block (c), comprising an area north of the Lurio Block (NLB), a southern Irumide Belt Block (SIB) and Malawi (Mal), and (d) Nampula Block igneous crystallization and metamorphic zircons. The lines for 600 Ma and 1100 Ma are shown for reference in all the figures.

(NLB), Malawi (Mal) and the southern Irumide areas of southern Zambia and NW Mozambique (SIB) and the Nampula Block (see Fig. 2). Data available include those generated during the mapping programme (Bingen et al. 2006; Ma¨ntta¨rri et al. 2006; Grantham et al. 2007a, b; Macey et al. 2007) along with additional data from Costa et al. (1994), Kro¨ner et al. (1997, 2001), Sacchi et al. (2000), Manhica et al. (2001), Jamal (2005) and Johnson et al. (2005, 2006). The data are dominantly derived from sensitive high-resolution ion microprobe (SHRIMP), inductively coupled plasma mass spectrometry (ICP-MS) or thermal ionization mass spectrometry (TIMS) analyses, except for the Pb– Pb evaporation data of Kroner et al. (2001) from southern Malawi. A few mineral – whole-rock Sm – Nd data are included from the Southern Irumide Block. In addition, only the data from the Southern Irumide Belt of Johnson et al. (2005, 2006) have been used for the geochronological comparisons below. The data used for the crystallization age

analysis from the Namuno Block are shown in Table 1 whereas the age data from the Nampula Block are summarized in Table 2. Comparison of the histograms and probability density distribution curves shows that all areas have broadly extensive Mesoproterozoic crystallization ages of c. 900–1150 Ma as well as Neoproterozoic–Cambrian ages of c. 650–450 Ma. However, a significant difference is that the Namuno Block (Fig. 3a) is characterized by crystallization ages from c. 650 to 900 Ma, which are absent from the Nampula Block (Fig. 3b). In the Nampula Block three ages of c. 630 Ma are recorded from the Monapo (two) and Mugeba klippen (Jamal 2005; Grantham et al. 2007b; Macey et al. 2007). These data clearly demonstrate that the Monapo and Mugeba klippen have age characteristics similar to rocks exposed north of the Lurio Belt in the Namuno Block. Other rock types with ages of c. 800 Ma possibly located south of Lurio Belt extensions are the granitic rocks intruded into the Rushinga area of NE Zimbabwe and western Mozambique

98


Table 1. Crystallization age data used to construct figures for the Namuno Block from the southern Irumide Belt and Malawi Unit and sample number Biotite–hornblende gneiss (MA16) Pelitic paragneiss (MA8) Chewore Ophiolite plagiogranite (sample SJ106.1) Kaourera Arc meta-dacite (sample SJ220) Kadunguri Whiteschists Chewore Inlier Granulite Terrane (sample ADC) Chewore Inlier Zambezi Terrane orthogneiss (sample AF) ZM007 meta-dacite Chongwe River CH6 banded mafic gneiss Chowe River CH7 meta-tuff Chowe River CH7 meta-tuff Chowe River CH9 meta-dacite Chowe River CH9 meta-dacite Chowe River CH10 K-feldspar augen gneiss Chowe River CH10 K-feldspar augen gneiss Chowe River Charnockite associated with Chipera gabbro–anorthosite Garnet –spinel–cordierite gneiss (Chipata Gneiss) Porphyritic granite (EP26 Petauke–Sinda Terrane) Deformed K-feldspar augen gneiss CHP2a Undeformed syenite CHP2c Moderately deformed coarsegrained syenite CHP3 Opx-bearing granulite from Madzimoyo quarry CHP4a Garnet –opx-bearing mafic layer Madzimoyo quarry CHP4b Opx-bearing granulite roadside Madzimoyo quarry CHP5 Coarse-grained hbl–biotite equigranular granite CHP6a Coarse-grained K-feldspar porphyritic granite CHP8 Coarse-grained K-feldspar porphyritic granite CHP10 Foliated or banded qtz–feldspar migmatite, leucosome portion CHP11a K-feldspar porphyritic granite CHP12 K-feldspar porphyritic granite CHP13 Magmatically banded K-feldspar porphyritic granite PS17 Coarse-grained bt-poor qtz–plag granite PS18

Method

Age (Ma) Error

Block

Source

SHRIMP

664

27

S. Malawi

Kro¨ner et al. (2001)

SHRIMP SHRIMP

576 1393

11 22

S. Malawi S. Irumide Belt

Kro¨ner et al. (2001) Johnson et al. (2005)

SHRIMP

1082

7

S. Irumide Belt

Johnson et al. (2005)

SHRIMP SHRIMP

1066 1071

21 8

S. Irumide Belt S. Irumide Belt

Johnson et al. (2005) Johnson et al. (2005)

SHRIMP

1083

8

S. Irumide Belt


SHRIMP

1088

20

S. Irumide Belt


SHRIMP

1051

12

S. Irumide Belt


SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP

1064 1037 1040 1105 1094

15 8 21 22 2

S. Irumide Belt S. Irumide Belt S. Irumide Belt S. Irumide Belt S. Irumide Belt

Johnson et Johnson et Johnson et Johnson et Johnson et

SHRIMP

1105

9

S. Irumide Belt


TIMS

1050

20

S. Irumide Belt


TIMS

1046

3

S. Irumide Belt


LA-ICP-MS

1125

15

S. Irumide Belt


SHRIMP

1046

4

S. Irumide Belt


SHRIMP SHRIMP

1050 543

7 6

S. Irumide Belt S. Irumide Belt


SHRIMP

1076

6

S. Irumide Belt


SHRIMP

1977

11

S. Irumide Belt


SHRIMP

1047

20

S. Irumide Belt


SHRIMP

1038

6

S. Irumide Belt


SHRIMP

1061

13

S. Irumide Belt


SHRIMP

1076

14

S. Irumide Belt


SHRIMP

1950

67

S. Irumide Belt


SHRIMP

1038

9

S. Irumide Belt


SHRIMP

1058

34

S. Irumide Belt


SHRIMP

479

9

S. Irumide Belt


SHRIMP

510

6

S. Irumide Belt


al. (2005) al. (2005) al. (2005) al. (2005) al. (2005)

(Continued)


99

Table 1. Continued Unit and sample number Fine-grained magmatically banded syenite PS19 Medium-grained, equigranular qtz–plag syenite PS 28 Undeformed equigranular coarse qtz–Kfs –bt granite PS65 Undeformed K-feldspar porphyritic granite PS71b Patch equigranular granite in coarse pegmatites PS73 Equigranular medium-grained qtz–plag –bt granite PS76 Strongly deformed quartzofeldspathic gneiss PS78 Garnet-bearing pelitic migmatite SZ16 Deformed qtz–feld gneiss with thin amphibolite SZ23 Progressively mylonitized porphyritic granite SZ25c Strongly deformed quartzofeldspathic gneiss SZ26 Strongly deformed hornblende – biotite gneiss SZ27 Tonalite Angonia Complex GGZ238 Metabasite Angonia Complex GGZ229 Monzonite GGZ256 Desaranhama Granite Monte Capingo Suite Sinda granite Monte Capirimpica Suite Cassacatiza Suite Monte Sanje Suite Granito Castanho Chipera Complex (Tete Suite) Macanga Granite Mussata Granite Ocua Complex Charnockite Marrupa Complex Tonalitic Gneiss Granitic gneiss (MA1) Diorite granulitic gneiss (MA2) Trondhjemite gneiss (MA3) Quartz monzonite gneiss (MA4) Charnockitic gneiss (MA6) Charnockitic gneiss(MA7) Pelitic paragneiss (MA8) Charnoenderbitic gneiss (MA9) Biotite–hornblende gneiss (MA10) Biotite–hornblende gneiss (MA13) Biotite–hornblende gneiss (MA13) Biotite–hornblende gneiss (MA14)

Method

Age (Ma) Error

Block

Source

SHRIMP

494

5

S. Irumide Belt


SHRIMP

495

10

S. Irumide Belt


SHRIMP

1043

14

S. Irumide Belt


SHRIMP

720

12

S. Irumide Belt


SHRIMP

504

7

S. Irumide Belt


SHRIMP

474

8

S. Irumide Belt


SHRIMP

742

13

S. Irumide Belt


SHRIMP

1984

21

S. Irumide Belt


SHRIMP

1008

17

S. Irumide Belt


SHRIMP

1023

12

S. Irumide Belt


SHRIMP

1961

31

S. Irumide Belt


SHRIMP

647

11

S. Irumide Belt


SHRIMP

1104

11

S. Irumide Belt

Grantham et al. (2007a)

SHRIMP

1058

11

S. Irumide Belt

Grantham et al. (2007a)

SHRIMP

SHRIMP SHRIMP

568 1041 1201 502 1086 1077 1050 1050 1047 470 1046 994 951

5 4 10 8 7 2 8 2 29 14 20 61 44

S. Irumide Belt S. Irumide Belt S. Irumide Belt S. Irumide Belt S. Irumide Belt S. Irumide Belt S. Irumide Belt S. Irumide Belt S. Irumide Belt S. Irumide Belt S. Irumide Belt N. of Lurio N. of Lurio

Grantham et al. (2007a) Ma¨ntta¨rri et al. (2006) Ma¨ntta¨rri et al. (2006) Ma¨ntta¨rri et al. (2006) Ma¨ntta¨rri et al. (2006) Ma¨ntta¨rri et al. (2006) Ma¨ntta¨rri et al. (2006) Ma¨ntta¨rri et al. (2006) Ma¨ntta¨rri et al. (2006) Ma¨ntta¨rri et al. (2006) Ma¨ntta¨rri et al. (2006) Macey et al. (2007) Grantham et al (2007b)

Pb/Pb Pb/Pb Pb/Pb Pb/Pb Pb/Pb Pb/Pb Pb/Pb Pb/Pb Pb/Pb

evap. evap. evap. evap. evap. evap. evap. evap. evap.

602.7 644.9 582.9 577.5 590.5 928.9 576.7 1012.5 998.9

1 0.9 1 1 1 0.9 1 0.8 0.8

Southern Southern Southern Southern Southern Southern Southern Southern Southern

Kro¨ner et al. Kro¨ner et al. Kro¨ner et al. Kro¨ner et al. Kro¨ner et al. Kro¨ner et al. Kro¨ner et al. Kro¨ner et al. Kro¨ner et al.

Pb/Pb evap.

738.7

0.9

Southern Malawi Kro¨ner et al. (2001)

Pb/Pb evap.

576.1

1


Pb/Pb evap.

1040.6

0.7


Sm –Nd

Malawi Malawi Malawi Malawi Malawi Malawi Malawi Malawi Malawi

(2001) (2001) (2001) (2001) (2001) (2001) (2001) (2001) (2001)

(Continued)

100


Table 1. Continued Unit and sample number

Method

Charnoenderbitic gneiss (MA15) Biotite–hornblende gneiss (MA16) Biotite gneiss (MA17) Biotite gneiss (MA17) Biotite gneiss (MA17)

Age (Ma) Error

Block

Source

Pb/Pb evap. Pb/Pb evap.

554.7 667.5

1 0.9

Southern Malawi Kro¨ner et al. (2001) Southern Malawi Kro¨ner et al. (2001)

Pb/Pb evap. Pb/Pb evap. Pb/Pb evap.

710.5 556.1 772.5

0.9 1 0.5

Southern Malawi Kro¨ner et al. (2001) Southern Malawi Kro¨ner et al. (2001) Southern Malawi Kro¨ner et al. (2001)

The data from north of the Lurio Belt are from Jamal (2005) and Bingen et al. (2006).

Table 2. Crystallization ages from the Nampula Block Rock type and sample number Granulite Migmatitic granite gneiss (sample MS5) Leucocratic granite (sample MS6) Augen gneiss sample NHF Tonalitic gneiss sample CVGN Granite Megacrystic charnockite Augen gneiss Tonalite Mocuba Suite Augen gneiss Mocuba Gneiss Augen gneiss Augen gneiss Tonalitic gneiss Augen gneiss Granitic gneiss Calc-silicate Granite Granitic gneiss Granite Augen gneiss Tonalitic gneiss Granite Syenite Granite Granite Granite NB21-8 Granite NB21-1 Granite NB5-1 Syenite Granulite

Unit Mocuba Complex

Nhansipfhe Megacrystic Gneiss Chimoio Granodiorite Gneiss Murrupula Suite Culicui Suite Culicui Suite Mocuba Complex Mocuba Complex Culicui Suite Mocuba Complex Culicui Suite Culicui Suite Rapale gneiss Culicui Suite Mamala Gneis Molucue Grp Murrupula Suite Molucue Grp Murrupula Suite Culicui Suite Rapale gneiss Murrupula Suite Murrupula Suite Murrupula Suite Murrupula Suite Murrupula Suite Murrupula Suite Murrupula Suite Ramiane Suite Monapo Complex

Additional data are available from Jamal (2005).

Method

Age (Ma)

Error

Source

SHRIMP SHRIMP

1028 1094

7 13

Costa et al. (1994) Kro¨ner et al. (2001)

SHRIMP

1009

13


SHRIMP

1112

18

Manhica et al. (2001)

SHRIMP

1108

12

Manhica et al. (2001)

SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP

495 1074 1082 1078 1123 1085 1129 1077 1092 1091 1076 1092 1127 533 1090 504 1073 1095 521 527 507 497 516 505 514 634 637

2 13 26 16 14 10 9 26 42 14 8 13 11 5 22 12 16 8 4 4 7 4 3 5 4 8 6

Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Macey et al. (2007) Grantham et al. (2007b) Grantham et al. (2007b)


(Barton et al. 1993; Dirks et al. 1998; Vinyu et al. 1999; Hargrove et al. 2003; Fig. 2) as well as the Guro Bimodal Suite (Ma¨ntta¨ri et al. 2006; Westerhof 2006). These rocks are located along the margin of the Zimbabwe Craton and consequently their relationship to the Mozambique Belt is uncertain. Koistinen et al. (2006) related the c. 850 Ma magmatism to extensional processes at the margin of the Zimbabwe Craton, whereas Westerhof (2006) suggested that these ages may be related to detachment thrusting. The c. 850 Ma ages are therefore geographically restricted to areas at or close to the Zimbabwe Craton and consequently are anomalous in the Nampula Block. The difference in ages between the Namuno and Nampula Blocks was recognized by Pinna (1995), although at that time the nature and origin of the age differences was unclear. Another important difference between the three areas is that the Nampula Block and southern Irumide area have some samples with crystallization ages between 1100 and 1200 Ma whereas these ages are absent in the north of the Lurio Namuno area. Another difference is that the southern Irumide area appears to have dominantly 1000–1100 Ma ages and fewer ages in the ,600 Ma range. It is uncertain whether this is real or an artefact of sampling. Metamorphic age data are summarized in Tables 3 (Namuno Block) and 4 (Nampula Block). Data sources include Kro¨ner et al. (1997, 2001), Jamal (2005), Johnson et al. (2005, 2006), Bingen et al. (2006), Grantham et al. (2007a, b), and Macey et al. (2007). Comparison of the metamorphic ages shows that from the Namuno Block (Fig. 3c) no evidence of Mesoproterozoic metamorphism has been recorded north of the Lurio Belt, with metamorphic ages in the Namuno Block being c. 750– 400 Ma with 1050–1100 Ma metamorphism being recorded in the southern Irumide area. In contrast, data from the Nampula Block (Fig. 3d) indicate that metamorphism occurred during the Mesoproterozoic between 1050 and 1100 Ma as well as during the time period c. 600–400 Ma. In conclusion, from the histograms shown in Figure 3, it is recognized that the Namuno and Nampula Blocks have different characteristics, some fairly distinct (e.g. the common presence of 600 –900 Ma ages in the Namuno Block and absence in the Nampula Block) and some subtle.

Antarctica Extending the patterns recognized from the Namuno and Nampula blocks to the adjacent areas of Dronning Maud Land (DML), Antarctica, in a Gondwana context (Fig. 4), the following aspects become apparent. The age distribution for the Nampula Block is virtually identical to that

101

observed in western DML (Sverdrupfjella þ Kirwanveggean) (Fig. 5a) with crystallization ages in both areas falling in the time periods 950 – 1200 Ma and 470 – 500 Ma and metamorphic ages being recorded for the periods 950 – 1100 Ma and 450 – 600 Ma (Fig. 5b). The data utilized in Figure 5a and b are summarized in Tables 5 and 6 with the data derived from Harris et al. (1995), Krynauw & Jackson (1996), Jackson & Armstrong (1997), Harris (1999), Jackson (1999), Board et al. (2005), Grantham et al. (2006) and G. H. Grantham & R. A. Armstrong (unpubl. data). In addition, equivalents of the megacrystic granitic augen gneiss Culicui Suite and heterogeneous medium-grained, equigranular tonalitic orthogneiss Mocuba Suite, both volumetrically significant lithological units in the Nampula Block, are recognized in western DML in the form of the Kirwanveggen megacrystic orthogneiss (Grantham et al. 1995) and the Kvervelknatten orthogneiss (Grantham et al. 1995, 1997; Groenewald et al. 1995; Wareham et al. 1998), respectively. The chemistries of these two units are comparable, as are the ages, which are typically c. 1070–1090 and c. 1110– 1140 Ma, respectively. Progressing further eastwards into central DML (excluding Schirmacher Hills), the western Mu¨hlig-Hofmannfjella has crystallization ages and metamorphic ages largely comparable with those of the Nampula and western DML areas of Antarctica (Fig. 5c and d). The data from central DML are summarized in Tables 7 and 8, with the data sources including Jacobs et al. (1998, 2003a–c), Paulsson & Austrheim (2003), Bisnath et al. (2006). The crystallization ages recorded for central DML are in the range of 450–650 Ma and 1050– 1200 Ma and the metamorphic ages 500–600 and 1000–1100 Ma. The 550–650 Ma crystallization ages from central DML are all collected from the extreme eastern end of central DML from charnockites and anorthosites in the Wolthaat Massif area and are not recognized in the western Mu¨hlig-Hofman Mountains. Charnockites and anorthosites of Neoproterozoic age are not recognized in the Nampula Block. In contrast, limited data from the granulites exposed in Schirmacher Hills in NE Mu¨hligHofmannfjella (Table 8) dominantly have ages in the range c. 550–700 Ma with a few older ages between 800 and 1150 Ma being recognized (Ravikant et al. 2004, 2008). Consequently, the Schirmacher Hills has ages comparable with those of the north of the Lurio Namuno Block (Fig. 5e). Further east in the Sør Rondane area of Antarctica, SHRIMP (Shiraishi et al. 2008) and chemical Th –U –total Pb isochron method (CHIME) data (Asami et al. 2005) indicate that the NE Sør Rondane has age distributions similar

102


Table 3. Metamorphic ages used for the Namuno Block Rock type and sample number

Method

Age (Ma)

Error

Area

Source

Charnockitic gneiss (MA8) Felsic granulite (MA12) Biotite–hornblende gneiss (MA13) Hofineir Gneiss deformed quartzofeldspathic gneiss Garnet-bearing pelitic migmatite Nyamadzi Gneiss deformed quartzofeldspathic gneiss Wutepo Gneiss deformed hornblende –biotite gneiss Titanite in mafic gneiss Ocua Complex MZ05045a charnockite Mugeba Complex

SHRIMP SHRIMP SHRIMP

572 547 564

9 10 4

S. Malawi S. Malawi S. Malawi

Johnson et al. (2005) Johnson et al. (2005) Johnson et al. (2005)

SHRIMP

536

10

S. Irumide area


SHRIMP SHRIMP

1942 1065

5 13

S. Irumide area S. Irumide area


SHRIMP

555

11

S. Irumide area


SHRIMP SHRIMP

549 555

7 5

S. Irumide area N. of Lurio

Grantham et al. (2007a) Macey et al. (2007)

SHRIMP

614

8

Mugeba


Additional data are from Jamal (2005) and Bingen et al. (2006).

to those of the Namuno Block (Fig. 5f). Almost all of these samples are from the NE Sør Rondane. The SW Sør Rondane is separated from the NE Sør Rondane by a c. 10 km wide shear zone (Shiraishi et al. 1991; Shiraishi & Kagami 1992). The crystallization ages from Sør Rondane range between 1200 and 500 Ma with most being between 500 and 650 Ma whereas metamorphic ages are between

500 and 650 Ma with no Mesoproterozoic metamorphism being recognized.

Sri Lanka Very few SHRIMP zircon U/Pb or single-grain zircon data from single suites are available from Sri Lanka. The available data (Kro¨ner et al. 1987,

Table 4. Metamorphic ages from the Nampula Block Unit Mugeba Complex Mocuba Suite Mocuba Suite Molocue Group Culicui Suite Culicui Suite Culicui Suite Rapale Gneiss Culicui Suite Mamala Fm. Leucosome S3 in Culicui Suite Mocuba Suite Rapale Gneiss Monapo Complex Monapo Complex

Rock type

Method

Granulite zircon rim Quartzofeldspathic gneiss zircon rim Migmatitic vein Quartzofeldspathic gneiss zircon rim Augen gneiss zircon rim Augen gneiss zircon rim Charnockite zircon rim Tonalitic gneiss lower intercept Augen gneiss zircon rim Quartzofeldspathic gneiss zircon rim Leucosome

SHRIMP SHRIMP

591 527

4 18

Macey et al. (2007) Macey et al. (2007)

SHRIMP SHRIMP

1063 502

47 90


SHRIMP SHRIMP SHRIMP SHRIMP

538 525 513 449

8 20 10 95

Macey et al. Macey et al. Macey et al. Macey et al.

SHRIMP SHRIMP

505 555

10 12


SHRIMP

490

8

Macey et al. (2007)

Tonalitic gneiss zircon rim Tonalitic gneiss lower intercept Granulite zircon rim Syenite zircon rim

SHRIMP

1090

34

Macey et al. (2007)

SHRIMP

608

42

Macey et al. (2007)

SHRIMP SHRIMP

579 596

11 5

Grantham et al. (2007b) Grantham et al. (2007b)

Additional data are available from Jamal (2005).

Age (Ma) Error

Source

(2007) (2007) (2007) (2007)


103

Fig. 4. Schematic map of the various crustal blocks belonging to the Namuno and Nampula-type blocks. VC, Vijayan Complex; CHC, Central Highland Complex; Mo, Monapo Klippen; Mg, Mugeba Klippen; Kg, Kataragama Klippen; L, Lurio Belt; SH, Schirmacher Hills; MH, Mu¨hlig-Hofmannfjella; GV, Gjelsvikfjella; HUS, Sverdrupfjella; Kvg, Kirwanveggen; Hmf, Heimefrontfjella; GC, Grunehogna Craton; Ur, Urungwe Klippen; PCM, Prince Charles Mountain.

