Developments in Precambrian Geology 2 PRECAMBRIAN OF THE SOUTHERN HEMISPHERE
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Developments in Precambrian Geology 2 PRECAMBRIAN OF THE SOUTHERN HEMISPHERE
DEVELOPMENTS I N PRECAMBRIAN GEOLOGY Advisory Editor
B.F. Windley
Further t i t l e s in this series
1.
B.F. WINDLEY and S.M. NAQVl (Editors) Archaean Geochemistry
3. K.C. CONDIE Archean Greenstone Belts
DEVELOPMENTSIN PRECAMBRIAN GEOLOGY 2
PRECAMBRIANOF THE SOUl'HmN HEMISPHERE Edited by
D.R. HUNTER Department of Geology, University of Natal, Pietermaritzburg, South Africa
E LSEVl E R SCI E N T l F IC PUB LlSH I N G COMPANY Amsterdam - Oxford - New York 1981
ELSEVIER SCIENTIFIC PUBLISHING COMPANY 1, Molenwerf 1014 AG Amsterdam P.O. Box 21 1, Amsterdam, The Netherlands Distributors for the United States and Canada:
ELSEVIER NORTH-HOLLAND INC. 52, Vanderbilt Avenue New York, N.Y. 10017
Library of Congrr\5 Cataloging in Publication Data
Main e n t r y u n d e r t i t l e : b e c m b r i a n of t h e S o u t h e r n Hemisphere. (Developments i n Precambrian geology ; 2 ) I n c l u d e s b i b l i o g r a p h i e s and i n d e x . 1. Geology, S t r a t i g r a p h i c- - P r e - C m b r i an. 2. Geology--Southern Hemisphere. I. H u n t e r , Donald Raymond. 11. S e r i e s .
QE653.P735 551.7'1'091814 ISBN 0-444-41862-8
80-16064
ISBN 0-444-41862-8 (Val. 2) ISBN 0-444-41719-2 (Series) 0 Elsevier Scientific Publishing Company, 1981 A l l rights reserved. N o part of this publication may be reproduced, stored in a retrieval system or transmitted i n any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without the prior written permission of the publisher, Elsevier Scientific Publishing Company, P.O. Box 330,Amsterdam, The Netherlands
Printed in The Netherlands
V
INTRODUCTION
In the preparation of this volume, the various contributors have attempted to synthesize the very considerable mass of data that has been and is still being accumulated on the Precambrian rocks of the southern continents. The details of the geology are frequently complex and not yet susceptible t o unequivocal interpretation. Furthermore, outcrop quality differs greatly from, for example, the equatorial rain forests of South America to the more arid regions of Australia and Southern Africa. The consequence is inevitably that some areas are better understood than others. The authors have perforce had to present their data in a succinct form so that the volume would be of manageable proportions. Within these constraints, this volume therefore summarizes the main geological features of the three continents. However, the authors have provided extensive bibliographies for those readers who wish t o obtain more information on specific topics. Readers may well be surprised that the great continent of Antarctica has been excluded from this volume. Antarctica forms a critical building block in the reconstruction of Gondwanaland as Fig. 1.1 demonstrates. Knowledge of its geology is fundamentally important. At the present time, geological data are being accumulated by many workers from several nations. The difficulties involved in working in Antarctica together with the multi-national nature of this research makes it difficult t o prepare a synthesis of the geology. Although knowledge of the Phanerozoic rocks of Antarctica is fairly extensive, the Precambrian is less well-documented and it would be perhaps premature at this time t o make such an attempt. The geology of each continent has been discussed insofar as is possible in terms of provinces, following introductory chapters on the broad structural framework of each continent. Accounts of the granitoid shields with their remnants of supracrustal rocks are followed by descriptions of the nature of the lithic fill of successive sedimentary basins developed on these shields, and of subsequent plutonic, metamorphic and deformational events. Each section concludes with a review of the palaeontological record during the Precambrian . Figure 1.1 illustrates in a simplified and schematic form the broad relationships between the disposition of orogens and zones of tectonic reworking in
VI the Gondwanan reconstruction of Smith and Hallam (1970). In view of the parallels drawn by Goode in this volume between the Proterozoic evolution of Western Australia and southern Africa, it seems probable that there was a closer spatial relationship between these continents than is suggested by the Smith-Hallam (1970) model. There are a number of striking similarities between, for example, the Early Proterozoic Hammersley Basin of Australia and the Transvaal Basin of Southern Africa. Button (1976) suggests that the correspondence in development of these basins can be best explained if both developed on a common continental block, subject to the same tectonic, climatic and atmospheric history. Button (1976) envisages the closer juxtaposition of Australia and Africa as in the Gondwanan reconstruction of Piper (1974), with the modification that Madagascar is left in its present position. Such a reconstruction would also bring the -1100 Ma Namaqua-Natal orogenic zone of Southern Africa into alignment with those of apparently similar age in Australia. However, the relationships of more recent orogens between South America and Antarctica would be less clear. These problems must await geologic and geochronologic data from the east Antarctica shield. As several authors wish t o acknowledge assistance from colleagues, the editor has taken the liberty of collating these into a consolidated list which follows this introduction. The editor would also like to thank all contributors for their cooperation in the compilation of this volume.