1994; Baur et al. 1991; Holzl et al. 1994) are largely derived from multigrain TIMS studies on single rock units or come from SHRIMP studies on metasediments aimed at provenance determinations. None the less, histograms summarizing data from the Highland Complex (HC) (Fig. 5g) and Vijayan Complex (VC) (Fig. 5h) of Sri Lanka show that the available data from the HC have a similar pattern to those for the Namuno Block of Mozambique and the Sør Rondane and Schirmacher Hills areas of Antarctica, whereas the VC has a pattern of ages comparable to those of the Nampula Block and DML, Antarctica. The data reported from the Vijayan Complex by Kro¨ner et al. (1987) have large analytical uncertainties, resulting in the broad curves defined by the probability density distribution (Fig. 5h), whereas the absolute ages show age ranges of 500– 600 and 1000– 1250 Ma. In conclusion, the chronological data combined with lithological varieties facilitate recognition of two age groups of tectonic blocks; namely, those with significant volumes of rock with ages

between 600 and 900 Ma and those without (Fig. 4). The former group comprises the north of the Lurio Namuno Block, Malawi and southern Irumide Block as well as the Mugeba and Monapo klippen, the Highland Complex in Sri Lanka, the NE Sør Rondane, the far eastern Mu¨hlig-Hofmannfjella and the Schirmacher Hills. The latter group comprises the Nampula Block, the Vijayan Complex in Sri Lanka, the SW Sør Rondane, the western Mu¨hlig-Hofmannfjella, Sverdrupfjella and its extensions into Kirwanveggen (Fig. 4). In the remainder of the paper, we will refer to these grouped blocks as the Namuno and Nampula Blocks, respectively.

Structural data In most cases the boundaries between the Namuno and Nampula Blocks (when exposed) are defined by highly sheared rocks. In NE Mozambique the boundary is represented by the highly sheared Lurio Belt and the circumferential mylonites

104


Fig. 5. Histograms of geochronological data from western DML (a, b), central DML (c, d), Schirmacher Hills (e), Sør Rondane (f), Sri Lanka (Highland Complex) (g) and Sri Lanka (Vijayan Complex) (h). The 600 Ma and 1100 Ma lines are shown for reference. The data from Sør Rondane are subdivided into SHRIMP-based ages of crystallization and metamorphism and CHIME-based ages of crystallization and metamorphism.


105

Table 5. Crystallization ages from Sverdrupfjella and Kirwanveggen, western Dronning Maud Land Unit

Method

Age (Ma)

Error

Kyanite leucogneiss Bt–grt migmatite Intrusive leucogneiss Megacrystic orthogneiss Pegmatite vein Late felsic dyke Sveabreen Granite Fugitive Granite Fugitive Granite Roerkulten Granite Rootshorga Paragneiss Granite gneiss tonalitic Wbsv065 Tabular granite Wbsv073 Granite dykes Wbsv069 Kvervelkatten Gneiss Kvervelkatten Amphibolite Pod Cjk151 Megacrystic augen gneiss Cjk158 Leucopegmatite Cjk155 Porphyritic granite dyke Cjk 103 Leucogranite Cjk152 Granite dyke Cjk 159 Amphibolite dyke Hallgrens Cjk160 Banded bt gneiss Hallgrens Cjk56 Banded bt gneiss Hallgrens Cjk59 Titanite Cjk 159 Grey gneiss Dalmatian Granite Brattskarvet Monzonite Midbressrabben Diorite

SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP

1096 1157 1101 1088 1079 1011 1127 1131 1061 1103 1092 1132 1072 480 1134 1139 1074 1050 1011 990 980 986 1081 994 1003 1143 489 474 1140

10 10 13 10 6 8 12 25 14 13 13 16 10 10 11 10 11 10 8 12 13 6 4 22 9 11 10 10 10

Grey gneiss Jutulrora Jw4

SHRIMP

1139

12

Augen gneiss Sa 10

SHRIMP

1096

14

around the Monapo and Mugeba klippen. In the west the Sanangoe thrust zone (Barr & Brown 1987) and the shear zones defining the allochthonous Tete Complex (Koistinen et al. 2006; Westerhof 2006) represent possible extensions of the Lurio Belt. In Sri Lanka, the boundary between the Highland Complex is interpreted as a complex thrustfault zone (Kleinschrodt 1994) in which the granulite-facies Highland Complex has been thrust-faulted over the amphibolite-facies Vijayan Complex. In the shear zone, which is reportedly hundreds of metres wide, a strong north–south-oriented stretching lineation is developed; however, kinematic indicators are sparse. In Sør Rondane, NE Sør Rondane is separated from SW Sør Rondane by a c. 10 km wide shear zone (Shiraishi et al. 1991). The boundary in Mu¨hlig-Hofmannfjella is, however, not exposed. The differences in reported geochronology and

Source Harris (1999) Harris (1999) Harris (1999) Harris (1999) Harris (1999) Moyes & Harris (1996) Moyes & Harris (1996) Moyes & Harris (1996) Moyes & Harris (1996) Moyes & Harris (1996) Moyes & Harris (1996) Board et al. (2005) Board et al. (2005) Board et al. (2005) Jackson (1999) Jackson (1999) Jackson (1999) Jackson (1999) Jackson (1999) Jackson (1999) Jackson (1999) Jackson (1999) Jackson (1999) Jackson (1999) Jackson (1999) Jackson (1999) Krynauw & Jackson (1996) Krynauw & Jackson (1996) G. H. Grantham & R. A. Armstrong (unpubl. data) G. H. Grantham & R. A. Armstrong (unpubl. data) G. H. Grantham & R. A. Armstrong (unpubl. data)

lithologies in eastern Mu¨hlig-Hofmannfjella imply that the boundary between the two blocks probably passes immediately east of the Wolthaat Anorthosite Massif, the most easterly nunatak group in Mu¨hlig-Hofmannfjella, whose age and extensional affinity suggest that it belongs to the Namuno Block along with the granulites at Schirmacher Hills and Mramornye nunataks. This is possibly also supported by the differences in structural histories between the Wolthaat Massif and rocks further east described by Bauer et al. (2004). Immediately east of the Wolthaat Massif, the nunataks are reportedly underlain by c. 550 Ma granites after which, progressing eastwards, the lithologies are typical of the Nampula Block (Jacobs et al. 2003c; Bauer et al. 2004). The lithologies astride the Orvinfjella shear zone in Mu¨hlig-Hofmannfjella reportedly are not different (Jacobs et al. 2003c; Bauer et al. 2004) and the orientation of the Orvinfjella Shear Zone suggests that it may

106


Table 6. Metamorphic ages from Sverdrupfjella and Kirwanveggen, western Dronning Maud Land Unit

Method

Age (Ma)

Error

Source

Leucosome in garnet migmatite gneiss Kyanite leucogneiss Granite gneiss rim tonalitic wbsv065 Metapelite wbsv025 Metapelite wbsv025 Tabular granite rim wbsv073 Tabular granite rim wbsv073 Leucosome wbsv071 Leucosome wbsv071 lower intercept Leucosome wbsv074 Leucosome wbsv113 Leucosome wbsv113 Leucosome wbsv114 Leucosome wbsv116 Rim Kvervelknatten

SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP

1098 1096 1031 1044 540 565 996 1035 499 515 1032 503 525 519 1060

5 10 47 47 6 11 17 31 17 7 15 35 35 4 22

Leucosome CJK153 Monazite CJK 149 Titanite CJK56 Rim SA10

SHRIMP SHRIMP SHRIMP SHRIMP

1031 956 1015 538

6 17 16 25

Mafic dyke RK55

SHRIMP

523

21

Harris (1999) Harris (1999) Board et al. (2005) Board et al. (2005) Board et al. (2005) Board et al. (2005) Board et al. (2005) Board et al. (2005) Board et al. (2005) Board et al. (2005) Board et al. (2005) Board et al. (2005) Board et al. (2005) Board et al. (2005) Krynauw & Jackson (1996) Jackson (1999) Jackson (1999) Jackson (1999) G. H. Grantham & R. A. Armstrong (unpubl. Data) Grantham et al. (2006)

continue into Mozambique as the Namama ShearZone (Grantham et al. 2003), where it dissects similar rock types. Piazolo (2004) has described the structural evolution of the Mramornye nunataks and Schirmacher Hills. Both areas are characterized by strong shear fabrics, with those at Mramornye nunataks suggesting thrust-faulting toward the south during D3.

Planar fabric data Grantham et al. (2003) described one of the enigmas of the correlation of northern Mozambique with Dronning Maud Land as being the opposing structural facing directions in the two areas. Planar structures in northern Mozambique were described as dipping dominantly to the north and NW whereas those in western Dronning Maud Land were described as dipping dominantly to the SE. The improved densities of structural observation and geochronology permit the conclusion that the ages of fabrics in the two areas are not the same. The strong planar NW-dipping fabrics in and adjacent to the Lurio Belt clearly affect rocks with crystallization ages of c. 630 Ma along with older rocks in the Nampula Block. Consequently, the strong fabric-producing event in these areas has to be younger than c. 630 Ma. The structural data from the Mugeba and Monapo klippen, particularly

the latter, are discordant to their structurally underlying rocks, supporting their interpretation as klippen (Grantham et al. 2007b; Macey et al. 2007). The Monapo Complex, in particular, is interpreted as a circular synformal remnant in which at least two phases of deformation defined by the layered granulites form a large type 2 interference fold structure (Grantham et al. 2007b). In contrast to the younger than 630 Ma deformation in the klippen and the Lurio Belt, a detailed study in southern Kirwanveggen by Jackson (1999) showed that most of the deformation occurred there before c. 900 Ma. This is circumstantially supported in southern Kirwanveggen at Skappelknabben, where strongly sheared augen gneisses with greenschist-grade, planar fabrics are in relative close proximity to the virtually undeformed (brittlefaulted) sandstones of the Urfjell Group at Drapane. The Urfjell Group is less than c. 550 Ma old (Moyes et al. 1997), the age of the youngest detrital zircon recorded in it (Croaker 1999). Grantham et al. (2006) have shown that a c. 950 Ma mafic dyke at Roerkulten in Sverdrupfjella post-dates an earlier migmatitic fabric-forming event (D1), has a planar fabric (D2) that has been deformed about D3 folds, and was metamorphosed in the upper amphibolite facies at c. 500 Ma. Similarly, the syntectonic emplacement of granitic sheets at c. 480 Ma during top-to-the-SE-directed deformation in NW Sverdrupfjella (Grantham et al. 1991) demonstrates


107

Table 7. Crystallization ages for Mu¨hlig-Hofmannfjella in central Dronning Maud Land Rock unit or type and sample number

Method

Age (Ma)

Error

Source

Stabben gabbro Granite aplite dykes Augen gneiss Grey migmatite augen gneiss Granite gneiss Banded gneiss Granite gneiss Homogeneous migmatite Stabben syenite Lamprophyre dyke Risemedet/2312/2 Charnockite Hochlinfjellet 1301/2 Augen gneiss 2412/4 Grey migm gneiss 2712/4 Augen gneiss 1512/1 Grey gneiss 1701/2 Migmatitic augen gneiss Stabben gabbro Granite dyke Gygra Felsic gneiss j1671 Felsic gneiss j1704 Felsic gneiss j1795 Orthogneiss j1736 Orthogneiss j1797 Metagranodiorite j1698 Metaleucogranite j1695 Felsic gneiss j1838 Charnockite j1886 Anorthosite j1955 Anorthosite j1958 Zwiesel gabbro Zwiesel gabbro

SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SIMS SIMS SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP

487 497 1104 1124 1133 1091 1130 1163 500 523 521 1096 1115 1123 1142 1137 483 487 1073 1137 1076 1086 1087 530 527 1130 608 600 583 527 521

4 5 8 11 16 16 19 6 8 5 3 8 12 21 21 14 11 4 9 21 14 20 28 8 6 12 9 12 7 5 6

Bisnath et al. (2006) Bisnath et al. (2006) Bisnath et al. (2006) Bisnath et al. (2006) Bisnath et al. (2006) Bisnath et al. (2006) Bisnath et al. (2006) Paulsson & Austrheim (2003) Paulsson & Austrheim (2003) Jacobs et al. (2003a) Jacobs et al. (2003a) Jacobs et al. (2003b) Jacobs et al. (2003b) Jacobs et al. (2003b) Jacobs et al. (2003b) Jacobs et al. (2003b) Jacobs et al. (2003a) Jacobs et al. (2003a) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (2003c) Jacobs et al. (2003c)

a direction of deformation similar to that in Mozambique. These data support the work of Grantham et al. (1995), who concluded that there were two major periods of deformation in Sverdrupfjella in western DML; namely, one during the Mesoproterozoic at c. 900 –1000 Ma and the other during the Neoproterozoic and into the Cambrian at c. 550 –490 Ma. The data do not support the suggestion by Board et al. (2005) that the dominant deformation in Sverdrupfjella is Neoproterozoic to Cambrian in age. A more detailed study of planar structures (Figs 6–8) in the various areas of Mozambique and DML, Antarctica provides a better understanding of the structural relationships. In northern Mozambique, planar fabrics (Fig. 6a and b) dip and plunge dominantly toward the north to NW in the Lurio Belt as well as in the area along the southern margin of the Lurio Belt. Similarly the lineations in the same areas (Fig. 6e) define an arc with lineations plunging between north and west with concentrations towards the north and WNW. Limited fold-axis data are similar to the lineations

shown in Figure 6e. Progressing southwards toward the northern Mozambique coast the structural patterns become more complex and bimodal in nature, with both north- to NW-dipping planar structures as well as south- to SE-dipping structures being common (Fig. 6c and d). Similarly, the lineations along the southern margin of the northern Mozambique coast also show a bimodal variation with west- and ESE-dipping orientations (Fig. 6f). The data in Figure 5f are skewed by the high number of readings collected in the broad vicinity of the Namama shear zone by Aquater in the early 1980s (Aquater 1983). Bimodal patterns in the planar fabrics are observed in the western Mu¨hlig-Hofman Mountains of Antarctica (Fig. 7f; data from Jacobs et al. 2003a) and the Gjelsvikfjella (Fig. 7g, A. Bisnath unpubl. data), with gneisses dipping broadly to the NE and SW. In contrast, the structural data from Mramornye nunataks (Fig. 7h; from Piazolo 2004) is unimodal, with the gneisses dipping dominantly shallowly to steeply toward the ENE. Lineations from Gjelsvikfjella (Fig. 7a) show at least three

108


Table 8. Metamorphic ages from Mu¨hlig-Hofmannfjella and Schirmacher Hills Subject or sample and sample number Age of migmatization Zircon overgrowths Migmatite gneiss Lower intercept Aba/32 Leucosome-1301/2 Leucosome-0801/3 Grey gneiss 1701/2 Grey gneiss 1701/2 Augen gneiss 1512/1 Leucosome Rim j1704 Rim j1795 J1886 charnockite Anorthosite rim j1955 Metamorphic rim j1704 Whole-rock minerals Whole-rock minerals Whole-rock minerals Monazite Monazite Titanite Monazite Monazite Monazite Monazite Monazite Monazite Monazite Titanite Titanite

Method

Age (Ma)

Error

Source

SHRIMP SIMS SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP SHRIMP TIMS Sm –Nd TIMS Sm –Nd TIMS Sm –Nd TIMS U/Pb TIMS U/Pb TIMS U/Pb TIMS U/Pb TIMS U/Pb TIMS U/Pb TIMS U/Pb TIMS U/Pb TIMS U/Pb TIMS U/Pb TIMS U/Pb TIMS U/Pb

504 504 529 527 521 558 1061 528 1049 516 522 557 544 555 1084 616 632 554 629 639 580 809 656 676 580 613 1044 920 589 1146

4 6 4 50 3 6 56 10 19 5 10 11 15 11 8 52 8 16 5 4 5 6 2 12 5 4 25 5 9 3

Paulsson & Austrheim (2003) Paulsson & Austrheim (2003) Bisnath et al. (2006) Bisnath et al. (2006) Jacobs et al. (2003a) Jacobs et al. (2003a) Jacobs et al. (2003b) Jacobs et al. (2003b) Jacobs et al. (2003b) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Jacobs et al. (1998) Ravikant et al. (2004) Ravikant et al. (2004) Ravikant et al. (2004) Ravikant et al. (2008) Ravikant et al. (2008) Ravikant et al. (2008) Ravikant et al. (2008) Ravikant et al. (2008) Ravikant et al. (2008) Ravikant et al. (2008) Ravikant et al. (2008) Ravikant et al. (2008) Ravikant et al. (2008) Ravikant et al. (2008) Ravikant et al. (2008)

significant concentrations toward the NE, SE and NW. This variation either suggests numerous shearrelated lineation-producing events or reflects the folding of earlier unimodal lineations to produce multiple directions. Grantham et al. (1995) recorded bimodal planar structure dipping patterns in NW Sverdrupfjella as well as northern Kirwanveggen (Fig. 7a and b). In contrast, planar fabrics (Fig. 7c and d) in southeastern Sverdrupfjella and southern Kirwanveggen are unimodal and dip dominantly toward the SE (Grantham et al. 1995). Similarly, lineations in northern Sverdrupfjella (Fig. 8b) plunge dominantly eastwards with a subordinate roughly north-plunging group. In northern Kirwanveggen lineations plunge dominantly roughly north and south (Fig. 8c). In southeastern Sverdrupfjella a unimodal lineation direction is recognized plunging dominantly toward the SE (Fig. 8d). The data suggest a zone along the southern coast of northern Mozambique and along the northern coast of DML in which bimodal structural patterns are seen in planar and linear structures. In NW

Sverdrupfjella it is apparent in the field from the wonderful 3D exposures available in Antarctica that the bimodal pattern arises from the refolding of earlier D1 and D2 SE-dipping planar fabrics about near-horizontal NE-oriented D3 fold axes (Fig. 9a and b). The D3 folds are relatively open and commonly verge toward the SE or have top-to-the-SE geometries. In contrast, the D1 and D2 folds are tight to isoclinal and commonly verge toward the NW (Fig. 9a). The folded dyke at Roerkulten reported by Grantham et al. (2006) is also typical of a D3 fold. Other examples of D3 folds with NW-dipping axial planes are seen at the southern end of Brekkerista and the western end of Roerkulten (Grantham 1992).

Metamorphic history In Mozambique the mineral assemblages and grades of metamorphism also show significant differences. Except for the Mugeba and Monapo klippen, the grade of metamorphism in the Nampula Block


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Fig. 6. Contoured stereographic projections of poles to planar structures (a–d) and stereographic projections of contoured lineations (e, f) from various areas mapped in northern Mozambique. The approximate position of the Lurio Belt is shown. The arrows connecting the stereonets to the map area borders show the approximate areas from which the data have been collected.

is typically upper amphibolite facies, with orthopyroxene being seen only rarely as relict grains in magmatic charnockite granitoids or as rare localized diffuse fluid-driven(?) vein charnockitization. The Mocuba Suite gneiss is extensively migmatized, locally showing both Mesoproterozoic and the Pan-African generations of migmatization. The post-Mocuba Suite gneisses preserve only the weakly developed Pan-African migmatization. A significant aspect of the Nampula Block, however, is the abundance of undeformed to weakly deformed granitoids and pegmatites whose ages vary between c. 495 and 530 Ma. Absolute pressure constraints are difficult to constrain for the Nampula Block because of the absence of rock types (meta-pelites, metabasites) with suitable mineral assemblages. The only reliable constraints are provided by sillimanite-bearing migmatitic quartzofeldspathic gneisses implying temperatures of at least c. 700 8C and

pressures ,c. 7–8 kbar. Consequently, no P–T path for the Nampula Block is presented here. However, the weak migmatization and development of granitic melts with ages of c. 495–530 Ma imply an increase in temperature probably to 650– 750 8C during this time period. Implicit in the temperature increase is an increase in depth, because no widespread extensive source of advective heat is recognized. The granitoids are concentrated in the Nampula Block, with only a limited number of small granitoids being recognized north of the Lurio Belt. In contrast, the rocks in the Mugeba and Monapo klippen contain granulite-grade orthogneisses and paragneisses. The orthogneisses are typically ultramafic, mafic to felsic in composition. The ultramafic rocks in the Monapo klippen comprise clinopyroxenites with ,c. 5% modal plagioclase but c. 20% normative plagioclase, implying a

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Fig. 7. Contoured stereographic projections of poles to planar structures from various localities in western and central DML.

substantial omphacitic component. Subtle vermicular intergrowths of Pl þ Cpx (c. 5% Al2O3) (mineral abbreviations after Kretz 1983) are locally developed at the margins of coarse cpx grains with c. 10% Al2O3. These are interpreted as decompression exsolution intergrowths (G. H. Grantham, unpubl. data). The mafic rocks contain Pl–Opx–Cpx–Grt, and some samples have decompression textures defined by garnet with vermicular rims of Cpx/Hbl þ Pl. In the felsic granulites idiomorphic post-tectonic garnet (þ Qtz) after Opx/Cpx define isobaric cooling reactions. The metapelites contain Grt þ Sill þ Pl þ Rt assemblages. Thermobarometry on these assemblages from Mugeba (Roberts et al. 2005) and Monapo (Grantham et al. 2007b) have facilitated the construction of P–T loops (Fig. 10b and c). The P– T loops from Mugeba and Monapo have initial isothermal decompression from c. 900 to 1000 8C and .c. 10 kbar, followed by isobaric cooling at c. 700 8C and c. 6–7 kbar. These P –T loops are comparable with those from Schirmacher Hills (Fig. 10e), Sør Rondane (Fig. 10d) and the

Highland Complex of Sri Lanka (Fig. 10a). All these P–T loops suggest early isothermal decompression followed by isobaric cooling with P– T conditions of c. 6-7 kbar and c. 600–700 8C at c. 550 Ma, although the P–T path from Sør Rondane shows additional complexities not recognized from other areas. The P– T conditions for Schirmacher Hills and Sør Rondane are from Baba et al. (2006) whereas those from Sri Lanka are from Schumacher et al. (1990), Hiroi et al. (1994) and Raase & Schenk (1994). The P–T conditions in Gjelsvikfjella are from Bisnath & Frimmel (2005; Fig. 10f) and show an isothermal decompression path from c. 10 kbar and 700–800 8C during the Mesoproterozoic toward c. 5 kbar and c. 650 8C at 550 Ma. Extensive granitoids of c. 550 Ma age are recognized in Gjelsvikfjella and Mu¨hlig-Hofmannfjella as in the Nampula terrane, also implying a thermal increase at this time. The P–T path described by Grantham et al. (1995) (Fig. 10g) is similar to that described for Gjelsvikfjella and is also characterized by


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Fig. 8. Stereographic projections of lineations from various localities in western Dronning Maud Land.

granitic magmatism at c. 6 kbar and c. 700 8C at c. 490 Ma (Grantham et al. 1991). Progressing southward in Sverdrupfjella to Kirwanveggen, P– T conditions at Neuemayerskarvet in the north of Kirwanveggen (Fig. 10h) described by Grantham et al. (2001) are of the order of c. 6.5 kbar and c. 700 8C. Significantly, c. 500 Ma granitoids are absent in northern Kirwanveggen, and the P–T estimates described by Grantham et al. (2001) were based on a thermally driven dehydration reaction of Hbl þ Pl þ Qtz ! Grt þ Ab þ H2O. Progressing southwards to southern Kirwanveggen, at Drapane, c. 530 Ma sandstones and grits of the Urfjell Group are exposed, implying that at c. 530 Ma the southern Kirwanveggen was exposed at surface. These data imply a crustal depth gradient between north and south Kirwanveggen of c. 6 kbar or c. 20 km. The most important and fundamental difference between these various terranes is that in those areas with Namuno Block age signatures (Mugeba, Monapo, Highland Complex of Sri Lanka, NE Sør Rondane and Schirmacher Hills) the P –T evolution at c. 550 Ma is interpreted as involving significant isobaric cooling. In contrast, rocks with Nampula Block age signatures (Nampula, Sverdrupfjella, northern Kirwanveggen, Gjelsvikfjella and western Mu¨hlig-Hofmannfjella and southwestern Sør Rondane) are largely characterized by extensive magmatism at 500 –550 Ma, implying thermal

heating. This aspect has been recognized by Baba et al. (2008).

Discussion and conclusions This integrated study of geochronological and structural data and metamorphic P–T paths supports an interpretation that rocks north of the Lurio Belt have been thrust southwards over the Nampula (Mozambique)–Maud (Antarctica) block as summarized in Figure 11. Figure 11 represents a schematic cross-section from northern Mozambique to southern Kirwanveggen with the staggered horizontal line representing current exposure levels in Africa and Antarctica superimposed on the inferred 600–500 Ma topographic profile. At the northern end, the Geci Group sediments were deposited at c. 580 Ma (Melezhik et al. 2006; Fig. 2). Progressing southward, metamorphic conditions increase toward the Lurio Belt, with the Namuno Block being thrust over the Nampula Block. In the footwall, depression to greater depths resulted in thermal increase and partial melting, resulting in anatexis and granite genesis in the Nampula Block in Mozambique and in western DML, Antarctica. In this context, the numerical modelling of crustal melting in continental collision zones by England & Thompson (1986) is applicable; they modelled the

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Fig. 9. Field photographs from Jutulrora, Sverdrupfjella, western DML. The upper photograph shows a NW-vergent F1 isoclinal recumbent fold with axial planar (AP) foliation, a NW-vergent F2 isoclinal fold in which the banding is clearly folded and a small SE-vergent F3 concentric fold. The lower photograph shows a larger-scale SE-vergent F3 fold with NW-dipping axial plane. (Note also the SE-dipping thin granitic sheets in the lower photograph.)

P–T evolution in a setting where crustal thickness is doubled by large-scale thrust faulting. Their model predicts that anatexis in the footwall would result in granite genesis c. 40 Ma after the thrust-related thickening, depending on the level of anatexis. This modelling provides a plausible explanation of why the c. 530– 495 Ma granites appear to be mostly undeformed and younger than the metamorphic ages, which start at c. 590 Ma. Following the thrust-related crustal thickening, isostatic rebound with associated inversion and extensional collapse would follow, with concomitant genesis and intrusion of the granites and pegmatites. The extensional structures resulting from the collapse of the orogenic pile would be oriented at c. 908 to the s3 direction of the orogenic compression, as demonstrated in the Himalayas (Dewey 1988). Undeformed pegmatite dykes in the Nampula Block south of the Lurio Belt are correlated with the granites and have strikes

dominantly toward the NNW, implying ENE – WSW extension (Grantham et al. 2007b). The section from northern to southern Kirwanveggen represents a progression from midcrustal levels to the surface, with the deposition of the Urfjell Group at c. 530 Ma. Figure 11 also shows that the c. 550–580 Ma top-to-the-SE deformation was superimposed on an older top-tothe-NW deformation recognized in western DML. The reorientation of structures immediately below the suture zone in the footwall is shown in Figure 11 and is consistent with the structures shown in Figure 9. The interpretation of a large-scale thrust of East African Orogen rocks onto Antarctica has already been proposed by Ravikant et al. (2004, 2008), although their study did not examine the structural and metamorphic details. The interpretation presented in Figure 10 requires different levels of erosion between Africa and Sri Lanka and Antarctica. This difference in level of erosion is supported by the klippen of granulite remnants at Mugeba and Monapo on top of the Nampula Block footwall, as well as similar klippen remnants also preserved at Kataragama in Sri Lanka (Kriegsman, 1995, amongst others) (Fig. 4). Additional klippen along the northern Kalahari Craton margin that have similar geochronology (where data are available) include the Naukluft Mountains in Namibia (Ahrendt et al. 1978; Gray et al. 2006), the Urungwe Klippen in northern Zimbabwe (Shackleton et al. 1966), the Makuti Group (Dirks et al. 1999) and the allochthonous Masoso Suite (Dirks & Jelsma 2006). Within the Zambezi Belt, the main phase of deformation involved transcurrent shearing and SW-vergent thrusting (Hanson et al. 1994; Wilson et al. 1997). In contrast, Antarctica has not been subject to the same level of erosion, resulting in the preservation of a largely continuous slab of East African Orogen rocks from the nunataks at Mramornye c. 718S, c. 88E (Piazolo 2004) via Schirmacher Oasis (Ravikant et al. 2002, 2004, 2008; Baba et al. 2006, 2008), the eastern Mu¨hlig-Hofman mountains, NE Sør Rondane and beyond. We have not explored detailed aspects of the geology east of Sør Rondane other than to note that we are not aware of any rocks, east of SW Sør Rondane, with Nampula Block type lithological, geochronological and metamorphic characteristics. This may suggest that it is possible that rocks of the Nampula Terrane do not extend beyond the SW Sør Rondane. Using the broad geographical grid references in Figure 4, the width of the overthrust block approximates to 4–58 of latitude, equivalent to c. 480– 600 km, representing the distance from the Lurio Belt in Mozambique to the sheared boundary


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Fig. 10. Figure summarizing P– T loops from various localities in Mozambique, Sri Lanka, central DML, western DML and Sør Rondane as shown by the arrows linking the P –T loops with the geographical locality. Abbrevations as in Figure 4.

between the two crustal blocks in Sør Rondane and Mu¨hlig-Hofmannfjella. At the western extremity of the proposed belt, Martin (1974) estimated that the c. 560 –570 Ma (Gray et al. 2006) Naukluft nappes had been transported at least 60 km towards the SE onto the Kalahari Craton. For purposes of comparison, it should be noted that recent studies on the collision zone between

Peninsular India and Asia in the Himalayan Orogen have proposed crustal shortening of c. 550 km (Ratsbacher et al. 1994). If the crustal shortening had begun at c. 590 Ma as suggested by the metamorphic ages in the Nampula Block, at a plate movement rate of c. 4 cm a21, similar to that recorded in current studies of the Himalayas (Ratsbacher et al. 1994), the emplacement of the

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Fig. 11. Schematic cross-section representing the geological relationships from northern Mozambique to southern Kirwanveggen, Antarctica. The cross-section summarizes the deformational structures, the distribution of the 495–530 Ma granitic intrusions in the footwall resulting from burial heating, and the setting for the isothermal decompression followed by isobaric cooling paths of the hanging-wall Namuno Block. The cross-section also shows the relative difference in erosion levels between Africa and Antarctica.

mega-nappe would have required c. 15 Ma to cover the possible 600 km envisaged. The nappe emplacement would then have been followed by inversion, isostatic uplift and erosion, with anatexis at depth c. 40 Ma after emplacement of the obducted slab, as suggested by England & Thompson (1986). The model of England & Thompson (1986) demonstrates that it would require c. 100 Ma for the thermal gradients of the footwall rocks to rise and achieve equilibrium with the hanging-wall rocks and that, depending on depth, it would take c. 40 Ma for rocks at middle crustal level of the footwall to heat into the field of anatexis where melting would be initiated. The model also demonstrates that, depending on the relative depth of melting, one can generate a wide range of granitoids depending on whether the melts are ‘minimum melts’ or are produced by vapour-absent dehydration melting resulting in charnockitic K-rich anhydrous melts. Although there appear to be different erosion levels in Africa and Antarctica, the implications of this model are that a substantial block of crust has been eroded and removed from this belt since c. 550 Ma, particularly in southern Africa. The recognition of reoriented fabrics in Sverdrupfjella and possibly northern Kirwanveggen suggests that these areas were probably in the footwall as well, providing an estimate of the area underlain by the nappe that is significantly larger than that reflected in Figure 4. The recognition of zircon populations with ages typical of the Namuno Block in the Transantarctic Mountains (Goodge 1997; Goodge et al. 2004) and similar rocks in Australia (Veevers et al. 2006) suggests that the sedimentary rocks now exposed in the Transantarctic Mountains and its extensions and the Ellsworth –Whitmore

Mountains (Flowerdew et al. 2007) were probably the depository of the erosion products from the collision of North and South Gondwana along the Damara –Zambezi–Lurio –Sri Lanka–central Dronning Maud axis. Another example of such deposition is in the Urfjell Group, in which the detrital zircon population (Fig. 12) has an age pattern unlike that of its surrounding Nampula Block type floor rocks (data from Croaker 1999) but a distribution more typical of the Namuno Block rocks from north of the Lurio Belt. The overthrust block of tectonic units belonging to the Namuno Block as envisaged in this paper would have placed the detrital source for the Urfjell Group significantly closer to the depository and, in view of the age of the Urfjell Group of c. 530 Ma, the location of the Urfjell Group provides an absolute southern limit of the extent of the overthrust block. This model also provides insights into the geochronological differences between the East African Orogen (Stern 1994; Meert 2003) and the Kungu Orogeny proposed and described by Meert (2003). Stern (1994) initially described the north – south-oriented East African Orogen based on fieldwork in North Africa, through Kenya and Tanzania and suggested that its timing was between c. 900 Ma and c. 550 Ma. The model presented here suggests that, except for the structural outliers in the Nampula Block, Antarctica and Sri Lanka, the East African Orogen is terminated along the Lurio Belt, where it is overprinted and reworked by the Kungu Orogen. The collisional front along the Damara –Zambezi–Lurio –Sri Lanka–central Dronning Maud axis is therefore equivalent to the Kungu Orogen described by Meert (2003). The rapid erosion and uplift of the Nampula Block is recorded in a titanite fission-track study by Daszinnies et al. (2006), who showed that the titanite fission-track ages north of the Lurio Belt

Fig. 12. Histogram and relative probability density curve of zircon ages from the Urfjell Group. Data from Croaker (1999).


are c. 300 Ma and become progressively younger to the coast, to c. 240 Ma. Erosion to surface was completed by c. 180 Ma, the age of the Karooage Angoche Andesite lavas on the northern Mozambique coast (Grantham et al. 2007b). This model explains many of the correlation conundrums that have puzzled scientists in Gondwana reconstructions. It needs to be tested with much additional work. This work should probably involve constraining the precise ages of deformation in the vicinity of the suture zones between the Namuno type blocks and Nampula type blocks in Mozambique, Malawi, Sri Lanka and Antarctica. The extent of the overthrust blocks should also be tested with detailed geochronology and mapping on the Urungwe klippe in northern Zimbabwe, as well as geochronology and P –T work on the nunataks at Mramornye in Dronning Maud Land. Another potential implication is that, after this collision, extensive parts of the northern Kalahari Craton may have been in the footwall of a large nappe. It is an open question as to whether there are other .550 Ma lithological units on the Zimbabwe Craton that may be allochthonous. Possible candidates include the Makuti Groups and the Rushinga Group in Zimbabwe and the Frontier Formation in Mozambique. The origin and distribution of the c. 850 Ma rocks in Mozambique adjacent to those in the vicinity of the NE corner of Zimbabwe need to be studied in detail. Geochronology focusing on low-temperature closure systems in Zimbabwe may provide some idea of how far south the nappe travelled over Zimbabwe, if at all. Similarly, the grits, sandstones, shales and conglomerates of the pre-Karoo-age Sijarira Group in northwestern Zimbabwe may represent the Cambrian-age detritus eroded from the overthrust Zambezi Belt rocks and, if so, then their current positions constrain the southern limit of the mega-nappe (P. Dirks pers. comm.;), just as the Urfjell Group do in southern Kirwanveggen, western DML. A sedimentological study combined with a zircon provenance study of the Sijarira Group would be particularly illuminating. The uncertainty of the position of the southwestern end of the Lurio Belt in central northern Mozambique and southern Malawi also needs further fieldwork and geochronology. We are sure that aspects of this model will be require revision and modification as new data are produced from further studies in Mozambique, Antarctica and Sri Lanka. New comprehensive geochronological and petrological data from Sri Lanka are vitally important. The model also needs to be tested with palaeomagnetic studies, particularly in view of the ‘destruction’, by crustal duplication followed by erosion, of large crustal blocks from the Namuno Block potentially involving 58 of latitude or 600 km.

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We would like to acknowledge all the people who have over the years assisted in the field in various capacities, and who have assisted with data and sample collection in Antarctica and Mozambique. In Antarctica these include geologists R. Thomas and B. Groenewald, and in Mozambique these include R. Thomas, G. de Kock, M. du Toit, P. Botha, M. Kota, R. Opperman, M. Rohwer, J. C. Nolte, M. Cronwright, I. Haddon, J. Miller, S. De Azevedo, S. Fernando, R. Matola, G. Cune and S. Kagashima. We would also like to acknowledge the assistance of H. Kaidan and D. Dunkley with SHRIMP analyses at NIPR, Tokyo. Critical reviews by R. Hansen and S. Johnson significantly improved the manuscript. Discussions with P. Dirks provided valuable insights into aspects related to the Zimbabwe Craton cover. Correspondence with C. Reeves led to the recognition that the Lurio Belt shear zone could also have a horizontal attitude. Permission to publish the summarized geochronological data from Mozambique was also granted by E. Daudi, Director of Direcaõ Nacional Geologia, Maputo, Mozambique. This paper is dedicated to the pioneers R. Sacchi and P. Pinna, who probably would have developed this model had they had the access to extensive single zircon geochronology.

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An overview of geological studies of JARE in the Napier Complex, Enderby Land, East Antarctica HIDEO ISHIZUKA Department of Geology, Kochi University, Kochi 780-8520, Japan (e-mail: [email protected]) Abstract: Subsequent to the reconnaissance fieldwork in 1982, the Japanese Antarctic Research Expedition (JARE) carried out extensive geological studies that focused on structural and tectonic aspects, petrology, geochemistry and geochronology of the Napier Complex in Enderby Land, East Antarctica. Detailed field investigations in several key areas, including geological mapping of the Mt. Riiser-Larsen area and Tonagh Island, revealed that the Napier Complex comprises layered and massive gneiss units, of which the layered unit is composed of garnet felsic gneiss, orthopyroxene felsic gneiss, pelitic and basic gneisses, impure quartzite, and minor metamorphosed banded iron formation, whereas the massive unit consists mainly of orthopyroxene felsic gneiss. The boundary between the units is transitional in the Mt. Riiser-Larsen area, in which metamorphosed anorthosite and ultramafic rocks occur as thin layers, or blocks or pods, but on Tonagh Island the boundary is closely associated with the shear zone. Nine deformation episodes (D1 –D9) were suggested for Tonagh Island. These results of fieldwork were presented in detail in two geological maps. Geochemical studies showed that (1) garnet–sillimanite gneisses and garnet-rich felsic gneisses were derived from mudstone and sandstone, respectively, both enriched in MgO, Cr and Ni; (2) orthopyroxene felsic gneisses have a close REE affinity with Archaean tonalite–trondhjemite–granodiorite (TTG); (3) basic gneisses were derived from light rare earth element (LREE)-enriched or -depleted basalts; (4) meta-ultramafic rocks are comparable with komatiite and related depleted mantle peridotite. This suite of protoliths is reminiscent of Archaean greenstone–granite belts. Precise analyses of physical conditions of metamorphism were carried out by using reliable approaches such as feldspar thermometry, alumina content of orthopyroxene, inverted pigeonite and bulk-rock compositions, and clinoand orthopyroxene compositions with different textures (porphyroblastic and neoblastic), and the results suggested that the maximum metamorphic temperature might have reached 1130 8C (i.e. ultrahigh-temperature (UHT) metamorphism). P– T evolution of the Napier UHT metamorphism was examined by analyses of reaction textures combined with fluid inclusion studies, suggesting both clockwise (Bunt Island) and counterclockwise (Mt. Riiser-Larsen and Tonagh Island) P– T– t paths. U–Pb sensitive high-resolution ion microprobe and secondary ionization mass spectrometry zircon ages from the Mt. Riiser-Larsen area and Tonagh Island indicate three stages of protolith formation at around 3.28 –3.23, 3.07 and 2.68–2.63 Ga, and two contrasting ages for the timing of peak UHT metamorphism at either c. 2.55 or c. 2.51– 2.45 Ga. On the basis of these results, more comprehensive studies on the Napier Complex are essential in the future for understanding (1) the role and age of TTG protolith and (2) the origin and timing of UHT metamorphism.

In Enderby Land, East Antarctica, the Napier Complex is situated in the northern part and the Rayner Complex in the southern part (Fig. 1). After a reconnaissance investigation by Soviet pioneers (Kamenev 1975; Ravich & Kamenev 1975), regular fieldwork on the Napier Complex was carried out by Australian expeditions during 1970–1980. The results have been summarized by Grew & Manton (1979), Sheraton et al. (1980, 1987), Grew (1982a) and Harley & Hensen (1990). These studies have revealed that (1) the Napier Complex underwent extremely hightemperature metamorphism (.900 8C), classified as ultrahigh-temperature (UHT) metamorphism (e.g. Spear 1993), during the late Archaean era,

and (2) a part of the protoliths of the Napier Complex shows a very old radiometric age, going back to early Archaean time, and thus belongs to one of the oldest continental crusts in the world. It is, therefore, most likely that the Napier Complex offers an excellent opportunity to study the processes of continental birth and subsequent growth during the Archaean to Proterozoic era. With the aim of further understanding such processes, the 23rd Japanese Antarctic Research Expedition (JARE-23) started geological investigations in the Napier Complex in 1982, and this work was followed by several expeditions in later years (Table 1). In particular, during 1996–1999 full summer field studies were carried out in the


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Fig. 1. Locality map of the Napier and Rayner Complexes, Enderby Land, East Antarctica.

Napier Complex (JARE-38, 39, and 40) under the auspices of a multidisciplinary project of the National Institute of Polar Research (NIPR), Japan, to study the Structure and Evolution of East Antarctic Lithosphere (‘SEAL’ project). In this study, an overview of important results of JARE studies on the Napier Complex obtained mainly as part of the SEAL project is provided, along with a brief consideration of future studies.

Lithologies and geological structures On the basis of field studies in the Napier Complex supported by JARE-38 (Ishizuka et al. 1997a, b), -39 (Moriwaki 1998) and -40 (Motoyoshi et al. 1999), detailed geological maps of the Mt. Riiser-Larsen

area (Ishizuka et al. 1998; Ishikawa et al. 2000) and Tonagh Island (Osanai et al. 1999, 2001a) were published (Fig. 2). These geological maps are accompanied by explanatory text that describes the geological structure and lithology of the area, and includes preliminary data on petrology, geochemistry and geochronology, which are briefly outlined below. The Mt. Riiser-Larsen area is underlain by UHT metamorphic rocks and minor dolerite dykes; the dykes are provisionally termed ‘Amundsen dykes’ after Sheraton et al. (1987), and were apparently emplaced after the UHT metamorphism (Ishizuka et al. 1998; Ishikawa et al. 2000). Suzuki et al. (2008) studied the Amundsen dykes in this area in detail, and revealed that: (1) there are NE–SWand north– south-striking dykes, of which the

Table 1. Localities where JARE parties have surveyed, and localities where icebreaker Shirase helicopters reconnaissance flights have landed in the Napier and Rayner Complexes JARE

Date

23 29 31

1982.2 1988.2 1990.2

34

1993.2

35 36 37 38 39

1994.2 1995.2 1996.2 1996.12 –1997.2 1998.1 –1998.2

40

1998.12 –1999.1

41 42 46

2000.2 2000.12 –2001.2 2005.2

Research areas Mt. Riiser-Larsen Mt. Riiser-Larsen Mt. Riiser-Larsen, Edward Is., Mt. Oldfield, Tonagh Is., Mt. Pardoe, Hydrographer Is., McIntyre Is. Hydrographer Is., McIntyre Is., Mt. Riiser-Larsen, Forefinger Point, Raggatt Mts, Dick Peaks,* Mt. Humble,* Mt. Maslen* Mt. Riiser-Larsen Mt. Riiser-Larsen Mt. Riiser-Larsen Mt. Riiser-Larsen, Tonagh Is. Tonagh Is., Bunt Is., Priestley Peak, Beaver Is., Bowl Is., Mt. Trail, Mt. Tod, Mt. Riiser-Larsen Mt. Riiser-Larsen, Tonagh Is., Mt. Pardoe, Howard Hills, Edward Is., Christmas Point Tonagh Is. Mt. Riiser-Larsen, Tonagh Is. Condon Hills, Mt. Lira, Mt. Yuzhnaya, Mt. Bergin, Forefinger Point, Fyfe Hills, Mt. Cronus

*Areas where icebreaker Shirase helicopters landed.