Fig. 1.1. Reconstruction of Gondwandand.
VII Since the manuscripts were received, Rhodesia has been renamed Zimbabwe. Alterations have been made in the text, but it was not possible to amend the text-figures, in which references to Rhodesia should now read Zimbabwe. The term Rhodesian craton has been retained as a formal geological name because of its established place in geological literature. D.R. HUNTER Pietermaritz burg October 1979 REFERENCES Button, A., 1976. Transvaal and Hammersley Basins - a review of basin development and mineral deposits. Miner. Sci. Eng., 8 : 262-293. Piper, J.D., 1974. Proterozoic crustal distribution, mobile belts and apparent polar movements. Nature, 251: 331-384. Smith, A.G. and Hallam, A., 1970. The fit of the southern continents. Nature, 225: 139.144.
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IX
ACKNOWLEDGEMENTS
The following authors wish t o record their appreciation for assistance received in preparing their contributions: (1)C.R. Anhaeusser wishes t o thank C.F. Vermaak and A. Bernasconi for their help in providing data on the Precambrian geology of Angola. ( 2 ) A . Button wishes to acknowledge the helpful comments and assistance received from H. de la R. Winter, N.J. Grobler, J.N.J. Visser, N.J. Beukes, and K.A. Eriksson. ( 3 ) K . Bloomfield wishes t o acknowledge the assistance received from R.N. Annels, E.A. O’Connor and B.A. Klinck who reviewed his contribution. (4)A.D. T. Goode is deeply grateful to colleagues in The Broken Hill Proprietary Co Ltd (BHP), particularly W.D.M. Hall, for their assistance and encouragement in preparing his contribution. Many were also involved in various investigations in recent years which helped to form the basis of his contribution; these include C.D. Arndt, S.V. Bell, G. Davis, W.D.M. Hall, J.E. Harms, J.S. MacFarlane, H.S. Madan, H.C. Meyer, and I.W. Reid. Geologists from other organizations who were most helpful include A.T. Brakel, J.A. Bunting, R.D. Gee, P.C. Muhling, A.F. Trendall and I.R. Williams (Geological Survey of Western Australia), M.R. Walter (Bureau of Mineral Resources), and R.C. Horwitz (C.S.I.R.O.). The assistance of G. Beavis, J. Bradley and S. Varley in the preparation of the manuscript is greatly appreciated. (5) M.F. Glaessner and M.R. Walter wish t o thank J.H. Oehler and M.D. Muir for their helpful suggestions for improvement of the manuscript of their contribution. M.R. Walter publishes with permission of the Director, Bureau of Mineral Resources, Geology and Geophysics, Australia. ( 6 )J. Hallberg and A . Y. Glikson wish t o acknowledge the helpful comments on their contribution from R.C. Honvitz, R.D. Gee, A.H. Hickman, and W.E. Ewers. They also wish to thank C.R. Steel for drafting the figures, and C.K. Harris for typing the manuscript. ( 7 )K . A . Plumb, G.A. Derrick, R.S. Needham and R.D. Shaw wish to record that their contribution is the result of the cumulative efforts of numerous BMR colleagues, both past and present, who have carried out the systematic regional mapping of northern Australia which has provided the data and con-
X cepts on which their synthesis is based. Particular thanks are due t o those involved in current BMR projects and who have freely discussed their ideas and provided both published and unpublished data for their contribution: A.J. Stewart, L.A. Offe and R.G. Warren (Arunta Inlier); A.W. Wells and M.R. Walter (central Australian basins); J.H.C. Bain, B.S. Oversby, D.E. Mackenzie, and I.W. Withnall (GSQ) (Georgetown Inlier); M.J. Jackson and M.D. Muir (McArthur Basin), I.H. Wilson (GSQ), D.H. Blake, R.J. Bultitude, R.M. Hill, I.P. Sweet (Mount Isa); I.H. Crick and P. Stuart-Smith (Pine Creek Inlier); L.P. Black and R.W. Page (Geochronology). The manuscript has been reviewed by J.H.C. Bain, D.H. Blake, I.P. Sweet, R.G. Warren and A.W. Wells. ( 8 ) J.F. Wilson wishes t o acknowledge the considerable help he received from the Geological Survey of Zimbabwe in preparing his contribution.