JARE STUDIES IN THE NAPIER COMPLEX

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Fig. 2. Locality map of the northwestern part of the Napier Complex.

north–south-striking dykes interrupt the NE –SWstriking ones; (2) these dykes are geochemically grouped into tholeiite basalt (THB), highmagnesian andesite (HMA), alkaline basalt (AL), and enriched mid-ocean ridge basalt (E-MORB)like rock (THB-m), of which the THB and HMA belong to the NE –SW-striking dykes, and the AL and THA-m to the north–south-striking varieties; (3) the Sm–Nd and Rb –Sr bulk-rock isotope ratios of the THB dykes define an isochron ages of 2.0–1.9 Ga, whereas the AL dykes yield an isochron age of 1.2 Ga in the Rb –Sr system. Significantly, the AL and THB-m dykes are very rare in other areas of the Napier Complex (Sheraton et al. 1987), and the emplacement age of 2.0–1.9 Ga is the first report for the Amundsen dykes from the Napier Complex (Sheraton & Black 1981; Sheraton et al. 1987). Somewhat surprisingly, however, the 2.0–1.9 Ga period coincides with a global peak in mantle-derived magmatism (Condie 1997). Suzuki et al. (2008) consequently suggested that: (1) the north–south-striking dykes may occur in restricted areas in the Napier Complex, whereas the

NE–SW-striking dykes are regional; (2) the 2.0–1.9 Ga magmatism of the NE–SW-striking dykes may have been related to the formation of continental crust of the Rayner Complex. Furthermore, on the basis of geochemical affinities of the THB and HMA dykes, such as large ion lithophile element (LILE) and light rare earth element (LREE) enrichment and negative anomalies of Nb, Ti and/or P in a spider diagram normalized to primitive mantle, Suzuki et al. (2008) demonstrated that the origin of these dykes is analoguous to modern subduction-related arc volcanism. The UHT metamorphic rocks in the Mt. RiiserLarsen area commonly display metamorphic foliation that strikes NE –SW to east– west, and dips at a moderate to gentle angle (20 –408) to the south or SE, and they can be divided into layered and massive gneiss units (Ishizuka et al. 1998; Ishikawa et al. 2000). The layered gneiss unit, occurring in the central to northwestern part of the area, is characterized by layering composed of garnet felsic gneiss with subordinate amounts of orthopyroxene felsic gneiss, pelitic and basic

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gneisses, impure quartzite, and metamorphosed banded iron formation. The massive gneiss unit, developing in the southern to southeastern part of the area, consists mainly of orthopyroxene felsic gneiss, in which the layering is not conspicuous and the lithology is rather monotonous. As noted by Sheraton et al. (1987), this essentially follows the Soviet geologists’ subdivision of the Napier Complex into two major gneiss groups: layered garnet –quartz –feldspar gneiss with subordinate pelitic, psammitic, and ferruginous metasediments, and massive pyroxene –quartz –feldspar gneiss with minor mafic granulite (Kamenev 1975). However, these two gneiss units are not mutually exclusive. Indeed, transitional varieties occur between them, in which metamorphosed anorthosite and metaultramafic rocks (clinopyroxenite, orthopyroxenite and peridotite) occur characteristically as thin layers, or blocks or pods. The most prominent geological structure in the Mt. Riiser-Larsen area could be denoted by shear zones, in which the metamorphic rocks have been sheared to mylonites or sometimes pseudotachylites. These shear zones are mostly near-vertical and tend to follow pre-existing structures such as dyke margins. The width of the shear zones ranges from several centimetres to a few metres, but in the western part of the area a shear zone with the maximum width of 400 m occurs, striking north to south; Ishizuka et al. (1998) and Ishikawa et al. (2000) called this shear zone the Riiser-Larsen Main Shear Zone (RLMSZ). The lithology and geological structure are discontinuous from the western to the eastern parts of the RLMSZ. Ishikawa et al. (2000) found that two phases of folding are present in both parts of RLMSZ: the first is a fold with the axes aligned NNE–SSW and wavelengths up to 100 m, and the second is a fold that forms a broad dome structure after the UHT conditions. On Tonagh Island the metamorphic rocks are divided into five lithological units (Units I –V from north to south) based on their lithologies and geological structures (Osanai et al. 1999, 2001a). Thrust-shear zones accompanied by remarkable mylonite and later pseudotachylite–cataclasite bound each unit. Of these units, Unit I consists mainly of layered gneisses similar to the layered gneiss unit of the Mt. Riiser-Larsen area, whereas Units II, III and IV comprise two-pyroxene gneiss and garnet –orthopyroxene gneiss with minor layered gneisses. Unit V is mainly composed of orthopyroxene- and garnet-bearing quartzofeldspathic gneisses with subordinate layered gneisses. This lithological contrast between Unit V and other units led Osanai et al. (1999, 2001a) to suggest that the most prominent tectonic boundary on Tonagh Island may be the shear zone between Units V and IV. Subordinate amounts of

unmetamorphosed alkali-dolerite and granitic pegmatite cut across the sequence of metamorphic rocks, and the alkali-dolerite is also discordant to the unit boundary shear zone. Toyoshima et al. (1999) analysed the geological structure of Tonagh Island, and divided the deformation history into D1 –D9; the D1 structure would have been formed under non- or weakly deformational conditions during the thermal peak of prograde metamorphism, the D2 –D6 structures would have been produced under retrograde granulite-facies conditions, and subsequently the D7 –D9 brittle faulting modified the structures in part. For the whole Napier Complex, Sheraton et al. (1987) previously proposed three tectonothermal episodes; the first episode (D1 –M1), characterized by the formation of the present foliation, intrafolial folds and lineation, was synchronous with the peak of the granulite-facies metamorphism; the second episode (D2 –M2) produced tight to isoclinal, commonly asymmetric folds under granulite-facies conditions; and the third episode (D3 –M3) formed major asymmetric folding during the waning stage of high-grade metamorphism. Toyoshima et al. (1999) interpreted their D1, D2 –D4, and D5 – D6 stages as corresponding to the first (D1 –M1), second (D2 –M2) and third (D3 –M3) episodes, respectively. Toyoshima et al. (2001) further suggested multiple stages of pseudotachylite formation, related to D3, D6 and D8 deformations. Thus, detailed geological information for two key areas in the Napier Complex, the Mt. Riiser-Larsen area and Tonagh Island, were obtained during the JARE field studies, which will aid our understanding of the processes of Archaean crustal formation and UHT metamorphism. However, the regional geological map of Sheraton et al. (1987) is the only available regional geotectonic information on the Napier Complex, and there is still a lack of detailed geological information for other regions in Enderby Land. In this regard, the structural data for the Napier Complex as shown by Toyoshima et al. (2008, Figs 1 and 5) show that: (1) the Napier Complex can be subdivided into several units or blocks separated by east –west, NE –SW-, and NW– SE-striking faults including the RLMSZ; (2) the general strike changes from east –west in the western part to NNE– SSW or NE–SE in the central to eastern part; (3) several folds trending east –west and NE– SW are developed; (4) a dome-and-basin fold pattern on a regional scale is characteristic in some areas; (5) these structural subdivisions may be closely related to P–T –t evolution such as a clockwise or counterclockwise P–T –t path. These structural features are constructed from attitude data for foliations shown in the geological maps of Sheraton et al. (1987), Ishikawa et al. (2000),


and Osanai et al. (2001a), which will be useful for deducing a regional tectonic framework of the Napier Complex in the future.

Protoliths and their formation ages Geochemical studies on protoliths of the Napier UHT metamorphic rocks in the Mt. Riiser-Larsen area by Suzuki et al. (1999) revealed that: (1) garnet –sillimanite gneisses and garnet-rich felsic gneisses are of sedimentary origin, such as mudstone and sandstone, respectively, which are characteristically enriched in MgO (c. 4.1 wt%), Cr (c. 680 ppm) and Ni (c. 130 ppm) in the garnet –sillimanite gneisses, and in MgO (c. 2.7 wt%), Cr (c. 170 ppm) and Ni (c. 70 ppm) in the garnet-rich felsic gneisses; (2) orthopyroxene felsic gneisses and garnet-poor felsic gneisses are chemically comparable with CIPW normative tonalite to granodiorite and granite, respectively, and have a close REE affinity with Archaean tonalite –trondhjemite– granodiorite (TTG) (Luais & Hawkesworth 1994); (3) mafic gneisses are divided into a quartz-free and LREE-depleted type and a quartz-bearing and LREE-enriched type, suggesting a different source material for these two types; (4) phlogopite-free and phlogopitebearing meta-ultramafic rocks were derived from depleted mantle peridotites and komatiitic rocks, respectively, and the latter exhibit a magmatic differentiation controlled by olivine fractionation. The sedimentary precursors of these rocks (mudstone and sandstone enriched in MgO, Cr and Ni) are very similar to the sedimentary rocks reported from Archaean terranes (Condie 1997) such as the Kaapvaal Craton in southern Africa (Condie & Wronkiewicz 1990), where enrichment in these components appears to reflect the existence of komatiite –high-Mg basalt sources. For Tonagh Island, Owada et al. (1999, 2000) also studied the geochemistry of mafic gneisses and meta-ultramafic rocks (pyroxenite, websteritic peridotite and hornblende-bearing lherzolitic peridotite), and showed that (1) a majority of mafic gneisses are tholeiitic basalts in composition, and (2) some of the mafic gneisses and meta-ultramafic rocks are enriched in MgO (up to 31 wt%) and LREE, and resemble komatiitic basalts to komatiites. Of particular interest in the Riiser-Larsen area and on Tonagh Island is the presence of the TTG –komatiite association that is commonly reported from other Archaean greenstone –granite belts (e.g. Condie 1994). Sheraton et al. (1987) previously found that the less-fractionated high-Mg metamorphosed dykes are characterized by high MgO, Cr and Ni, but rather low TiO2, Na2O, P2O5, Zr, Nb and Y, and suggested that in these

125

respects these dykes have some chemical affinities with basaltic komatiites. Also, Sheraton et al. (1987) described the precursors of felsic gneisses throughout the Napier Complex, and suggested that they are comparable with the TTG protolith. It follows that, although the rocks were metamorphosed under UHT conditions, the protolith of the Napier Complex has an important role in understanding the continental evolution. On the other hand, the two types (LREE-enriched and -depleted types) of mafic gneisses described by Suzuki et al. (1999) and Owada et al. (1999, 2000) were reported by Sheraton et al. (1987) from other areas of the Napier Complex. This means that there are two types of source materials for the mafic gneisses throughout the Napier Complex, which is also important for understanding the origin and evolution of the Napier continental crust. Ages of protoliths were determined by means of the Sm–Nd bulk-rock method, U –Pb sensitive high-resolution ion microprobe (SHRIMP) and secondary ionization mass spectrometry (SIMS) mineral isotope methods, and the chemical U –Th –Pb isochron method (CHIME) (Table 2). The results show four age clusters at c. 3.7, 3.3, 3.0 and 2.6 Ga. Of these, the U –Pb SHRIMP and SIMS zircon ages are restricted to 3.28 –3.23 Ga (Shiraishi et al. 1997; Hokada et al. 2003), 3.07 Ga (Hokada et al. 2003) and 2.68 –2.63 Ga (Carson et al. 2002; Crowe et al. 2002). These U –Pb data are usually obtained from zircons with oscillatory zoning pattern and/or magmatic Th/U ratio, and are interpreted as the ages of protoliths. In the literature, apart from JARE studies, there are many geochronological studies focusing on the ages of protoliths in the Napier Complex. After reports of Pb –Pb ages of c. 4.0 Ga for orthogneisses from Fyfe Hills (Ravich & Kamenev 1975; Sobotovich et al. 1976), the subsequent studies by using U– Pb SHRIMP zircon dating indicated tonalitic igneous events at 4.0–3.8 Ga from Fyfe Hills, Mt. Sones and Gage Ridge (Black et al. 1986a; Harley & Black 1997; Kelly & Harley 2005). These very old ages have not been detected during JARE studies, although Owada et al. (1994) and Asami et al. (1998) reported an Sm–Nd bulk-rock age of 3.71 Ga for mafic gneiss from Tonagh Island and a CHIME zircon age of 3.65 Ga for felsic gneiss from Mt. Cronus, respectively. In contrast the U –Pb SHRIMP zircon ages of 3.28 –3.23 Ga from Tonagh Island (Shiraishi et al. 1997) and 3.27 Ga from Mt. Riiser-Larsen (Hokada et al. 2003) have not been found in other areas, although other radiometric data such as conventional zircon data, Rb–Sr bulk-rock data and Sm–Nd bulk-rock data showed the ages of protoliths to be 3.2 –3.0 Ga from Proclamation Island (Black et al. 1986b; Sheraton & Black

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Table 2. Age data estimated as protolith ages of the Napier Complex measured by the JARE geology group Sample locality

Method

Lithology

Age (Ga)

Tonagh Island

Sm –Nd bulk-rock

Tonagh Island

U –Pb (SHRIMP) mineral (core of zircon) U –Th–Pb (CHIME) mineral (core of zircon) Sm –Nd bulk-rock

Mafic gneiss Felsic gneiss Felsic gneiss

3.71 2.46 3.28– 3.23

Shiraishi et al. (1997)

Felsic gneiss

3.65

Asami et al. (1998)

Mafic gneiss Felsic gneiss Mafic gneiss Felsic gneiss

2.92 3.02 2.63 2.63

Suzuki (2000) Carson et al. (2002)

Felsic mylonite

2.68

Crowe et al. (2002)

Felsic gneiss

3.27

Hokada et al. (2003)

Felsic gneiss

3.07

Mt. Cronus Mt. Riiser-Larsen Tonagh Island Tonagh Island Mt. Riiser-Larsen

U –Pb (SIMS) mineral (core of zircon*) U –Pb (SHRIMP) mineral (core –margin of zircon*) U –Pb (SHRIMP) mineral (core of zircon*)

Reference Owada et al. (1994)

*Zircon with oscillatory zoning.

1983) and Fyfe Hills (Black et al. 1983, 1984). U– Pb SHRIMP zircon ages of 2.98–2.92 Ga, similar to that (3.07 Ga) from Mt. Riiser-Larsen (Hokada et al. 2003), have been also reported from Proclamation Island and Dallwitz Nunatak (Harley & Black 1997; Kelly & Harley 2005). As suggested by Hokada et al. (2003), there are at least three magmatic events in the Napier Complex, at c. 3.8 Ga in Fyfe Hills, Mt. Sones and Gage Ridge, 3.3 Ga in Tonagh Island and Mt. Riiser-Larsen and 3.0 Ga in Proclamation Island, Dallwitz Nunatak and Mt. Riiser-Larsen. In addition, it is most likely that the U–Pb SHRIMP zircon ages of about 2.63 and 2.68 Ga of felsic gneisses reported from Tonagh Island (Carson et al. 2002; Crowe et al. 2002) are the youngest ages of protoliths in the Napier Complex. It is, therefore, stressed that the JARE studies add new evidence of U–Pb SHRIMP zircon ages (3.3 and 2.6 Ga) to the field of magmatic activities of the Napier Complex. Of these, the 3.3 Ga event has been, however, reported as the age of protoliths from the Archaean terranes in eastern Australia (Kemp et al. 2006) and South Africa (Shirey et al. 2003).

UHT metamorphism and its peak age After the formation of protoliths, the Napier Complex underwent UHT metamorphism characterized by the presence of such diagnostic minerals and mineral assemblages as sapphirine þ quartz (Dallwitz 1968; Ellis 1980; Grew 1980; Harley & Hensen 1990), osumilite (Ellis 1980; Grew 1982b), inverted pigeonite (Sandiford & Powell 1986a; Harley 1987), and sillimanite þ orthopyroxene þ quartz (Harley 1985; Sheraton

et al. 1987). On the basis of mineral assemblages and conventional geothermobarometry, the highest metamorphic temperature (950 –1020 8C) was restricted to the northwestern area of the Napier Complex, that is, the area around Amundsen Bay, whereas the pressure conditions increased from the northern area (c. 5 kbar) to the southern area (11 kbar) of the Napier Complex (e.g. Harley & Hensen 1990; Harley 1998). Also, it has been proposed that the peak conditions of the Napier UHT metamorphism were followed by a period of near-isobaric cooling, for which the pressure was low in the north (,8 kbar) and high in the south (9–10 kbar) (Ellis 1980; Harley & Hensen 1990; Harley 1998). This gradation of pressures is suggested by regional differences in core Al2O3 contents of orthopyroxene equilibrated with garnet þ sillimanite (Harley 1985), differences in reaction coronas on sapphirine þ quartz such as cordierite þ sillimanite þ garnet in the north and orthopyroxene þ sillimanite þ garnet in the south (Sheraton et al. 1987), and the stability of osumilite þ garnet in the area to the north of Amundsen Bay compared with orthopyroxene þ sillimanite þ K-feldspar þ quartz southwards and eastwards (Hensen & Motoyoshi 1992; Harley 1998). These characteristics of UHT metamorphism in the Napier Complex were also reported by the JARE reconnaissance fieldwork in the Mt. Riiser-Larsen area (Motoyoshi & Matsueda 1984, 1987; Makimoto et al. 1989; Motoyoshi & Hensen 1989; Motoyoshi et al. 1990), and reviewed by Motoyoshi (1998). To clarify the more detailed metamorphic characteristics of the Napier Complex, the mineral assemblages, mineral chemistries, and geological structure in the Mt. Riiser-Larsen area (Ishizuka


et al. 1998; Hokada 1999; Ishikawa et al. 2000), Tonagh Island (Hokada et al. 1999; Osanai et al. 1999, 2001a; Owada et al. 1999; Toyoshima et al. 1999; Tsunogae et al. 1999), Howard Hills (Yoshimura et al. 2000; Miyamoto et al. 2004), and Christmas Point (Yoshimura et al. 2001) were studied in detail. On the basis of these descriptive studies, especially at the Mt. Riiser-Larsen area and Tonagh Island, several new approaches were applied for the determination of metamorphic temperature (Hokada 1999, 2001; Ishizuka et al. 1999, 2002; Harley & Motoyoshi 2000; Hokada & Suzuki 2006). Hokada (1999, 2001) calculated chemical compositions of reintegrated perthitic, mesoperthitic and antiperthitic feldspar minerals from felsic gneiss and garnet – orthopyroxene– sapphirine gneiss by the modal proportions and the chemical analyses of host and lamellar domains formed through exsolution, and estimated minimum equilibrium temperatures of 1000 – 1100 8C by using ternary feldspar solvus models. Similarly, Hokada & Suzuki (2006) recovered pre-exsolution single-phase compositions separately for the feldspar core and whole feldspar grain from felsic gneisses with tonalitic and granodioritic compositions, and calculated a temperature range of 940 –1100 8C for the feldspar core and 850 –1070 8C for the whole feldspar grain. Harley & Motoyoshi (2000) used the alumina content of porphyroblastic orthopyroxene (12.2 + 0.5wt% Al2O3) from the sapphirine þ orthopyroxene þ quartz granulite and calculations in the MAS and FMAS systems with theoretical considerations, and deduced a temperature in excess of 1120 8C. Ishizuka et al. (1999) analysed compositions of inverted pigeonite (exsolution intergrowth of clino- and orthopyroxenes) and bulk rocks from quartz –magnetite rocks, and, using ternary pyroxene phase diagrams, estimated a temperature of 1130 8C, at which one phase pyroxene (pigeonite) was once stable. Ishizuka et al. (2002) determined clino- and orthopyroxene compositions with different textural features (porphyroblastic and neoblastic) from meta-ultramafic rock, and used pyroxene thermometry to calculate a temperature that ranged up to 1130 8C for the porphyroblastic pyroxenes. The results of these new approaches revealed that the maximum temperature of the Napier UHT metamorphism reached 1130 8C at the Riiser-Larsen area and Tonagh Island. To further characterize the Napier UHT metamorphism, Hensen & Osanai (1994) and Motoyoshi & Hensen (2001) have carried out analyses of fluorine in mica minerals and related experiments. Hensen & Osanai (1994) found that synthetic F-rich phlogopite (Mg/(Fe þ Mg) ¼ 0.75) with F/(F þ OH) ¼ 0.6 (equivalent to 5–6 wt% F) disappeared at 1045 8C and 9 kbar, and indicated that F-rich phlogopite can be a stable phase in the

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sapphirine þ quartz stability field in appropriate bulk compositions by the fractionation of available F into biotite during partial dehydration melting. This experimental result was supported by Motoyoshi & Hensen (2001), who analysed high fluorine contents (up to 8.2 wt% F) of phlogopites of pelitic granulite and quartz-free garnet-granulite from the Mt. Riiser-Larsen area and documented that very fluorine-rich phlogopite (6–8 wt% F) can be stable in UHT mineral assemblages with aluminous orthopyroxene, osumilite, pyrope-rich garnet, and sapphirine þ quartz. Similarly, Tsunogae et al. (2000) analysed high fluorine (up to 2.6 wt% F) and chlorine (up to 1.45 wt% Cl) contents of amphibole minerals in mafic granulites from Tonagh Island, and showed that amphibole minerals containing halogen elements could also be stable at UHT metamorphic temperatures. Subsequently, those workers considered that the H2O component in mica and amphibole minerals was replaced by a fluorine component at a temperature that is usually sufficient to melt mica and/or amphibole minerals in ordinary continental crust. These studies strongly demonstrate that such replacement may have played an important role in the very small degree of crustal melting during UHT metamorphism in the Napier Complex. For the P–T evolution of the Napier Complex, Motoyoshi & Hensen (1989) suggested a counterclockwise P– T path based on the detailed petrography of sapphirine þ quartz þ orthopyroxene symplectite, probably after cordierite, from the Mt. Riiser-Larsen area, whereas Osanai et al. (2001b) described sapphirine-consuming and osumilite-producing reactions from Bunt Island and suggested isothermal decompression in part of the clockwise P–T path. Tsunogae et al. (2001, 2002) also suggested that the CO2 isochores estimated by fluid inclusions in sapphirine and quartz from Tonagh Island intersect the counterclockwise P –T trajectory at around 6– 9 kbar at 1100 8C, whereas Tsunogae et al. (2003) described two types of CO2-rich fluid inclusions in garnet from Bunt Island (high-density and early entrapped inclusion, and low-density and late entrapped inclusion) and proposed a clockwise P–T evolution from 10 kbar at 1050 8C to ,7 kbar at ,950 8C. From other areas of the Napier Complex, both clockwise (Ellis 1987; Harley 1989) and counterclockwise (Sandiford & Powell 1986a, b; Hollis & Harley 2002) P–T –t paths have been also proposed, as summarized by Harley (2004). There is a wealth of textural descriptions of the post-peak stage as described above, but the prograde history of the Napier UHT metamorphism has been largely obliterated by the high-temperature conditions. This may partly result in the contrasting views as shown above. On the other hand, Hokada (1999) and Hokada et al. (2008) described two

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Table 3. Age data estimated as metamorphic ages of the Napier Complex measured by the JARE geology group Sample locality

Method Sm –Nd bulk-rock

Tonagh Island

U –Pb (SHRIMP) mineral (margin of zircon)

Mt. Riiser-Larsen Mt. Riiser-Larsen

40

Mt. Riiser-Larsen

Ar – 39Ar bulk-rock U –Th–Pb (CHIME) mineral (zircon) U –Th–Pb (CHIME) mineral (monazite) U –Th–Pb (CHIME) mineral (zircon) U –Th–Pb (CHIME) mineral (monazite) U –Th–Pb (CHIME) mineral (zircon) U –Th–Pb (CHIME) mineral (monazite) Sm –Nd bulk-rock and mineral

Tonagh Island

U –Pb (SHRIMP) mineral (zircon) U –Pb (SHRIMP) mineral (monazite) U –Th–Pb (CHIME) mineral (zircon) Sm –Nd bulk-rock and mineral

Mt. Riiser-Larsen Christmas Point Tonagh Island Tonagh Island Mt. Riiser-Larsen

Sm –Nd bulk-rock and mineral U –Th–Pb (CHIME) mineral (xenotime) U –Pb (SIMS) mineral (margin of zircon) U –Pb (SHRIMP) mineral (margin of zircon) U –Pb (SHRIMP) mineral (zircon)

Mt. Riiser-Larsen

Age (Ga)

Reference

Felsic gneiss Mafic gneiss Felsic gneiss Felsic gneiss Felsic gneiss Felsic gneiss Felsic gneiss Felsic gneiss Felsic gneiss Felsite Felsic gneiss Felsic gneiss Mafic gneiss Garnet gneiss Felsic gneiss Felsic gneiss Felsic gneiss Alminous gneiss Felsic gneiss Garnet gneiss Felsic gneiss Pegmatite Felsic gneiss Felsic mylonite Felsic gneiss

2.56 2.60 2.66 2.55– 2.44 2.0 2.41 2.44 2.6– 2.3 2.6– 2.1 2.6– 2.4 2.5– 2.4 2.38 2.38, 2.30 2.36 2.83– 2.80, 2.64 – 2.44 2.46– 2.39 2.9– 2.8, 2.6– 2.3 1.87 1.90 1.56 2.20 2.17 2.47– 2.45 2.55– 2.47 2.52– 2.45

Tainosho et al. (1994, 1997) Shiraishi et al. (1997) Takigami et al. (1998) Asami et al. (1998) Hokada (1999)

Suzuki (2000)

Owada et al. (2001, 2003) Suzuki et al. (2001) Grew et al. (2002) Carson et al. (2002) Crowe et al. (2002) Hokada et al. (2003)

H. ISHIZUKA

Mt. Pardoe

Lithology

Mt. Riiser-Larsen Beaver Island Reference Peak Mt. Riiser-Larsen Mt. Riiser-Larsen

Bunt Island

U –Pb (SHRIMP) mineral (zircon) U –Pb (SHRIMP) mineral (monazite) U –Pb (conventional) mineral (zircon) U –Pb (conventional) mineral (rutile) Rb –Sr mineral (phlogopite) Sm –Nd bulk-rock and mineral Lu –Hf bulk-rock and mineral

Granulite Granulite Quartz gneiss Skarn Felsic gneiss Felsic gneiss Felsic gneiss Paragneiss Paragneiss Psammitic gneiss Mafic gneiss Granitic gneiss Granitic gneiss Granitic gneiss Felsic gneiss Granulite Granulite Granulite Layered gneiss

2.40 2.44 2.42 2.41 2.43 2.42 2.43 2.49– 2.48 2.47– 2.45 2.36 2.38 2.38 2.51– 2.47 2.47 2.44 1.5 1.85 1.85 2.40

Asami et al. (2002)

Hokada et al. (2004) Suzuki et al. (2006)

Miyamoto et al. (2006)

Choi et al. (2006)


Howard Hills

U –Th–Pb (CHIME) mineral (monazite) U –Th–Pb (CHIME) mineral (zircon) U –Th–Pb (CHIME) mineral (zircon) U –Th–Pb (CHIME) mineral (monazite) U –Th–Pb (CHIME) mineral (monazite) U –Th–Pb (CHIME) mineral (monazite) U –Th–Pb (CHIME) mineral (xenotime) U –Pb (SHRIMP) mineral (zircon) U –Th–Pb (CHIME) mineral (zircon) Sm –Nd bulk-rock and mineral

129

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H. ISHIZUKA

types of reaction products after sapphirine þ quartz in the Mt. Riiser-Larsen area such as orthopyroxene þ sillimanite in the western area and orthopyroxene or cordierite in the eastern area of the shear zone (RLMSZ), and interpreted it as showing a slight difference in pressure conditions during the isobaric cooling between the western (higher-P) and eastern (lower-P) areas. This implies that the structural movement responsible for the formation of the shear zone could juxtapose rocks with different path of isobaric cooling; that is, rocks derived from different depths. Experimental studies have been applied to examine the P– T conditions of the Napier UHT metamorphism and related magmatism (Motoyoshi et al. 1993; Sakai & Kawasaki 1997; Kawasaki & Motoyoshi 2000, 2006, 2007; Hokada & Arima 2001; Sato & Kawasaki 2002; Sato et al. 2004, 2006). Motoyoshi et al. (1993) carried out a highpressure experiment in the K2O– MgO–Al2O3 – SiO2 (KMAS) system to determine the stability and phase relation of osumilite, and showed that (1) Mg-osumilite breaks down to enstatite, sillimanite, K-feldspar and quartz at pressures between 11 and 12 kbar over a temperature range of 950– 1100 8C with no appreciable dP/dT slope, and (2) the reactions osumilite ¼ cordierite þ enstatite þ K-feldspar þ quartz and osumilite ¼ sapphirine þ enstatite þ K-feldspar þ quartz limit the osumilite stability at lower and higher temperatures, respectively. Consequently, they suggested that in the Napier Complex an inferred ‘isograd’, marking the disappearance of osumilite, is the trace of a nearly isobaric surface at a palaeo-depth of c. 35 km, beyond which osumilite breaks down to orthopyroxene, sillimanite, K-feldspar and quartz. On the other hand, to study the origin of the (A-type?) granite that occurs as several small bodies or subconcordant veins in relatively lowtemperature regions in the Napier Complex (e.g. Napier Mountain and Tange Promontory; Sheraton et al. 1987), Hokada & Arima (2001) carried out a melting experiment on a mixture of feldspar (antiperthite) þ quartz þ orthopyroxene separated from the felsic gneiss of the Mt. Riiser-Larsen area at 1000– 1150 8C and 10 kbar under dry conditions, and showed that a granitic glass (,10 vol.%) with a chemical composition of A-type granite was detected along boundaries between lamella-free plagioclase and quartz grains in the run at 1150 8C. Therefore, they proposed that some genetic links might exist between UHT metamorphism and A-type granite magmatism in the Napier Complex. The outcome of JARE geological investigations also includes determinations of metamorphic ages of the Napier Complex by means of isotope dating methods for Sm –Nd bulk-rock and minerals,

Lu –Hf bulk-rock and minerals, 40Ar – 39Ar bulkrock, U –Pb SHRIMP and SIMS mineral, and chemical dating by CHIME (Table 3), showing that the data are scattered from 2.9 to 1.6 Ga. Of these, the U– Pb SHRIMP and SIMS mineral data include the ages 2.66 and 2.55– 2.44 Ga (Shiraishi et al. 1997), 2.83–2.80, 2.64–2.44 and 2.46 – 2.39 Ga (Suzuki 2000), 2.47 –2.45 Ga (Carson et al. 2002), 2.55–2.47 Ga (Crowe et al. 2002), 2.52 –2.45 Ga (Hokada et al. 2003), 2.49–2.48 Ga (Hokada et al. 2004) and 2.51 –2.47 and 2.47 Ga (Suzuki et al. 2006). From these data, there are two interpretations on the timing of the UHT Napier metamorphism: c. 2.55 Ga (Crowe et al. 2002) and c. 2.51 –2.45 Ga (Carson et al. 2002; Hokada et al. 2004; Suzuki et al. 2006). Asami et al. (2002) reported CHIME ages of 2.40 – 2.44 Ga and interpreted these ages to be the peak event of UHT metamorphism. In addition to the JARE studies, U –Pb SHIRIMP and SIMS zircon data also suggest two further ages: c. 2.59 – 2.55 Ga (Harley et al. 2001; Kelly & Harley 2005) and c. 2.50 –2.45 Ga (e.g. Grew 1998). Harley & Black (1997) proposed c. 2.84 Ga for the peak age of the Napier UHT metamorphism, but this age was later considered by Kelly & Harley (2005) to relate to a low-P/high-T type of metamorphism and not to the c. 8–11 kbar UHT metamorphic event in the Napier Complex. Younger ages in Table 3 were interpreted as reset ages caused by local thermal effects such as intrusion or pegmatite activities, or by local structural effects such as the development of shear zones.

Discussion As summarized above, the geological studies of JARE have provided additional observations and data, and some new aspects and interpretations for the geology, protolith and metamorphism of the Napier Complex. However, there are still several topics of considerable debate, such as the role and age of TTG protolith, and the origin and timing of UHT metamorphism, which are briefly considered below.

Role and age of TTG protolith The TTG protolith as reported from the Mt. RiiserLarsen area (Suzuki et al. 1999) was also described from other areas of the Napier Complex (Sheraton & Black 1983; Sheraton et al. 1987). The origin of the modern TTG suite has been explained as the product of magmatic processes, namely melting of basaltic source rocks (low-Mg amphibolites) principally occurring in a subduction zone (e.g. Foley et al. 2002). A recent study of seismic


experiments demonstrated that TTG varieties (tonalitic rocks), derived from the anatexis of the differentiated basaltic lower crust (Tatsumi 2000), exist in the middle to lower crust of the Mariana intra-oceanic island arc (Takahashi et al. 2007). This is supported by the occurrence of obducted tonalitic rocks at the northern tip (the Tanzawa Mountain) of the Izu arc where the arc collides with the Japan arc (Kawate & Arima 1998). These facts are indicative of the presence of a ‘modern intra-oceanic island arc-like’ tectonic setting that may have played an important role in the formation of Napier TTG protoliths. However, the Archaean TTG suite is slightly different from the modern variety in its geochemistry (Martin 1994; Luais & Hawkesworth 1994), such as high chondrite-normalized (La/Yb)N ratio and low YbN content, and depleted heavy REE (HREE) in the Archaean TTG suite, which has also been described in the Napier TTG protolith (Sheraton et al. 1987; Suzuki et al. 1999). These geochemical features of the Archaean TTG suite require a source material with residual amphibole and/or garnet, because of the strong partitioning of HREE into these minerals (e.g. Martin 1994). This difference in geochemistry between the Archaean and modern TTG suites is considered to result from the difference of physical conditions for the production of TTG magma, such as a difference in the geothermal gradient at a subduction zone from the high T/P gradient during Archaean time to the low T/P variety at present; that is, at similar depths the high T/P geothermal gradient favours the stability of garnet and amphibole (e.g. Condie 1997). The assumption that the Archaean subducted plate was relatively young (,30 Ma) and warm is a major basis for this model. There is, however, no field evidence for this model, and it is still debated. Therefore, further study on the Napier TTG protolith, especially study of its field occurrence and field relations with surrounding rocks, may provide constraints for understanding the birth of Archaean continental crust. In this respect, the regional field relations need to be investigated in more detail in the Napier Complex. The age of protoliths is another subject requiring further research. As described above, there are at least four clusters of protolith ages (c. 3.8, 3.3, 3.0 and 2.6 Ga) reported from the Napier Complex, all of which were U –Pb SHRIMP or SIMS zircon ages from the TTG or granitic protolith (now represented by felsic gneisses or orthogneisses). These ages possibly indicate the magmatic episodes by which the protoliths were generated. However, they do not always indicate the formation ages of juvenile (mantle-derived) crust; that is, there is a possibility that the protoliths were derived from preexisting materials such as basalts or sediments by

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remelting processes. Distinction between juvenile crust and remelting crust is essential to evaluate the rate of crustal growth during Earth’s history (Condie 1997), and it can be examined by means other than the U –Pb isotope method alone. Recent progress in U –Pb zircon isotope geochemistry combined with oxygen and Lu –Hf isotope systematics analysed using laser-ablation multicollector inductively coupled plasma-mass spectrometry (MC-ICP-MS) shows that (1) the model age of Lu –Hf isotopes is useful to examine the formation age of juvenile crust, and (2) the oxygen isotope can distinguish mantle-derived magma (low d18O, e.g. ,6.5‰) from remelting magma (high d18O, e.g. .6.5‰), because higher d18O values reflect a component of 18O-enriched, supracrustal material in the magma from which the zircon precipitated, which could include recycled sedimentary rocks or hydrothermally altered oceanic crust (Cavosie et al. 2005; Hawkesworth & Kemp 2006; Kemp et al. 2006). It is, therefore, important that the protolith ages of the Napier Complex are re-examined by using in situ analysis of U –Pb, oxygen, and Lu –Hf zircon isotopes, which will provide important information that will help in modelling the origin and evolution of the Napier continental crust.

Origin and timing of UHT metamorphism The origin of UHT metamorphism, such as heat source and related tectonics, has a key role in understanding the process that stabilized the continental crust. In particular, the fact that the Napier Complex showed little sign of melting at temperature up to 1130 8C provides a constraint on the thermal structure of the Archaean continental crust and related behaviour of volatile components, because the temperature at similar depths of the modern lower continental crust is assumed to be 800–900 8C. On the basis of the thermal as well as pressure gradations, and the P –T evolution, several models have been proposed for the origin of the Napier UHT metamorphism, such as an asthenosphere heat advection model related to continental collision and thickening (Ellis 1987; Harley 1989, 2004; Arnold et al. 2001) or a magma intrusion model (Sandiford & Powell 1986b; Motoyoshi & Hensen 1989; Hensen & Motoyoshi 1992) or a combination model (Suzuki et al. 2006). There is, however, little field evidence supporting these speculative models, and, as described above the P –T evolution is still a matter of debate. In this respect, very recently, Soyama & Ishizuka (2007a, b) described a UHT granulite sample from the Mt. Riiser-Larsen area, which is composed of osumilite, sapphirine, garnet, orthopyroxene (Al2O3 9.4– 14.9 wt%), sillimanite, K-feldspar, biotite, spinel, and

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H. ISHIZUKA

quartz. They examined the microstructures using a combination of confocal micro-Laser Raman Spectroscopy (LRS) and electron microprobe analysis (EMPA) to find the presence of kyanite þ orthopyroxene (Al2O3 5.0–7.4 wt%) association occurring as an inclusion in osumilite. This finding is critical as the P– T conditions change from the kyanite þ low-Al orthopyroxene stability field to the osumilite þ sillimanite þ high-Al orthopyroxene stability field (lower-P and higher-T field) during the prograde clockwise P –T path in the Riiser-Larsen area. Furthermore, the fact that this sample was collected from the eastern region of the RLMSZ is consistent with the regional subdivision of the Napier Complex into a clockwise and counterclockwise P–T –t path metamorphic unit as shown by Toyoshima et al. (2008, Fig. 5). It is, therefore, suggested that detailed analyses of P –T conditions and evolution of the UHT metamorphism in the Napier Complex require further detailed microstructural studies using new techniques such as LRS. The precise timing of the UHT Napier metamorphism is also controversial at present, and on the basis of the U – Pb SHRIMP and SIMS zircon ages, there are at least two contrasting ages as described above; c. 2.59 –2.55 Ga and c. 2.50– 2.45 Ga. The use of cathodoluminescence (CL) and backscattered electron (BSE) image analyses, scanning electron microscope (SEM) imaging, and analyses of trace elements such as Th and U have led to attempts that link zoning patterns, internal textures and compositional features of zircon grains in distinguishing igneous or metamorphic origins (Carson et al. 2002; Hokada et al. 2004; Kelly & Harley 2005). This technique has been extensively applied for the dating of protoliths of the Napier Complex, as shown in Table 2. It is, however, not always entirely clear what kinds of zircon textures have been formed during metamorphism; for instance, by growth of new zircon or recrystallization of pre-existing zircon (e.g. Schaltegger et al. 1999; Kelly et al. 2002). Consequently, it is unclear which minerals were in equilibrium with the zircon, and what kinds of metamorphic reactions result in the production of zircon. These problems seem to have largely contributed to the contrasting views on the timing of the UHT metamorphism in the Napier Complex. In this regard, the analyses of zirconium for constituent minerals such as garnet, osumilite, cordierite and sillimanite carried out by Fraser et al. (1997), Pan (1997) and Degeling et al. (2001) are important to obtain a constraint on zircon formation. For example, Degeling et al. (2001) proposed that (1) in the decompression reaction garnet (Zr 24.75 ppm) þ sillimanite (0.08 ppm) þ quartz ¼ cordierite (1.57 ppm), Zr released during

garnet breakdown cannot be incorporated in the cordierite structure, resulting in zircon nucleation and growth, and (2) for the reaction garnet (Zr 62.10 ppm) þ biotite þ sillimanite þ quartz ¼ osumilite (38.71 ppm) þ orthopyroxene (31.80 ppm) þ spinel (0.21 ppm) þ magnetite (0.09 ppm), no new zircon growth takes place. On the other hand, recent studies demonstrate that an analysis of zircon for trace or rare earth elements combined with CL and SEM images, and with U–Pb SHRIMP or SIMS dating is a potentially powerful chemical approach to link zircon growth and recrystallization to metamorphic events (Harley et al. 2001; Hokada & Harley 2004; Kelly & Harley 2005). Hokada & Harley (2004) analysed REE in zircon and coexisting garnet from the UHT gneiss of the Mt. Riiser-Larsen area, and, based on comparisons with recent estimates of equilibrium zircon–garnet REE distribution coefficients (Harley et al. 2001; Rubatto 2002; Whitehouse & Platt 2003), they inferred that the inner core of zircon did not grow with the garnet but grew within a garnet-absent melt that was then injected into the gneiss during the time interval 2.50–2.47 Ga in the waning stages of the Napier UHT metamorphism. Also, Kelly & Harley (2005) used a similar approach for samples from the Gage Ridge, Proclamation Island, Dallwitz Nunatak and Zircon Point, and proposed c. 2.55 Ga as the timing of the Napier UHT metamorphism. Consequently, they suggested that the high proportion of published zircon U–Pb data with ages between c. 2.49 and 2.45 Ga reflects late, post-peak zircon growth and does not date the timing of the peak Napier UHT metamorphism. These new approaches can, therefore, provide important clues as to when and how the zircon has been crystallized, which promises new vision for future study of the Napier Complex.

Conclusions The geological studies of JARE in the Napier Complex can be summarized as follows. 1.

2.

Field studies revealed that the Napier Complex comprises layered and massive gneiss units with nine stages of deformation history. The boundary between the units is transitional in the Mt. Riiser-Larsen area, but closely associated with the shear zone on Tonagh Island. Geochemical studies showed that (a) garnet– sillimanite gneisses and garnet-rich felsic gneisses were derived from mudstone and sandstone, respectively, both enriched in MgO, Cr and Ni, (b) orthopyroxene felsic gneisses have a close REE affinity with Archaean TTG, (c) basic gneisses were derived from LREE-enriched or -depleted basalts, and


3.

4.

(d) meta-ultramafic rocks are comparable with komatiite and related depleted mantle peridotite. This suite of protoliths is reminiscent of Archaean greenstone–granite belts. Precise analyses of physical conditions of the Napier UHT metamorphism by using new approaches indicated that the maximum metamorphic temperature reached 1130 8C. P–T evolution was examined by the analyses of reaction textures combined with fluid inclusion studies, and suggested both clockwise (Bunt Island) and counterclockwise (Mt. RiiserLarsen and Tonagh Island) P– T–t paths. U –Pb SHRIMP and SIMS zircon ages revealed three stages of protolith formation at c. 3.3, 3.0 and 2.6 Ga. The timing of peak UHT metamorphism is still unresolved between c. 2.55 and c. 2.51 –2.45 Ga.

On the basis of these results, further studies in the Napier Complex, such as regional structural and tectonic characterization based on geological field studies, microstructural studies that focus on the pre-UHT metamorphic imprints using advanced techniques such as LRS, and radiogenic and stable isotopes coupled with trace element imaging, should be carried out. The results of these studies will advance our understanding of the role of TTG protolith, causes and consequences of UHT crustal metamorphism and will help in modelling the Archaean continental evolution more accurately. I gratefully acknowledge all members of the JARE geology group, especially the Nishi–Higashi group, with whom I enjoyed frank discussions during the course of the study. I thank C. J. Carson, T. Hokada, M. Satish-Kumar and Y. Zhao for careful and constructive reviews of the manuscript.

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Antarctica: insights from zircon, monazite, and garnet ages. Journal of Geology, 114, 65–84. S UZUKI , S., I SHIZUKA , H. & K AGAMI , H. 2008. Early to Middle Proterozoic dykes in the Mt. Riiser-Larsen area of the Napier Complex, East Antarctica: tectonic implications as deduced from geochemical studies. In: S ATISH -K UMAR , M., M OTOYOSHI , Y., O SANAI , Y., H IROI , Y. & S HIRAISHI , K. (eds) Geodynamic Evolution of East Antarctica: A Key to the East– West Gondwana Connection. Geological Society, London, Special Publications, 308, 195–210. T AINOSHO , Y., K AGAMI , H., T AKAHASHI , Y., I IZUMI , S., O SANAI , Y. & T SUCHIYA , N. 1994. Preliminary result for the Sm–Nd whole-rock age of the metamorphic rocks from Mount Pardoe in the Napier Complex, East Antarctica. Proceedings of NIPR Symposium on Antarctic Geosciences, 7, 115– 121. T AINOSHO , Y., K AGAMI , H., H AMAMOTO , T. & T AKAHASHI , Y. 1997. Preliminary result for the Nd and Sr isotope characteristics of the Archaean gneisses from Mount Pardoe, Napier Complex, East Antarctica. Proceedings of NIPR Symposium on Antarctic Geosciences, 10, 92– 101. T AKAHASHI , N., K ODAIRA , S., K LEMPERER , S. L., T ATSUMI , Y., K ANEDA , Y. & S UYEHIRO , K. 2007. Crustal structure and evolution of the Mariana intra-oceanic island arc. Geology, 35, 203–206. T AKIGAMI , Y., I SHIKAWA , N. & F UNAKI , M. 1998. Preliminary 40Ar – 39Ar analyses of igneous and metamorphic rocks from the Napier Complex. Polar Geoscience, 11, 200– 207. T ATSUMI , Y. 2000. Continental crust formation by crustal delamination in subduction zones and complementary accumulation of the enriched mantle I component in the mantle. Geochemistry, Geophysics, Geosystems, 1, 2000GC000094, doi:10.1029/2000GC000094. T OYOSHIMA , T., O SANAI , Y., O WADA , M., T SUNOGAE , T., H OKADA , T. & C ROWE , A. W. 1999. Deformation of ultrahigh-temperature metamorphic rocks from Tonagh Island in the Napier complex, east Antarctica. Polar Geoscience, 12, 29– 48. T OYOSHIMA , T., Y AMAMOTO , K., O SANAI , Y., O WADA , M., T SUNOGAE , T., H OKADA , T. & C ROWE , W. A. 2001. Microstructures and deformation conditions of granulite-facies pseudotachylites from Tonagh Island, Napier Complex, East Antarctica. In: Abstracts of the 21st Symposium on Antarctic Geosciences, Tokyo, 75– 76 [in Japanese]. T OYOSHIMA , T., O SANAI , Y. & N OGI , Y. 2008. Macroscopic geological structures of the Napier and Rayner Complexes, East Antarctica. In: S ATISH -K UMAR , M., M OTOYOSHI , Y., O SANAI , Y., H IROI , Y. & S HIRAISHI , K. (eds) Geodynamic Evolution of East Antarctica: A Key to the East–West Gondwana Connection. Geological Society, London, Special Publications, 308, 139– 146. T SUNOGAE , T., O SANAI , Y., T OYOSHIMA , T., O WADA , M., H OKADA , T. & C ROWE , W. A. 1999. Metamorphic reactions and preliminary P– T estimates of ultrahigh-temperature mafic granulite from Tonagh Island in the Napier Complex, East Antarctica. Polar Geoscience, 12, 71– 86. T SUNOGAE , T., O SANAI , Y., T OYOSHIMA , T., O WADA , M., H OKADA , T. & C ROWE , W. A. 2000. Fluorine-rich

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2003. Fluid inclusions in an osumilite-bearing granulite from Bunt Island in the Archean Napier Complex, East Antarctica: implication for a decompressional P– T path? Polar Geoscience, 16, 61–75. W HITEHOUSE , M. J. & P LATT , J. P. 2003. Dating high grade metamorphism—constraints from rare-earth elements in zircon and garnet. Contributions to Mineralogy and Petrology, 145, 61–74. Y OSHIMURA , Y., M OTOYOSHI , Y., G REW , E. D., M IYAMOTO , T., C ARSON , C. J. & D UNKLEY , D. J. 2000. Ultrahigh-temperature metamorphic rocks from Howard Hills in the Napier Complex, East Antarctica. Polar Geoscience, 13, 60–85. Y OSHIMURA , Y., M IYAMOTO , T., G REW , E. D., C ARSON , C. J., D UNKLEY , D. J. & M OTOYOSHI , Y. 2001. High-grade metamorphic rocks from Christmas Point in the Napier complex, East Antarctica. Polar Geoscience, 14, 53–74.