XI
LIST OF CONTRIBUTORS
AUSTRALIA
G.M. Derrick, Bureau of Mineral Resources, Geology and Geophysics, P.O. Box 378, Canberra, A.C.T. 2601, Australia. M.F. Glaessner, Department of Geology and Mineralogy, University of Adelaide, Adelaide, South Australia 5001, Australia. A.Y. Glikson, Bureau of Mineral Resources, Geology and Geophysics, P.O. Box 378, Canberra, A.C.T. 2601, Australia. A.D.T. Goode, Exploration Department, The Broken Hill Proprietary Co Ltd., Melbourne Research Laboratories, Clayton, Victoria 3168, Australia. J.A. Hallberg, Division of Mineralogy, C.S.I.R.O., Wembley, Western Australia 6014, Australia. B. Murrel, Department of Geology and Mineralogy, University of Adelaide, Adelaide, South Australia 5001, Australia. R.S. Needham, Bureau of Mineral Resources, Geology and Geophysics, P.O. Box 378, Canberra, A.C.T. 2601, Australia. A.J. Parker, Department of Geology and Mineralogy, University of Adelaide, Adelaide, South Australia 5001, Australia. G.M. Pitt, Department of Geology and Mineralogy, University of Adelaide, Adelaide, South Australia 5001, Australia. K.A. Plumb, Bureau of Mineral Resources, Geology and Geophysics, P.O. Box 378, Canberra, A.C.T. 2601, Australia. W.V. Preiss, Department of Geology and Mineralogy, University of Adelaide, Adelaide, South Australia 5001, Australia. R.W.R. Rutland, Department of Geology and Mineralogy, University of Adelaide, Adelaide, South Australia 5001, Australia. R.D. Shaw, Bureau of Mineral Resources, Geology and Geophysics, P.O. Box 378, Canberra, A.C.T. 2601, Australia. M.R. Walter, Bureau of Mineral Resources, Geology and Geophysics, P.O. Box 378, Canberra, A.C.T. 2601, Australia. SOUTHERN AFRICA
C.R. Anhaeusser, Economic Geology Research Unit, University of the Witwatersrand, Jan Smuts Avenue, Johannesburg 2001, South Africa.