Macroscopic geological structures of the Napier and Rayner Complexes, East Antarctica T. TOYOSHIMA1, Y. OSANAI2 & Y. NOGI3 1

Graduate School of Science and Technology, Niigata University, 8050 Ikarashi-2-nocho, Niigata 950-2181, Japan, (e-mail: [email protected]) 2

Division of Evolution of Earth Environments, Graduate School of Social and Cultural Studies, Kyushu University, Ropponmatsu 4-2-1, Fukuoka 810-8560, Japan 3

National Institute of Polar Research, Kaga 1-chome, Itabashi-ku, Tokyo 173-8515, Japan

Abstract: This paper presents a form-line map of the Napier and Rayner Complexes, East Antarctica, constructed from attitude data for foliations shown on published geological maps, and discusses the macroscopic geological structures. The form-line map shows that the two complexes consist of several, structurally distinct, units or blocks bounded by east–west-, NE–SW- and NW– SE-striking faults. The major boundary between the two complexes, as indicated on the published geological maps, is a structural discontinuity shown as a large fault on the form-line map. On the form-line map, east– west- and NE– SW-trending folds are abundant and NW– SE-trending ones occur locally in both complexes. North– south-trending folds are also abundant in the Napier Complex. Dome-and-basin fold patterns on a regional scale occur in some regions. The regional strikes, macroscopic structures, and the major boundary between the two complexes are considered to have resulted from the same later deformation episode. The form-line map and distribution map of key mineral assemblages show that the Napier Complex is not uniform and includes at least two types of metamorphic units or fragments of the Archaean crust that were formed through distinct P– T– t evolutionary processes and divided by several faults.

The Napier Complex of Enderby Land and western Kemp Land (index map in Fig. 1) is one of the several known Archaean cratonic blocks in the East Antarctic Precambrian Shield (e.g. James & Tingey 1983; Sheraton et al. 1987; Black et al. 1992). The complex is characterized by ultrahightemperature (UHT) metamorphic rocks, and is bounded on the south by younger mobile belts (Rayner Complex), like the other Archaean cratons (Fig. 1) (e.g. Sheraton et al. 1987). The structures of the Napier and Rayner Complexes have been outlined by geologists of the Bureau of Mineral Resource, Geology and Geophysics, Australia, on the basis of 1974–1977 field investigations. They have confirmed the integrity of the Napier Complex as a single highgrade Archaean craton by their assertion that the unmetamorphosed ‘Amundsen dykes’ are restricted to the complex (Sheraton et al. 1980, 1987). Sheraton et al. (1987) has reported the geological outline and structural features for the whole region of the Napier and Rayner Complexes. Toyoshima (2001) has preliminarily constructed a strike-line map of the two complexes and discussed their macroscopic geological structures. However, macroscopic geological structures and regional strikes in these two complexes remain poorly understood. The structural

geology of the boundary between the two complexes is not well known. Sheraton et al. (1987) have shown the extents of the two complexes and Napier Complex rocks overprinted by the Rayner metamorphism in their geological map and figures 4 and 16. However, geological structural relationships between the two complexes and the remetamorphosed Napier Complex rocks are not clear in the published geological maps (e.g. James & Black 1981; Ellis 1983; Sheraton et al. 1987; Harley & Hensen 1990). Harley (1998) has suggested that sapphirinebearing UHT metamorphic rocks from the northern Napier Complex represent lower pressures during the cooling stage than those from the areas south of the Amundsen Bay region, which is the highgrade part of the complex. On the basis of key mineral assemblages and reaction textures of UHT metamorphic rocks, Osanai et al. (2001b, c) and Hokada et al. (2008) have recently distinguished two metamorphic evolutionary processes; one shows isothermal decompression in part of a clockwise P–T –t path, and the other indicates isobaric cooling in part of a counterclockwise P–T– t path. They also suggested that the former path was due to the ‘true’ Napier metamorphism but the


Fig. 1. Form-line map of the Napier and Rayner Complexes, constructed from attitude data for foliations shown on geological maps of Sheraton et al. (1987), Ishikawa et al. (2000) and Osanai et al. (2001a). The boundary between the Napier and Rayner Complexes is from geological maps of Sheraton et al. (1987). Extensions of faults in Amundsen Bay are inferred from magnetic anomaly data of Nogi et al. (2001), and the total intensities of magnetic anomalies that are used by Nogi et al. (2001) have been included in ADMAP (Antarctic Digital Magnetic Anomaly Project) (Golynsky et al. 2001), which is shown in Figure 4. ABF, Amundsen Bay Fault; AFF, Adams Fjord Fault; CNF, Crosby Nunataks Fault; RLMSZ, Riiser-Larsen Main Shear Zone; TMF, Tula Mountains Fault.

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latter path shows overprinting or reworking by the Rayner metamorphism (Osanai et al. 2001b, c). Geological structural relationships between these UHT metamorphic rocks with different P–T –t paths also remain unknown. The dominant layering and lithological boundary-parallel main foliation in the Napier Complex resulted from early deformation (pre-D1 or D1 of James & Black (1981), Black & James (1983) and Sheraton et al. (1987); D1 –D2 of Toyoshima et al. (1999)). Where not significantly affected by later folding, the layering and foliation are subhorizontal, and the enveloping surfaces of the later fold generations are also subhorizontal (Sandiford & Wilson 1984). In the Rayner Complex, early deformation (Da of Sheraton et al. 1987) was also responsible for the formation of the main foliation and associated structures. These early structures have been folded and faulted locally by the later superimposed deformations. This paper presents a form-line map of the Napier and Rayner Complexes, and discusses macroscopic geological structures, regional strikes and P–T –t paths of the complexes.

Construction of form-line map Key horizons or marker beds cannot be regionally recognized in the Napier and Rayner Complexes. Therefore, a form-line map is constructed from attitude data for foliations shown on the geological maps of Sheraton et al. (1987), Ishikawa et al. (2000) and Osanai et al. (2001a). Most of data used in this paper are from Sheraton et al. (1987). Ishikawa et al. (2000) and Osanai et al. (2001a) provided data on small areas such as Mount RiiserLarsen and Tonagh Island. The main foliations on their geological maps are parallel to lithological boundaries and derived from the early deformation in each complex (e.g. Sheraton et al. 1987), and therefore we have assumed the main foliation in each complex as a structurally equivalent surface. The form-line map can indicate the forms or shapes of the regional structures involving the complexes. The contour lines at a given locality are parallel to the strike, and the spacing of the contour lines is roughly proportional to the dip angle. If contour lines are more closely spaced, the dip angle of the interval being contoured is steeper. The spacing, s, of the contour lines on the form-line map is given as s ¼ i cot u where u is a dip angle of the foliation and i ¼ 2 km is the contour interval (Fig. 1). The contour lines of the form-line map represent approximate lines of equal elevation but cannot be assigned specific values.

Before the contouring, two types of structural domains have been roughly recognized on the basis of the attitude data for foliations and lithological boundaries shown on the geological maps. One has nearly homoclinal structure, and the other shows changes in attitude of foliations. The changes indicate an abundance of folds with halfwavelengths of less than 20 km. A few examples of these structural domains and their structural features are as follows. (1) A domain around the Raggatt to southern Scott Mountains shows a north- to NWdipping homocline. (2) A domain around the northern Scott Mountains shows north-dipping foliations. (3) The southeastern part of the Nye Mountains shows a SE-dipping homocline. (4) A domain around Mount McGhee shows a NW-dipping homocline and asymmetrical minor folds. (5) A domain around Mount Stregutt to the northern Schwartz Range shows a NNE-dipping homocline and north-trending folds. (6) Changes in the orientations of foliations as well as folds are abundant in many domains around the Amundsen Bay, Tula Mountains, McDonald Ridge, the western and southern parts of Napier Mountains, Nicholas Range, southern Schwartz Range, Edward VIII Gulf, Dismal Mountains, Nye Mountains, and Mount Christensen. Fold geometry appears to be different in each domain. An interpretive contouring technique (e.g. Marshak & Mitra 1988) is used for the contour line drawing. First, contour lines are drawn in each domain. If the contour map is constructed with no faults between the domains, many linear zones of very closely spaced contours develop in the form-line map. The form-line map is then re-examined to determine if any faults may be present. Smoothed and harmonic contour lines can often be drawn by inferring faults between the structural domains and by substituting a fault for a zone of closely spaced contours (Marshak & Mitra 1988, pp. 27–34). The contour lines and fold axis are displaced along the inferred faults. The form-line map has been drawn freehand, modified by trial and error. The contour lines are drawn parallel to the lithological boundary shown on the geological maps (thick light blue lines in Fig. 1). Each of the attitude data points is assumed to represent the geological structure of an area of 20 –30 km2, if the location of one of the data points is far enough from that of other data points. We have also constructed a strike-line map of the complexes to check the inferred faults in the form-line map (Fig. 2). The strike-line map is constructed on the assumption that the thickness

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Fig. 2. Strike-line maps of some areas in the Napier and Rayner Complexes, constructed from attitude data for foliations shown on geological maps of Sheraton et al. (1987), Ishikawa et al. (2000) and Osanai et al. (2001a).

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of the imaginary bed is constant. The spacing, s, of the contour lines on the strike-line map is defined by s ¼ cosec u where u is a dip angle of the foliation and the thickness of the layer is 1 km. The assumption of constant layer thickness is the general method for constructing a strike-line map. Missing, ending-off and discontinuous contours on the strike-line map correspond to the disharmonic linear zones of closely spaced contours on the form-line map drawn with no faults (Figs 2 and 3). We have inferred faults on the form-line map if the contour lines and map satisfied the following criteria (Figs 2 and 3): (1) missing, ending-off and discontinuous of fold axes; (2) disharmonic linear zones of closely spaced contours; (3) contrasting structural features in two adjacent regions and domains; (4) sudden and disharmonious changes in orientations of foliations and contours; (5) missing, ending-off and discontinuous contours on the strike-line map. The inferred faults result in drawing smoothed and harmonic contour lines (Figs 1–3).

Interpretation of form-line map The form-line map (Fig. 1) suggests structural features as follows. (1) The Napier Complex consists of several structurally distinct units or blocks separated by east– west-, NE –SW- and NW–SE-striking faults. (2) Numerous faults were recognized in the Napier Complex. We call these faults Newman Nunataks Fault, Wilkinson Peaks Fault, Napier Mountains Fault, Tippet Nunataks Fault, Northern Tula Mountains Fault, Riiser-Larsen Main Shear Zone (Ishizuka et al. 1998), Adams Fjord Fault, Crosby Nunataks Fault, Tula Mountains Fault, Beaver Glacier Fault, Amudsen Bay Fault, Napier Fault, Scott Mountains Fault, Rayner Glacier Fault, and Nye Mountains Fault, approximately from east to west (Fig. 1). A north– south-trending fault at the western part of Mount Riiser-Larsen has been referred to as the Riiser-Larsen Main Shear Zone (Ishizuka et al. 1998), but we extend the fault to the north and to the eastern Amundsen Bay where the fault trends NE– SW (Fig. 1). Nogi et al. (2001) have compiled the vector geomagnetic anomaly data obtained on board the icebreaker Shirase in Amundsen Bay. The strikes of magnetic structures in Amundsen Bay are deduced from the vector magnetic anomalies using Seama et al. (1993) and the extensions of faults in Amundsen Bay are inferred from those. These structures are

also observed in total intensity magnetic anomalies of ADMAP (Golynsky et al. 2001) that include total intensity magnetic anomaly data from the vector magnetic anomalies obtained on board the icebreaker Shirase (Fig. 4). (3) The regional strike of the Napier Complex changes from east –west in the western part to NNE– SSE to NE–SW in the central to eastern part. Where not significantly affected by later folding, the foliation dips gently north or is subhorizontal. (4) The foliation of the Rayner Complex generally strikes NE –SW and dips south. (5) NE –SW- and east– west-trending faults occur in the Rayner Complex. (6) The major boundary between the two complexes (e.g. Sheraton et al. 1987) is a structural discontinuity regarded as a fault. In the eastern part (western Kemp Land), the transition between the two complexes appears to be more gradual, as noted by Sheraton et al. (1987). (7) East–west- and NE –SW-trending folds are abundant in both complexes, but north–southtrending folds occur in some areas. (8) A dome-and-basin fold pattern on a regional scale characterizes some areas of the two complexes, as shown by James & Black (1981). (9) Zones of most closely spaced contour lines on the form-line map can be regarded as intensely folded zones or faults. Some of these structural features on a regional scale have also been implied from the recognition of structural domains and their local structural features prior to the production of the form-line map.

Discussion and conclusions Comparison of structural data and interpretation with previous studies The regional strike of the Napier Complex, shown in Fig. 1, is similar to that in figure 52 of Sheraton et al. (1987). However, the structural domains separated by the faults in the form-line map are greatly different form the sub-areas (A– L) delineated in figure 52 of Sheraton et al. (1987). The folded and homoclinal areas are recognizable in the form-line map (Fig. 1). The generalized geological map pattern shown in figure 6 of Sheraton et al. (1987) is broadly similar to the structural pattern shown in the formline map (Fig. 1). In the area from Amundsen Bay to Mount McGhee, where the distribution pattern of layered gneisses and orthogneisses is complex in figure 6 of Sheraton et al. (1987), folds and faults are abundant in the form-line map (Fig. 1). The complicated distribution pattern can be

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Fig. 3. Form-line maps of some areas in the Napier and Rayner Complexes drawn with no faults, constructed from attitude data for foliations shown on geological maps of Sheraton et al. (1987), Ishikawa et al. (2000) and Osanai et al. (2001a).

Fig. 4. Total intensity magnetic anomaly map from ADMAP (Golynsky et al. 2001) and inferred geological structures in the Napier and Rayner Complexes. The magnetic anomaly patterns almost match the distribution of the inferred faults. Abbrevations as in Figure 1.

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explained by the pattern of folds and faults. In the area from the Scott Mountains to the Raggatt Mountains, where the distribution pattern of metamorphic rocks is simple on the geological map, a homoclinal structure occurs on the form-line map. Inferred faults form the boundaries between the different geological map patterns. The form-line map also indicates that the later folding and faulting gave rise to the transformation from a gently north-dipping or flat-lying structural state to a steeply inclined structural state, as shown by James & Black (1981) and Sandiford & Wilson (1984). The east– west- and NE– SWtrending folds inferred in Fig. 1 correspond to F3 of James & Black (1981), F3 of Sandiford & Wilson (1984), F3 of Sheraton et al. (1987), and B5 and B6 of Toyoshima et al. (1999), but north– south-trending folds to B4 of Toyoshima et al. (1999), on the basis of fold geometry. The east – west-, NE –SW- and NW –SE-striking faults

inferred on the form-line map can be correlated with D6 faults of Toyoshima et al. (1999), on the basis of their orientation.

Regional strikes and major structures of the Napier and Rayner Complexes The regional strike and dome-and-basin fold pattern of the Napier Complex in Fig. 1 are the products of the third deformation episode (D3 of James & Black (1981) and Sheraton et al. (1987); stage III of Toyoshima et al. (1999)). These faults and folds with near-horizontal axes have resulted from intense late-stage horizontal shortening (James & Black 1981; Toyoshima et al. 1999). The east–west- and NE–SW-trending folds, the north–south-trending folds, and the NE–SW- and east–west-striking faults also affect the south-dipping foliation of the Rayner Complex significantly (Fig. 1). The major boundary between the two complexes

Fig. 5. Distribution map of key mineral assemblages and inferred faults in the Napier Complex, complied from Sheraton et al. (1987), Ishizuka et al. (1998), Hokada et al. (1999), Osanai et al. (1999, 2001a– c) and Ishikawa et al. (2000). Abbreviations as in Figure 1.

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(e.g. Sheraton et al. 1987) can be correlated with the D6 fault of Toyoshima et al. (1999), on the basis of their orientations. Therefore, the regional strike, macroscopic structures of the Rayner Complex, and the major boundary between the two complexes are considered to have largely resulted from the third deformation episode. The locations of inferred faults correspond well to the borders between high and low magnetic anomalies, which represent the boundaries of magnetic structures, on the magnetic anomaly map of the Napier and Rayner complexes shown in Fig. 4 (Golynsky et al. 2001). Figure 4 also indicates that the northern Tula Mountains Fault is likely to extend northwestward over a distance of about 60 km. The Beaver Glacier Fault and Napier Fault may also extend westward for about 100 km.

Unity of Napier Complex as a high-grade Archaean craton The highest-grade metamorphic region (Harley & Hensen 1990, p. 327, high-grade region of fig. 12.3b) of the Napier Complex appears to be surrounded by the inferred faults on the form-line map (Fig. 1). A distribution map of the key mineral assemblages (Osanai et al. 2001b, c) and the inferred faults shows that, in the Napier Complex, the faults separate the clockwise P –T– t path metamorphic units from the counterclockwise P–T –t path metamorphic units (Fig. 5). The Napier Complex is considered to be not uniform in Archaean geological history, and includes at least two types of metamorphic units or crustal fragments that were formed through distinct P–T–t evolutionary processes and bounded by the inferred faults. Our conclusion is further supported by the results obtained from the Mt. Riiser-Larsen area by Hokada et al. (2008). The western boundary of the ‘true’ Napier metamorphic rocks is not drawn as a smooth line such as shown in figure 1 of James & Black (1981) but is composed of several faults such as the Tippet Nunataks Fault, Northern Tula Mountains Fault, Riiser-Larsen Main Shear Zone, Beaver Glacier Fault, Amundsen Bay Fault and Napier Fault (Fig. 5). We wish to thank F. Storti, K. Kano and M. Ishikawa for extremely helpful reviews, and M. Satish-Kumar for editorial handling of the manuscript. Their constructive comments greatly improved the paper. H. Ishizuka, K. Shiraishi, Y. Motoyoshi, M. Owada, T. Tsunogae and T. Hokada are also thanked for their invaluable discussions and helpful advice.

References B LACK , L. P. & J AMES , P. R. 1983. Geological history of the Napier Complex of Enderby Land. In: O LIVER ,

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R. L., J AMES , P. R. & J AGO , J. B. (eds) Antarctic Earth Science. Australian Academy of Science, Canberra, A.C.T., 11–15. B LACK , L. P., S HERATON , J. W. & K INNY , P. D. 1992. Archaean events in Antarctica. In: Y OSHIDA , Y., K AMINUMA , K. & S HIRAISHI , K. (eds) Recent Progress in Antarctic Earth Science. Terra, Tokyo, 1– 6. E LLIS , D. J. 1983. The Napier and Rayner Complexes of Enderby Land, Antarctica—Contrasting styles of metamorphism and tectonism. In: O LIVER , R. L., J AMES , P. R. & J AGO , J. B. (eds) Antarctic Earth Science. Australian Academy of Science, Canberra, A.C.T., 20– 24. G OLYNSKY , A., C HIAPPINI , M., D AMASKE , D. ET AL . 2001. ADMAP—Magnetic Anomaly Map of the Antarctic, 1:10 000 000 scale map. In: M ORRIS , P. & VON F RESE , R. (eds) British Antarctic Survey Miscellaneous, 10, British Antartic Survey, Cambridge. H ARLEY , S. L. 1998. On the occurrence and characterization of ultrahigh-temperature crustal metamorphism. In: T RELOAR , P. J. & O’B RIEN , P. J. (eds) What Drives Metamorphism and Metamorphic Reactions? Geological Society, London, Special Publications, 138, 81– 107. H ARLEY , S. L. & H ENSEN , B. J. 1990. Archaean and Proterozoic high-grade terranes of East Antarctica (40–808E): a case study of diversity in granulite facies metamorphism. In: A SHWORTH , J. R. & B ROWN , M. (eds) High-temperature Metamorphism and Crustal Anatexis. Unwin Hyman, London, 320– 370. H OKADA , T., O SANAI , Y., T OYOSHIMA , T., O WADA , M., T SUNOGAE , T. & C ROWE , W. A. 1999. Petrology and metamorphism of sapphirine-bearing aluminous gneisses from Tonagh Island in the Napier Complex, East Antarctica. Polar Geoscience, 12, 49– 70. H OKADA , T., M OTOYOSHI , Y., S UZUKI , S., I SHIKAWA , M. & I SHIZUKA , H. 2008. Geodynamic evolution of Mt. Riiser-Larsen, Napier Complex, East Antarctica, with reference to the UHT mineral associations and their reaction relations. In: S ATISH -K UMAR , M., M OTOYOSHI , Y., O SANAI , Y., H IROI , Y. & S HIRAISHI , K. (eds) Geodynamic Evolution of East Antarctica: A Key to the East– West Gondwana Connection. Geological Society, London, Special Publications, 308, 253–282. I SHIKAWA , M., H OKADA , T., I SHIZUKA , H., M IURA , H., S UZUKI , S., T AKADA , M. & Z WARTZ , D. P. 2000. Geological map of Mount Riiser-Larsen, Enderby Land, Antarctica. Antarctic Geological Map Series, Sheet 37. National Institute of Polar Research, Tokyo. I SHIZUKA , H., I SHIKAWA , M., H OKADA , T. & S UZUKI , S. 1998. Geology of the Mt. Riiser-Larsen area of the Napier Complex, Enderby Land, East Antarctica. Polar Geoscience, 11, 154–171. J AMES , P. R. & B LACK , L. P. 1981. A review of the structural evolution and geochronology of the Archaean Napier Complex of Enderby Land, Australian Antarctic Territory. In: G LOVER , J. E. & G ROVES , D. I. (eds) Archaean Geology. Geological Society of Australia, Special Publications, 7, 71– 83. J AMES , P. R. & T INGEY , R. J. 1983. The Precambrian geological evolution of the East Antarctic metamorphic shield—a review. In: O LIVER , R. L.,

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J AMES , P. R. & J AGO , J. B. (eds) Antarctic Earth Science. Australian Academy of Science, Canberra, A.C.T., 5–10. M ARSHAK , S. & M ITRA , G. 1988. Basic Methods of Structural Geology. Prentice Hall, Englewood Cliffs, NJ. N OGI , Y., S EAMA , N. & I NOKUCHI , H. 2001. Magnetic anomalies in the Amundsen Bay, East Antarctica. In: Program and Abstracts for 21st Symposium on Antarctic Geosciences. National Institute of Polar Research, Tokyo, 29 [in Japanese]. O SANAI , Y., T OYOSHIMA , T., O WADA , M., T SUNOGAE , T., H OKADA , T. & C ROWE , W. A. 1999. Geology of ultrahigh-temperature metamorphic rocks from Tonagh Island in the Napier Complex, East Antarctica. Polar Geoscience, 12, 1– 28. O SANAI , Y., T OYOSHIMA , T., O WADA , M. ET AL . 2001a. Geological map of Tonagh Island, Enderby Land, Antarctica. Antarctic Geological Map Series, Sheet 38. National Institute of Polar Research, Tokyo. O SANAI , Y., T OYOSHIMA , T., O WADA , M., T SUNOGAE , T., H OKADA , T., C ROWE , W. A. & K USACHI , I. 2001b. Ultrahigh-temperature sapphirine–osumilite and sapphirine–quartz granulites from Bunt Island in the Napier Complex, East Antarctica— Reconnaissance estimation of P –T estimation. Polar Geoscience, 14, 1–24. O SANAI , Y., T OYOSHIMA , T., O WADA , M., T SUNOGAE , T., H OKADA , T., S ASANO , K. & C ROWE , W. A. 2001c. Clockwise and counter-clockwise P –T paths of UHT metamorphic rocks from the Napier Complex, East Antarctica. In: Program and Abstracts for 21st Symposium on Antarctic Geosciences.

National Institute of Polar Research, Tokyo, 20 [in Japanese]. S ANDIFORD , M. & W ILSON , C. J. L. 1984. The structural evolution of the Fyfe Hills–Khmara Bay region, Enderby Land, East Antarctica. Australian Journal of Earth Sciences, 31, 403– 426. S EAMA , N., N OGI , Y. & I SEZAKI , N. 1993. A new method for precise determination of the position and strike of magnetic boundaries using vector data of the geomagnetic anomaly field. Geophysical Journal International, 113, 155– 164. S HERATON , J. W., O FFE , L. A., T INGEY , R. J. & E LLIS , D. J. 1980. Enderby Land, Antarctica—an unusual Precambrian high-grade metamorphic terrain. Journal of the Geological Society of Australia, 27, 1– 18. S HERATON , J. W., T INGEY , R. J., B LACK , L. P., O FFE , L. A. & E LLIS , D. J. 1987. Geology of Enderby Land and Western Kemp Land, Antarctica. Australian Bureau of Mineral Resources, Geology and Geophysics, Bulletin, 223. T OYOSHIMA , T. 2001. A strike line map and geological structures of the Napier and Rayner complexes, East Antarctica. In: Program and Abstract for 21st Symposium on Antarctic Geosciences. National Institute of Polar Research, Tokyo, 77–78 [in Japanese]. T OYOSHIMA , T., O SANAI , Y., O WADA , M., T SUNOGAE , T., H OKADA , T. & C ROWE , W. A. 1999. Deformation of ultrahigh-temperature temperature metamorphic rocks from Tonagh Island in the Napier Complex, East Antarctica. Polar Geoscience, 12, 29–48.

Pre-metamorphic carbon, oxygen and strontium isotope signature of high-grade marbles from the Lu¨tzow-Holm Complex, East Antarctica: apparent age constraints of carbonate deposition M. SATISH-KUMAR1, T. MIYAMOTO2, J. HERMANN3, H. KAGAMI4, Y. OSANAI5 & Y. MOTOYOSHI6 1

Institute of Geosciences, Faculty of Science, Shizuoka University, 836 Oya, Suruga-ku, Shizuoka, 422-8529, Japan (e-mail: [email protected]) 2

Department of Earth and Planetary Sciences, Faculty of Science, Kyushu University, Fukuoka, 812-8581, Japan 3

4

Research School of Earth Sciences, The Australian National University, Canberra A.C.T. 0200, Australia

Department of Geology, Faculty of Science, Niigata University, Ikarashi-2no-cho, Niigata, Japan 5

Division of Evolution of Earth Environment, Graduate School of Social and Cultural Studies, Kyushu University, Fukuoka, 810-8560, Japan

6

National Institute of Polar Research, Kaga, Itabashi-ku, Tokyo, 173-8515, Japan Abstract: C, O and Sr isotope geochemistry of high-grade marbles from the Lu¨tzow-Holm Complex, East Antarctica, has given clues on the depositional ages and post-depositional alterations. Dolomitic and calcitic marbles occur as thin layers with varying thickness (up to 100 m) in several outcrops in eastern Dronning Maud Land, most of which underwent postdepositional geochemical alterations. In particular, the Sr and O isotope alterations are extensive, with 87Sr/86Sr(550 Ma) ratios as high as 0.758 and d18O values as low as 25‰. These data suggest that multiple stages of fluid–rock interaction processes during diagenesis, prograde to peak and retrograde metamorphic events have altered the depositional isotopic signatures. However, some of the marble layers, exceptionally, preserve pre-metamorphic geochemical characteristics, such as low Sr isotope ratios, high d18O and d13C values, and well-equilibrated unaltered trace and rare earth element patterns. Lowest 87Sr/86Sr isotopic ratios of 0.7066 and 0.7053 with high d13C and d18O values suggest an apparent age of deposition around 730–830 Ma, although total geochemical resetting of carbonates by seawater of this age cannot be ruled out. The apparent depositional ages are consistent with carbonate deposition in the ‘Mozambique Ocean’ that separated East and West Gondwana.

Extensive metasedimentary supracrustal sequences exposed in the crustal fragments of the East Gondwana supercontinent, especially in East Antarctica, Sri Lanka, peninsular India and Madagascar, provide us with an opportunity to understand the geodynamic evolution of supercontinent assembly and breakup as well as extract key information on the depositional environments of palaeo-oceans that separated proto-continents. The closure of the Neoproterozoic ‘Mozambique Ocean’ (Hoffman 1988, 1991; Dalziel 1991; Stern 1994) is considered to be a consequence of supercontinental assembly of East Gondwana and West Gondwana during a protracted Pan-African Orogeny that spatially

extends from the Arabian–Nubian Shield to East Antarctica, through East Africa, Madagascar, Southern India and Sri Lanka (Bauer et al. 2003; Kusky & Matsah 2003). The difficulty in constraining the characteristics of the ‘Mozambique Ocean’ is mainly due to the high-grade metamorphism and tectonic reworking of the sediments during the regional Pan-African Orogeny. However, metacarbonate is a lithology that not only preserves important evidence on the metamorphic and geochemical evolution during tectonic activities (Bickle et al. 1995, 1997; Satish-Kumar et al. 1998) but also provides valuable information on the palaeo-ocean geochemistry and environment


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of deposition (Melezhik et al. 2005). To recognize such information and to properly interpret it, a multidimensional geochemical approach has to be employed. Given the advances in microgeochemical analytical techniques, it is now possible to fingerprint and differentiate multistage geological processes accurately in space and time. In this study, we review the progress made in the field of metacarbonate geochemistry and attempt to identify pristine geochemical sedimentary signatures from high-grade marbles distributed in the Lu¨tzow-Holm Complex (LHC), East Antarctica. Lower to middle continental crust exposed in the Lu¨tzow-Holm Complex, East Antarctica (Fig. 1) is an ideal terrain to investigate crustal processes, especially for geochemical studies, because bedrocks are exposed continuously for kilometres in a single outcrop. Geological, tectonic and fluidrelated processes were the focus of several earlier studies in this region (Hiroi et al. 1983, 1986, 1987; Motoyoshi et al. 1989; Shiraishi et al. 1994, 2003; Motoyoshi & Ishikawa 1997; Satish-Kumar et al. 1998). Previous studies that documented fluid-related processes were based on mineralogy, phase petrology and fluid inclusions (Santosh & Yoshida 1998; Satish-Kumar et al. 2006a), grainscale carbon and oxygen stable isotopes (SatishKumar et al. 1998), LA-ICPMS study of trace and

rare earth elements (Satish-Kumar et al. 2006a) and Sr isotopes (Satish-Kumar et al. 2006b). Carbon, oxygen and strontium isotope studies of marbles from several outcrops exposed in the c. 400 km coastal stretch of the LHC are reviewed here. Integrating the geochemical characteristics, we discuss the influence of metamorphic fluid – rock interaction in modifying the isotopic characteristics and attempt to identify marbles that preserve pre-metamorphic signatures. We also discuss the importance of metacarbonate rocks in understanding the palaeo-ocean that separated East and West Gondwana.

Geological background and previous studies The LHC extends c. 400 km in eastern Dronning Maud Land, bounded in the NE by the Rayner Complex and in the SW by the Yamato–Belgica Complex. This region is characterized by an increase in metamorphic grade from amphibolite facies in the NE to granulite to ultrahigh-temperature metamorphic facies in the SW (Hiroi et al. 1991; Motoyoshi & Ishikawa 1997; Yoshimura et al. 2008). The dominant lithological units that crop out in this region include pelitic to psammitic gneisses,

Fig. 1. The Lu¨tzow-Holm Complex in East Antarctica, showing the localities where thick marble layers are distributed. The numbers in parenthesis are 87Sr/86Sr ratios of marbles corrected to 550 Ma.

APPARENT DEPOSITIONAL AGES OF MARBLES

149

Fig. 2. P– T fluid evolution of the Skallen region, a representative region where thick marble horizons are exposed (modified after Satish-Kumar et al. 2006a).

basic to intermediate gneisses, and subordinate calcareous and ultramafic rocks. In general, the occurrence of kyanite inclusions within garnet surrounded by sillimanite in the matrix has been reported from several outcrops in the region, indicating a regional clockwise P–T evolution (Motoyoshi et al. 1989; Hiroi et al. 1991; Yoshimura et al. 2008). Based on detailed petrological and geochemical study on scapolite-bearing rocks from the Skallen region, Satish-Kumar et al. (2006a) identified multiple fluid–rock interaction events during a protracted metamorphic evolution (Fig. 2). Detailed geochronological studies have also been carried out in this region that revealed a prolonged late Proterozoic to early Cambrian tectonothermal event, in concurrence with the widespread Pan-African events reported in the adjoining East Gondwana continental fragments (Shiraishi et al. 1994, 2003; Fraser et al. 2000; Hokada & Motoyoshi 2006). Marbles in the LHC occur as layers, up to about 100 m thick, interlayered with metapelitic and quartzofeldspathic rocks (Fig. 3a). The marble layers are often separated from the adjoining metapelitic rocks by decimetre- to metre-scale coarse-grained skarn formations, which consist of diopside, phlogopite, amphibole and spinel. Hiroi et al. (1987) carried out a detailed study on the occurrence, mineralogy and reaction textures observed in the calc-silicate rocks and marbles in the Prince Olav Coast and Soya Coast

Fig. 3. (a) Field occurrence of a thick marble layer at Skallevikshalsen. (b) Rhythmic layering of pure dolomite marble and forsterite– spinel calcite marble.

150


and recognized wollastonite-in and epidote-out isograds. Humite-bearing assemblages were reported from Kasumi Rock in the Prince Olav Coast (Hiroi & Kojima 1988). Recently, several new occurrences of calc-silicate rocks and marbles have been identified from the LHC (Satish-Kumar et al. 2006c). Earlier research on stable isotope characteristics of marbles from the Skallen region identified important evidence for fluid –rock interaction process. Extremely large (.20‰) oxygen isotope heterogeneities at a sub-millimetre scale were discovered in a marble layer from Skallen, which were interpreted to have resulted from grain boundary migration of aqueous fluids of meteoric origin (Satish-Kumar & Wada 1997; Satish-Kumar et al. 1998). Marbles that were not affected by fluid – rock interaction were useful in identifying the peak metamorphic temperatures using carbon isotope thermometry between calcite and graphite (Satish-Kumar & Wada 2000). In addition, recent strontium isotope studies considered a possible fluid–rock interaction event during diagenetic or early prograde metamorphism (Satish-Kumar et al. 2006a). Thus, the majority of previous studies on stable isotopes of metacarbonates rocks were focused on constraining the metamorphic evolution and less attention has been paid to understanding the pre-metamorphic features of the carbonates.

Field relations and sample descriptions Mineralogical and geochemical studies of marble samples collected from several outcrops in the LHC were carried out in this study. In particular, marbles from the layered sequence at Skallen (sampled during the 39th Japanese Antarctic Research Expedition (JARE-39) and Skallevikshalsen (sampled during JARE-46) (southwestern LHC; Fig. 1) were selected for detailed study. Major rock types in this region are orthopyroxene-bearing felsic gneiss, garnet–sillimanite gneiss, garnet–biotite gneiss and marbles (Yoshida 1977, 1978; Osanai et al. 2004). Marble layers occur conformably with thickness up to several tens of metres (Fig. 3a). Several typical high-grade marble mineral assemblages are observed, which form thin decimetre-scale rhythmic layers of pure dolomite or calcite marbles (graphite-bearing) and forsterite þ diopsideþ spinel + phlogopite + amphibole-bearing marbles (Fig. 3b). Various kinds of skarn occurrences between marble and the adjoining gneisses are also observed (Matsueda et al. 1983). Decimetre-sized scapolite- or feldspar boudins surrounded by phlogopite- or amphibole-rich reaction rims are present within the marble (Satish-Kumar

et al. 2006b, c). After careful field and hand-specimen observations we selected representative pure marble samples comprising either calcite or dolomite. Pure marbles are likely to preserve pre-metamorphic features, as the rocks might not have been affected by metamorphic processes such as fluid–rock interactions and devolatilization reactions, except for coarsening and recrystallization. To compare and contrast the effect of metamorphic processes on pre-metamorphic geochemical signatures, two representative samples from the Skallen region are studied in detail here. Sample 2305E was collected from a layer comprising an interface between a pure dolomite layer and a phlogopite– forsterite–spinel-bearing calcite-rich marble layer (Fig. 4a). Notably, the modal proportion of calcite increases drastically with a corresponding increase in the content of silicate minerals. This implies that calcite might be a product of metamorphic reactions involving dolomite and silicate phases in the precursor sediments. Furthermore, during prograde metamorphism these metamorphic (decarbonation) reactions progress with an increase in temperature and/or infiltration of aqueous fluids (Ferry 1994). Sample 2305E can therefore be considered to have been affected by metamorphic fluid–rock interaction processes. Sample 602d, a coarse-grained pure dolomite marble, was collected from a layer in the SW part of the Skallen region. Other than dolomite, the marble contains minor amounts of graphite (Satish-Kumar & Wada 2000). Silicate phases are absent in this sample and therefore this sample is considered as a potential candidate that may preserve pre-metamorphic geochemical features. Additional results from other marble horizons in the LHC, namely, Skallevikshalsen (JARE-46), Kasumi Rock (JARE-20), Kabuto Rock (JARE46) and Breidvagnipa (JARE-35), are also reported in this study. Moreover, we also take into account the existing carbon, oxygen and strontium isotopic results for the marbles from the LHC reported in the literature for discussing the regional implications (Satish-Kumar et al. 1998, 2006b; Satish-Kumar & Wada 2000). Based on evaluation of the geochemical data, we attempt to distinguish potential marbles that might preserve premetamorphic geochemical signatures that help in constraining a minimum depositional age of sedimentation.

Analytical methods Petrography of polished thin sections of marble samples was carried out using cathodoluminescence (CL) microscopy. The CL system (ELM-3R


151

Fig. 4. (a) Example of a stained (Alizarin red-S) marble slab (Skallen-2305E) showing pink-coloured calcite and colourless dolomite. Sampling spots with results of carbon, oxygen and strontium isotope measurements are also shown. (b) Cathodoluminescence image of marble, showing the textural relations between calcite, dolomite, forsterite and diopside from Skallen region. The green rim of diopside is formed during retrograde metamorphism. (c) Purple-coloured homogeneous CL image of coarse-grained dolomite marble from Skallen. The fractures appear bright pink.

Luminoscopew) at Shizuoka University comprises an electron gun, which can generate an electron beam with an accelerating voltage up to 30 keV and a beam current of 1 mA, fixed on the stage of a normal petrological microscope equipped

with a digital photographic attachment (Nikonw Coolpix 995). Silicate and carbonate minerals were analysed for major elements using a wavelength-dispersive electron microprobe (JEOL JXA733; Shizuoka

152


University). Measurement conditions were 12 nA, 15 kV accelerating voltage and a focused beam for silicate minerals, whereas a defocused beam (diameter 10 mm) was used for carbonates. Bence & Albee (1968) correction with modified a factors of Nakamura & Kushiro (1970) was performed throughout. Trace element and rare earth element (REE) analyses were performed on polished thin sections using the laser-ablation inductively coupled plasma mass spectrometry (ICP-MS) facility at the Research School of Earth Sciences (RSES) at the Australian National University (ANU), Canberra. A pulsed 193 nm ArF excimer laser with 70 mJ energy at a repetition rate of 5 Hz coupled to an Agilent 7500 quadrupole ICP-MS system was used for ablation. During the timeresolved analysis, the contamination from fractures or included phases was detected by monitoring several elements and integrating only the ‘clean’ part of the signal. A spot size of 142 mm was used for analysing coarse minerals, whereas thin rims and reaction zones were analysed using reduced spot sizes of either 82 or 54 mm. An NIST-612 glass was used as the external standard and a BCR-2G glass was used as secondary standard. Internal standards employed CaO contents measured using electron microprobe analysis (EMPA) for respective minerals. Average REE contents of several spot measurements (2–6) for each mineral textural site of representative carbonates were normalized to chondrite (McDonough & Sun 1995). Relative 1s standard deviation of analyses is generally about 5–20%. Carbon and oxygen isotopic composition were measured for carbonates (both calcite and dolomite) that were separated using a knife edge by scraping from a polished slab of marble that was previously stained with Alizarin red-S, to distinguish between calcite and dolomite. Pulverized dolomite or calcite was then placed in small stainless steel thimbles and dropped into a reaction vessel containing concentrated phosphoric acid at 60 8C in vacuum to liberate CO2 (Wada et al. 1984). The liberated CO2 gas was then purified cryogenically for analysis. Stable isotope measurements were carried out with a Finnigan MAT-250 mass spectrometer (Shizuoka University). Machine standards calibrated to NBS-20 standard yield reproducibility of 0.03‰ for d13C and 0.05‰ for d18O. The results are reported in conventional d notation related to the V-PDB standard for carbon and V-SMOW standard for oxygen, and are presented in Table 1. Carbonates were then separated for strontium isotope analysis from the same sample slabs from which sampling for stable isotopes was performed (Fig. 4a). Sample powder (3– 10 mg) was either scraped using a knife edge or drilled using a dentists

drill under a microscope. Conventional isotope dilution methods were applied to determine Rb and Sr compositions of the samples. Powdered sample fractions were dissolved with HCl– (COOH)2 mixed acid, and then passed through a Dowex 50W-X8 cation exchange resin to separate Rb and Sr. Rb and Sr compositions were determined with a Hitachi RMU5G and a JEOL JMS05RB mass spectrometer at Kyushu University (Japan). The E & A strontium standard gave values of 87 Sr/86Sr ¼ 0.7080 + 0.0001 (1s) and NBS987 gave values of 87Sr/86Sr ¼ 0.71025 + 0.00010 (1s). The contamination levels of Rb and Sr are 1 1029 g and 3 10210 g per sample. The decay constant used for 87Rb is 1.42 10211 a21 (Steiger & Ja¨ger 1977). Additional Sr isotopic ratio and concentration were analysed using a MAT 262 mass spectrometer at Niigata University (Japan), following the same procedure. Sr isotope results are summarized in Table 1.

Mineralogy and textural features Marbles from the LHC are generally coarse-grained with well-formed crystals, and show hypidiomorphic and granoblastic texture. The marbles are predominantly composed of dolomite or calcite. Other constituent minerals are forsterite, diopside, phlogopite, spinel, apatite, pargasite and tremolite. Humite-group minerals were reported from the Kasumi Rock locality (Hiroi & Kojima 1988). A minor amount of graphite is present in some of the marble horizons. All silicate mineral phases have near end-member chemical composition with only limited variations (Hiroi et al. 1987; Hiroi & Kojima 1988). CL imagery is a useful petrological indicator for identifying fluid –rock interaction processes (e.g. Yardley & Lloyd 1989; Wada et al. 1998). The majority of the marbles from the LHC display normal red to dark red CL, whereas some samples show yellow CL along fractures and grain boundaries. Yellow CL in calcite is characteristic of domains that were affected by fluid interaction and/or recrystallization (Buick & Cartwright 2003). Detailed isotopic (C, O and Sr) studies of samples that show signatures of fluid–rock interaction processes during metamorphism have been discussed by Satish-Kumar et al. (2006b). However, for comparison, new results obtained from one sample (2305E) are discussed in detail here. Coarse dolomite and calcite in this sample display purple and orange–red CL (Fig. 4b). At places, along fractures and grain boundaries, calcite displays yellow CL. Thin diopside rims around forsterite show yellowish green CL (Fig. 4b). These CL textures are characteristic of retrograde fluid–rock interaction and related chemical

Table 1. Carbon, oxygen and strontium isotope composition of marbles from Lu¨tzow-Holm Complex, East Antarctica d13C(PDB)

d18O(SMOW)

Rb (ppm)

Sr (ppm)

Calcite Calcite Dolomite Calcite Calcite

20.23 20.24 20.24 20.61 20.42

14.92 14.93 13.35 14.36 14.36

0.02 0.05 0.04 0.03 0.05

198.3 206.9 116.1 116.3 127.6

0.72534 0.72609 0.72538 0.72537 0.72497

0.00002 0.00006 0.00007 0.00008 0.00009

Dolomite Dolomite Dolomite Dolomite

1.11 0.81 1.09 0.71

17.02 17.04 16.91 16.69

0.03

170.1

0.70702

0.00006

0.03

349.2

0.70663

0.00006

Calcite Calcite

24.54 24.44

11.71 11.66

— —

— —

0.71123 0.71136

0.00002 0.00006

Calcite Calcite Calcite

22.20 22.06 22.76

14.26 14.12 13.31

0.18

109

0.70731

0.00006

0.75 0.75 0.62 0.73 1.95 1.29 1.13

19.19 18.94 19.12 19.08 19.66 19.78 19.68

0.03 —

120 74

0.72556 0.73287

0.00004 0.00001

—

33

0.73057

0.00001

— — — —

— — 255 193

0.75821 0.75739 0.70999 0.70996

0.00002 0.00003 0.00001 0.00001

Sample

87

Sr/86Sr(550 Ma)

1s

Skallen

Breidvagnipa N94-2206-31M N94-2206-60M Kasumi Rock 79020411-CC1 CC-2 CC-3 Skallevikshalsen TM99-0206-02A SK05-120-1-1-2

SK05-119-2-3-1 SK05-116-1-14

Dolomite Calcite-1 Calcite-2 Calcite-3 Calcite-4 Dolomite-1 Dolomite-2 Dolomite-3 Calcite Dolomite Calcite Dolomite

0.06 0.22

19.21 19.97


A97-2305E CC1 CC2 CC3 CC4 CC5 Y69-0602d CC1 CC2 CC3 Y69-0602e

153

154


alterations. Furthermore, exsolution of dolomite in calcite is a common feature in calcite-rich domains. In contrast, coarse monomineralic dolomite and calcite marbles display characteristic homogeneous purple and orange– red CL as shown in Figure 4c (Sample 602d), except for healed fractures, which show brighter pink CL. Exsolution of calcite in dolomite is observed; however, this feature is less prominent than in calcite-rich marbles. Given the homogeneity in CL images, monomineralic dolomite and calcite marbles are considered to have less been affected by fluid –rock interaction processes.

Carbonate mineral chemistry and metamorphic conditions Appreciable mineral chemical compositional variations are found in dolomite and calcite. In particular, the calcite-rich marbles from the southwestern LHC show pronounced exsolution textures. The Ca–Mg covariations clearly suggest that exsolution dominated compositional re-equilibration during retrograde metamorphism (Fig. 5). Dolomite displays minor variations in Fe and Mn contents, with Fe contents marginally higher than the Mn contents. The Fe and Mn contents of calcite overlap, although a slightly higher Fe than Mn content is observed in calcite. Forsterite þ spinel is the common equilibrium assemblage in marbles from the southwestern LHC. This assemblage is indicative of granulitefacies conditions. Peak metamorphic temperature estimates of 860 + 15 8C were reported from marbles in the Skallen region using calcite –graphite carbon isotope thermometry (Satish-Kumar & Wada

2000). Mizuochi et al. (2006) estimated the metamorphic temperature condition of marbles at Skallevikshalsen using calcite –dolomite solvus thermometry. The matrix calcite yielded lower temperatures corresponding to retrograde exsolution (400 –750 8C), whereas reintegrated inclusions within forsterite or spinel gave temperatures ranging from 850 to 950 8C, comparable with peak metamorphic conditions estimated from adjoining metapelitic assemblages (Yoshimura et al. 2004).

Carbonate trace element composition Carbonate trace element compositions combined with petrography and CL images provide a powerful tool to distinguish between marbles that preserve fully equilibrated calcite and dolomite, and those altered by retrograde interaction with fluids. Figure 6a shows the REE patterns of calcite and dolomite from a dolomite marble (602d) that displays well-equilibrated grains and normal red CL colours. The patterns plot in a tight range and indicate that calcite and dolomite are in REE equilibrium. Narrow deformation bands in dolomite are clearly visible in CL as pink domains (Fig. 4b). These domains are decorated by tiny fluid inclusions and appear as dusty dolomite. The REE pattern of this dusty dolomite is indistinguishable from that of the normal dolomite domains. Figure 6b displays the REE patterns of calcite and dolomite from an olivine marble (2305E) that contains carbonate grains that are partly dusty and display optical evidence of alteration, abundant trails of fluid inclusions and domains with yellow CL. Interestingly, the REE patterns of calcite still define a tight range, indicating that REE have not

Fig. 5. Mineral chemistry variations of calcite and dolomite from Skallen marbles.


Fig. 6. Chondrite-normalized REE diagram for carbonate minerals in (a) a ‘normal’ marble (602d: 87 Sr/86Sr ¼ 0.707) and (b) a high-Sr marble (2305E: 87 Sr/86Sr ¼ 0.725) from the Skallen region.

been significantly mobilized by potential alteration. The wider range in REE observed in dolomite is mainly produced by different proportions of exsolved calcite within the dolomite. Carbonates also contain significant amounts of fluid-mobile elements such as Sr and Pb. In sample 602d, calcite and dolomite define a tight cluster in a Sr v. Pb plot (Fig. 7). Moreover, the dusty dolomite contains only a tiny amount more Pb than the clear domains. In contrast, there is a wide scatter in carbonate Sr and Pb contents from sample 2305E. The clean calcite domains do not define a cluster and the altered calcite domains, dusty calcite with yellow CL and calcite containing abundant fluid inclusions plot in a wide field with overall higher Pb and slightly higher Sr contents, indicating abundant alteration within this sample. A similar, although less pronounced pattern is observed in dolomite. Interestingly, the domains in dolomite that contain abundant calcite exsolutions plot mainly in the field between unaltered calcite and dolomite. This indicates that the alteration responsible for the scatter in Pb and Sr occurred after the exsolution event (i.e. during retrograde metamorphism). The influx of an H2O-rich fluid phase during retrograde metamorphism, enriched

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Fig. 7. Trace element variation (Pb v. Sr) observed in two representative marble samples (602d and 2305E) from Skallen. Sample 602d shows a tight cluster for both calcite and dolomite, whereas the data for 2305E are scattered, indicating alteration during metamorphism. Fl, fluid inclusions.

in Sr and Pb as schematically shown in Figure 2, might be responsible for the alteration.

Carbon and oxygen isotope geochemistry Data from earlier studies as well as newly obtained data on carbon and oxygen isotopic composition of marbles from various exposures in the LHC are complied in a C –O isotopic variation diagram (Fig. 8), where considerable scatter can be observed. In general, carbonate rocks deposited during late Proterozoic time have carbon and oxygen isotopic compositions of .0‰ and 25 + 5‰, respectively (Veizer et al. 1992; Halverson et al. 2005), except for narrow intervals of glaciation. In general, based on the C –O isotopic composition, marbles can be classified into three categories: marbles that possibly preserve premetamorphic isotope signatures (Type-I), marbles that follow extensive devolatilization-induced isotope trends (Type-II), and marbles that were affected by retrograde fluid infiltration events (Type-III). Type-I marbles are comparatively rare

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Fig. 8. Compilation of existing carbon and oxygen isotopic composition on a d13C v. d18O variation diagram, and the results obtained in this study. The curved and straight lines were decarbonation trends of marbles by Rayleigh and batch devolatilization calculated after Baumgartner & Valley (2001). Data source for Skallen marbles are Satish-Kumar & Wada (2000) and Satish-Kumar et al. (1998, 2006b).

in the LHC. Although some of the marble horizons display carbon isotopic composition characteristic of Neoproterozoic unmetamorphosed carbonate rocks (.0‰), their oxygen isotopic values are generally lower than 20‰. Indeed, it is rather difficult to preserve pre-metamorphic oxygen isotope composition of carbonates during high-grade metamorphism, because the interaction between carbonate and externally derived aqueous fluids can easily shift oxygen isotopes even during the earliest stages of diagenesis or prograde metamorphism. Therefore, oxygen isotopes serve as a first-hand proxy for defining the extent of geochemical alteration during metamorphism. Many marbles in the LHC display coupled C–O isotopic depletion, which we group as Type-II. This trend is shown by samples that contain an appreciable amount of silicate minerals or are spatially associated with skarns, indicating the progress of decarbonation reactions (Baumgartner & Valley 2001). In addition, external fluid infiltration events along lithological contacts (see Satish-Kumar et al. 2006a) might have also altered the isotopic signatures. Type-III marbles, especially those from the Skallen region, display major oxygen isotope heterogeneity in micro-domains (Satish-Kumar et al. 1998, 2006b). Based on the extensiveness of the negative shift in oxygen isotopic composition

and the small, but consistent, shift in carbon isotopic composition, it can be considered that aqueous fluid of meteoric origin percolated the marble along fractures and grain boundaries, during the brittle regimes of retrograde metamorphism (Fig. 8). Satish-Kumar et al. (1998), based on a microscale isotope zonation study in a sample from the same horizon, suggested that dissolution–reprecipitation and further diffusion resulted in heterogeneities of the order of 20‰. In high-grade metamorphic terranes, it is not uncommon to find fluid infiltration events during retrograde metamorphism (e.g. Buick & Cartwright 1996; Buick et al. 1997; Cartwright et al. 1997). In most cases, deep circulating surface fluids alter oxygen isotopic composition of marbles to a higher extent than those fluids derived from crystallizing magmas (Buick & Cartwright 2003). The Skallen marbles were affected by deep circulating surface fluids, which penetrated through selective grain boundaries and fractures (Satish-Kumar et al. 1998). New C– O isotope results from the marbles investigated from the LHC are included in Figure 8. In particular, the d18O values (13.5 – 14.9‰) of sample 2305E are considerably lower than those of sample 602d (17‰) reported by Satish-Kumar & Wada (2000). It is therefore clear that 2305E has been affected by metamorphic


fluid –rock interaction processes, corroborating the evidence seen in CL imagery and trace element variations. Furthermore, d13C and d18O results from Skallevikshalsen, Breidvagnipa and Kasumi Rock marbles cluster nearer the Neoproterozoic d13C and d18O values (Fig. 8) and retain the possibility of preserving premetamorphic signatures.

Strontium isotope geochemistry Strontium isotopic composition of marbles displays remarkable variations within single marble layers as well as between layers on a regional scale in the LHC. In particular, marble layers from Skallen and Skallevikshalsen have unusually high 87 Sr/86Sr(550 Ma) ratios, whereas some layers of dolomitic marbles exhibit ‘near-normal’ Proterozoic sedimentary Sr isotopic ratios. The highest (0.7582) and lowest (0.7066) 87Sr/86Sr(550 Ma) values were observed in a dolomite marble from Skallevikshalsen and Skallen localities, respectively (Table 1). Moreover, a calcite marble from the Kabuto Rock outcrop has a 87Sr/86Sr(550 Ma) value of 0.7053 (S. Kagashima, pers. comm.). Although not in all cases, the higher 87 Sr/86Sr(550 Ma) ratios can be correlated with a decrease in d13C and d18O values (Type-II and Type-II marbles), which can be attributed to the fluid –rock interaction process as discussed by Satish-Kumar et al. (2006b).

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For example, samples that show high Sr/86Sr(550 Ma) ratios have variations within single hand specimens (0.7250–0.7261 in sample 2305E; Fig. 4a). The d18O values are characteristically low (Fig. 4a), corresponding to the Type-II marble. However, within-sample variations in 87 Sr/86Sr(550 Ma) ratios could not be directly correlated to carbon and oxygen isotopic compositions (Fig. 4a). In contrast, monomineralic marble layers that display low strontium isotopic composition, such as 0.7070 for sample 602d, have relatively homogeneous 87Sr/86Sr(550 Ma) values and characteristically high oxygen and carbon isotope values (Fig. 8), and less scatter in trace element and REE contents (Figs 6 and 7). Therefore, marbles of Type-I have equilibrium homogeneous geochemical characteristics. In addition to the low 87Sr/86Sr(550 Ma) ratios, the Type-I marbles also have higher Sr contents (Fig. 9). However, the variation in the Sr contents could not be directly correlated with a simple mixing of sources such as those derived from the pelitic rocks, but in general, metamorphic fluid –rock interaction processes can be considered as causing the alteration of 87Sr/86Sr(550 Ma) ratios (Satish-Kumar et al. 2006b). 87

Discussion Carbon, oxygen and strontium isotopic compositions of metacarbonate rocks have been widely

Fig. 9. Sr isotopic compositions as a function of Sr contents. The scatter of Sr isotope ratios for different marble samples from the same locality should be noted. The Sr initial isotope ratio for metapelitic rocks from Skallen is after Miyamoto et al. (2008).

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utilized for deducing the fluid –rock interaction processes in the continental crust (e.g. Cartwright & Valley 1991; Richards et al. 1996; Bickle et al. 1997; Buick et al. 1997; Satish-Kumar et al. 1998). Alternatively, several studies have also pointed out the potential application of carbonate or metacarbonate rocks in chemostratigraphic studies (Baker & Fallick 1988; Melezhik et al. 2001a, b, 2002a, b, 2005, 2006), because they preserve pre-metamorphic isotope signatures (Baker & Fallick 1988; Wickham & Peters 1993; Vyhnal & Chamberlain 1996; Melezhik et al. 2005, 2006). In addition to stable and radiogenic isotopes, trace element and REE geochemistry is also useful in identifying metamorphic fluid–rock interaction in marbles (Boulvais et al. 2000; Korsakov & Hermann 2006; Satish-Kumar et al. 2006b). Recent studies have also found that high-grade calcite marbles have a higher geochemical preservation potential than dolomitic marbles (Melezhik et al. 2005). Isotope geochemistry (especially Sr isotopes of seawater) is believed to be largely controlled by the supercontinent cycle, reflecting a balance between the continental input and mantle input to the oceans (Veizer et al. 1992; Condie 2000, 2003). The scarcity of unaltered or unmetamorphosed carbonate sediments in the Precambrian places limitations on our understanding of palaeo-oceans. However, metacarbonate rocks are common in the Proterozoic mobile belts. Apparently, in the absence of interaction with fluids during post-depositional events, high-grade marbles are potential candidates for understanding the palaeo-seawater chemistry and may provide insight into the determination of an age of deposition and thereby be useful in correlation studies. Conventionally, shifts from typical sedimentary signatures of carbon and oxygen isotopic composition are caused by metamorphic decarbonationrelated fluid– rock interaction or exchange with external fluids (Baumgartner & Valley 2001). Pre-metamorphic alterations, such as diagenesis or dolomitization, can also induce isotopic shifts. Decarbonation and associated fluid –rock interaction can occur through batch volatilization or Rayleigh volatilization, and in these cases the C–O isotope fractionation can be traced through modelling, by taking into account the amount of remaining carbonates and final isotopic composition (Baumgartner & Valley 2001, and references therein). However, interaction with external fluids is often complicated and difficult to trace appropriately, unless obvious geological evidence is identified. In addition to stable isotopes, Sr isotopes can also provide valuable information on the source characteristics of carbonate sediments as well as post-depositional processes. Initial Sr isotopic composition of carbonates usually preserves

chemical characteristics of seawater and, if undisturbed, indicates an apparent depositional age. However, Sr isotopic composition is highly sensitive to fluid processes during metamorphism under crustal regimes (e.g. Bickle et al. 1997; Abart et al. 2002) and is commonly used in tracing post-depositional alteration events.

Geochemical signatures of post-depositional alterations Carbonate rocks are highly vulnerable to geochemical alteration after deposition. However, in regional highgrade terranes, alterations during diagenesis and prograde metamorphism are difficult to constrain, because of the geochemical homogenization during high-temperature metamorphism and post-peak metamorphic resetting. In a recent study, Melezhik et al. (2005) evaluated the post depositional isotopic resetting based on major and trace element geochemistry, and d18O, d13C and 87Sr/86Sr ratios in Neoproterozoic dolomitic and calcitic marbles metamorphosed under amphibolite-facies conditions. Whereas both dolomitic and calcitic marbles display a high degree of preservation of premetamorphic carbon and oxygen isotopic composition, Sr isotope ratios of dolomitic marbles were reset during post-depositional alteration. Comparatively, calcitic marbles were more capable of preserving lower 87Sr/86Sr ratios than dolomitic marbles, suggesting an apparent depositional age of 700–600 Ma. The large-scale heterogeneity in Sr isotopic composition of marbles from the LHC is considered to have resulted from the post-depositional alteration of carbonates. Marbles associated with skarns have generally higher Sr contents than pure marbles. Satish-Kumar et al. (2006b) reported a 87 Sr/86Sr(550 Ma) ratio of 0.735 for Cl-rich scapolite and 0.731 for CO3-rich scapolite. Metapelitic rocks associated with the marble layer have a 87 Sr/86Sr(550 Ma) ratio of 0.764 (Miyamoto et al. 2008). These variations could not be directly correlated with simple mixing of two sources such as pelitic and carbonate rocks, although some of the samples seem to follow that trend (Fig. 9). Extensive Sr isotopic variations are observed in pure dolomitic and calcite marbles. In particular, the enrichment of Sr isotopes with concomitant decrease in Sr content is most effective during diagenesis. Further, interaction with fluids released from adjacent pelitic lithologies during prograde metamorphism can also cause large-scale shifts. A simple interaction model between carbonate rocks and diagenetic or prograde metamorphic fluids cannot explain the extensive variation observed in the Sr isotope systematics. More detailed


microscale analyses and assessment of Sr isotopes are necessary for understanding the postmetamorphic processes and will be addressed elsewhere. Here, we tentatively consider the large-scale variations to have occurred in an open system during post-depositional processes (diagenesis þ metamorphic fluid–rock interaction) and focus our attention on the low- 87Sr/86Sr(550 Ma) marbles. In comparison with unmetamorphosed Neoproterozic carbonate sediments (Jacobsen & Kaufman 1999; Melezhik et al. 2005), most of the LHC marbles display a general shift toward lower oxygen isotope values, higher Sr isotope ratios, and variations in trace element and REE characteristics. These characteristics are indicative of substantial geochemical resetting during postdepositional diagenesis and metamorphism. Integrating the observed geochemical features with the evidence obtained from the occurrence of Cl-rich scapolite from Skallen and Skallevikshalsen (Satish-Kumar et al. 2006b), we consider that the marbles were subjected to interaction with 87 Sr-enriched hypersaline fluids. 87Sr enrichment in the hypersaline fluids could result from the interaction with adjacent pelitic lithologies (87Sr/86Sr(550 Ma) ¼ 0.764; Miyamoto et al. 2008) in the early stages of metamorphism. Although the timing of this process cannot be determined, we tentatively assign it to an early stage of prograde metamorphism (,500 8C), coeval with the formation of Cl-rich scapolite (Satish-Kumar et al. 2006a). Furthermore, marbles (Type-III) were also locally affected by retrograde fluid –rock interaction, which resulted in extreme negative shifts in oxygen isotope value. These retrograde fluid– rock interactions were limited to grain boundaries and fractures, and caused substantial geochemical alterations (Satish-Kumar et al. 1998, 2006b).

Preservation of pristine sedimentary geochemical features Despite the extensiveness of post-depositional geochemical alteration in many of the marble layers in the LHC, exceptional marble layers (Type-I) tend to preserve pre-metamorphic geochemical features. The following geochemical signatures are considered to support this assertion. Coarse crystalline monomineralic marbles have (1) typical homogeneous CL images; (2) wellhomogenized trace element compositions and tight REE patterns; (3) similar d13C values to nonmetamorphic Neoproterozoic carbonates; (4) the highest d18O values among the marbles in the LHC, which are comparable with those of the non-metamorphic Neoproterozoic carbonates;

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(5) low Sr isotope ratios and high Sr concentration, comparable with values for sedimentary carbonates. Early stages of chemical alteration during diagenesis and dolomitization of carbonate sediments, in the presence of external fluids, will result in an enrichment of Mn and depletion of 13C and 18O (Gao & Land 1991). In Type-I marbles we could not find any geochemical signatures that are indicative of alteration during diagenesis and/or dolomitization.

Apparent age of deposition The isotopic composition of Sr in the ocean is considered to be homogeneous, because of its long residence time. However, the 87Sr/86Sr ratio in the oceans varies with time and deviates from the mantle evolution curve in the early Proterozoic, because of the input of Sr derived from continents. The increase in 87Sr/86Sr ratio, however, is irregular during Neoproterozoic to early Palaeozoic time, and has been constrained rather precisely based on detailed studies on temporal trends in sedimentary carbonate rocks (Derry et al. 1994; Denison et al. 1998; Jacobsen & Kaufman 1999; Kuznetzow et al. 2003a, b; Melezhik et al. 2006). Although it is difficult to substantiate complete preservation of sedimentary isotopic compositions, an apparent age of deposition can be assigned from the least altered 87Sr/86Sr ratios, by comparing them with the temporal trends in 87Sr/86Sr in seawater prior to the peak metamorphism in the region (Fig. 10). In the SW LHC, from the least altered 87Sr/86Sr ratio of 0.7066, obtained from Skallen, an apparent age of c. 730 Ma could be deduced or c. 640 Ma, if the reference curve of Melezhik et al. (2006) is taken into account. Other less altered samples have 87Sr/86Sr ratios of 0.707, which indicate the possibility of a younger depositional age around 580 Ma; however, this age seems irrelevant, as discussed below. Apparently, the lowest 87Sr/86Sr ratio obtained from LHC rock is 0.7053, which might suggest carbonate deposition in the region as old as c. 830 Ma. Alternatively, these strontium isotope ratios might be preserving an event of interaction of earlier deposited carbonate sediments with seawater of those ages, which cannot be ruled out. A recent compilation of Neoproterozoic carbon isotope data of unmetamorphosed carbonate rocks in the Neoproterozoic by Halverson et al. (2005) suggests that the d13C values were positive (up to 5‰) before 670 Ma, except for short glaciation intervals. The highest d13C values of carbonates from the LHC mostly group around 0‰, with exceptional layers preserving positive values up to 2‰. The reference chemostratigraphic curve of Halverson et al. (2005) cannot be directly compared

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Fig. 10. Extrapolation of initial Sr isotopic composition for a possible depositional age of carbonate rocks in the marginal sea between East and West Gondwana. The temporal trend of 87Sr/86Sr in seawater is after Melezhik et al. 2006 and references therein.

with the carbon isotope values of the marbles in the present study, as most of the marbles contain graphite. The presence of graphite indicates that the original carbon isotope values of calcite were re-equilibrated with organic carbon having low d13C during metamorphism. Alternatively, the d13C values of carbonates now preserve the minimum values and do not contradict the depositional ages estimated using Sr isotope ratios. With the limitations described above, we tentatively assign the age spectrum of deposition of carbonate rocks in the LHC to be between 730 and 830 Ma. Further detailed geochemical and isotopic study of carbonate rocks is necessary to confirm and refine this age spectrum. The age of sedimentation obtained from the carbonate rocks is in broad agreement with the existing geochronological and tectonic framework of the LHC. Detailed U– Pb sensitive high-resolution ion microprobe (SHRIMP) studies on metamorphic gneisses in this region have suggested that the age of peak metamorphism is c. 530 Ma (Shiraishi et al. 1994, 2003), although earlier ages of around 650 –580 Ma have also been reported recently (Hokada & Motoyoshi 2006). The early metamorphic ages suggest that the deposition of carbonate rocks could not have occurred as late as 580 Ma, but instead might represent an alteration event during metamorphism. The age of sedimentation suggested in this study is also supported by the youngest reported inherited ages of magmatic zircons in metapelitic rocks in the LHC region. Shiraishi et al. (2003) reported

c. 1000 Ma age zircon cores from several localities in the LHC. They considered the protolith of the pelitic gneisses as detritus deposited at the margin of a pre-Grenvillian craton, which consists of various continental components of ages older than 1000 Ma. Moreover, Shiraishi et al. (2008) reported a youngest TDM age of 0.87 Ga and a major mode at 1.0–1.25 Ga for samples from the LHC. This provides an upper age limit for sedimentation in the region. The results of our study supplement this assertion that the deposition of sediments in the ‘Mozambique Ocean’ between East and West Gondwana might have occurred between 730 and 830 Ma. Carbonate rocks preserve important evidence on the marginal sea conditions of supercontinent assembly. In particular, the Sr isotope record of carbonate can give insights into the palaeo-ocean chemistry. A sharp decrease in Sr isotope ratio during the early Neoproterozoic (c. 850 Ma) is evident in many recent studies of unmetamorphosed carbonate sediments (see Melezhik et al. 2006, and references therein). Low Sr isotope ratios are indicative of enhanced mantle input of Sr to the seawater, which potentially relates to the breakup of supercontinents. In contrast, a steady increase in 87 Sr/86Sr ratio in the seawater might result from the supply of Sr during extensive continental erosion, following continental uplift during collision. For example, the steep increase in the Sr isotope ratio between 600 and 500 Ma corresponds to possible collision between East and West


Gondwana to form the Gondwana supercontinent (Condie 2000). Thus, strontium isotopes serve as an indirect proxy in understanding the supercontinent cycle. Thick metasedimentary sequences exposed in East Antarctica include metacarbonate rocks supposed to have been deposited in the ‘Mozambique Ocean’ that separated East and West Gondwana (Hoffman 1988, 1991; Dalziel 1991; Stern 1994). Similar to the results on high-grade marbles that we have presented here, geochemical studies may provide insight into the seawater geochemistry and may serve as valuable indicators for supercontinent evolution.

Conclusions (1) Metacarbonate rocks can provide important clues for understanding the palaeo-oceans that existed between proto-continents. (2) Geochemical studies of metacarbonate rocks from the Lu¨tzow-Holm Complex, East Antarctica, suggest that most of the metacarbonate rocks have been altered during protracted postdepositional tectonic events. (3) Most of the marbles display extreme variations in carbon, oxygen and strontium isotopic composition, suggesting an active role of fluids in the Lu¨tzow-Holm Complex continental crust. (4) Exceptional metacarbonates, which preserve possible near-depositional geochemical features, are found in the Lu¨tzow-Holm Complex. These metacarbonates constrain a possible apparent depositional age between 730 and 830 Ma. (5) There is need for more information on metacarbonates from adjacent Gondwana terranes to understand the Neoproterozoic evolution of the ‘Mozambique Ocean’ that separated East and West Gondwana. M.S.-K. acknowledges a grant (No. 18740319) from the Ministry of Education, Culture, Sports, Science and Technology, Japan. We thank the geological field survey group and the crew members of the Japanese Antarctic Research Expedition for their support during the field expedition. T. Tsuchiya (Tohoku University) and S. Kagashima (Yamagata University) are thanked for sparing marble samples from Breidvagnipa and Kabuto Rock, respectively. We thank V. Melezhik and an anonymous reviewer for constructive comments, and K. Shiraishi for editorial handling of our manuscript.

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