XI1
K. Bloomfield, Institute of Geological Sciences (Overseas Division), 154 Clerkenwell Road, London EClR 5DU, England. T.J. Broderick, Geological Survey, P.O. Box 8039, Causeway, Salisbury, Zimbabwe. A. Button, Department of Geology and Geological Engineering, South Dakota School of Mines and Technology, Rapid City, S.D. 57701, U.S.A. M.C. Du Toit, Department of Geology, Rand Afrikaans University, P.O. Box 524, Johannesburg 2000, South Africa. K.A. Eriksson, University of Texas at Dallas, Box 688, Richardson, Texas 75080, U.S.A. D.R. Hunter, Department of Geology and Mineralogy, University of Natal, P.O. Box 375, Pietermaritzburg 3200, South Africa. H. Jansen, Geological Survey, Private Bag X112, Pretoria, 0001, South Africa. P. Joubert, Precambrian Research Unit, University of Cape Town, Private Bag Rondebosch, Cape Town 7700, South Africa. C.A. Lee, Geology Department, Johannesburg Consolidated Investment Co. Ltd., P.O. Box 976, Randfontein 1760, South Africa. N.H. Lockett, Hunting Geology and Geophysics, P.O. Box 363, Eyswick, A.C.T. 2609, Australia. I.M. McGregor, Bernard Price Institute for Palaeontological Research, University of the Witwatersrand, Jan Smuts Avenue, Johannesburg 2001, South Africa. R. Mason, Department of Geology, Rhodes University, Grahamstown 6140, South Africa. P. Matthews, Department of Geology, University of Natal, King George V Avenue, Durban 4001, South Africa. F. Mendelsohn, 643 Maritime House, 26 Loveday Street, Johannesburg 2001, South Africa. D.A. Pretorius, Economic Geology Research Unit, University of the Witwatersrand, Jan Smuts Avenue, Johannesburg 2001, South Africa. I.D.M. Robertson, c/o CRA Exploration. P.O. Box 9, Cobar, N.S.W. 2835, Australia. J.G. Stagman, Geological Survey, P.O. Box 8039, Causeway, Salisbury, Zimbabwe. C.F. Vermaak, J.C.I. Exploration, Research and Development Division, James Park, P.O. Box 976, Randfontein 1760, South Africa. A.H. Wilson, Department of Geology and Mineralogy, University of Natal, P.O. Box 375, Pietermaritzburg 3200, South Africa. J.F. Wilson, Department of Geology, University of Zimbabwe (Rhodesia), P.O. Box M.P. 167, Mount Pleasant, Salisbury, Zimbabwe. SOUTH AMERICA
E. Suszczynski, Rua Joana Angelica 5/203, Ipanema, Rio de Janeiro, G.B., Brasil.
XI11
CONTENTS
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . List of Contributors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Contents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
V IX XI xi11
A . AUSTRALIA
Chapter I .
STRUCTURAL FRAMEWORK O F THE AUSTRALIAN PRECAMBRIAN R.W.R. Rutland
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Basement provinces and their subdivision . . . . . . . . . . . . . . . . . . . . . . . . . . . . Principles for recognition of provinces . . . . . . . . . . . . . . . . . . . . . . . . . . . . The chelogenic cycle and plate tectonics . . . . . . . . . . . . . . . . . . . . . . . . . . . The Archaean chelogen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Proterozoic chelogen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Musgrave-Fraser Province . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The block structure of Australia: basement reactivation . . . . . . . . . . . . . . . . . . . : . . :. . :. . :. . : The Archaean blocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Amadeus Transverse Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Kimberley Block . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Gawler Block . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Central Australian Platform Cover . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Trans-Australian Platform Cover . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Note added in proof . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
Chapter 2 .
1 3 3 9 12 16 19 19 20 20 25 26 26 27 27 29 32
ARCHAEAN GRANITE42REENSTONE TERRANES O F WESTERN AUSTRALIA J.A. Hallberg and A.Y. Glikson
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Yilgarn Block . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Subdivision . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy and rock associations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Lower mafic/ultramafic associations - 43 ; Lower sedimentary association 49 ; Upper maficlultramafic association - 51 ; Upper felsic volcanicsedimentary association - 52; Intrusive rocks - 54 Structure and metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Granitoids . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Economic geology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
33 36 36 37
55 58 62
XIV Evolution of the Yilgarn Block . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pilbara Block . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy of the greenstone belts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure and metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Granitoids . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geochronology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mineralization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Evolution of the Pilbara craton . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pilbara-Yilgarn relationships . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chupter 3 .
66 73 73 81 85 88 90 91 94 96
PROTEROZOIC GEOLOGY OF WESTERN AUSTRALIA A.D.T. Goode
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 105 Regional tectonic units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 107 Archaean granite-greenstone terrains . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 108 Intracratonic Proterozoic basins and fold belts . . . . . . . . . . . . . . . . . . . . . . . 108 109 Proterozoic metamorphic complexes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Intracratonic basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109 Early Proterozoic Hamersley Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 109 Fortescue Group . 111; Hamersley Group . 113; Turee Creek Group . 119; Wyloo Group . 120; Manganese Group . 122; Structure and metamor124; Geochronology . 126; History of sedimentation and tectophism . 127 ; Ore deposits . 130 nism . 134 Early Proterozoic Nabberu Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Earaheedy Sub-basin . 134; Glengarry Sub-basin . 143; Padbury Sub-basin . 144; Relationships between the sub-basins . 145; Structure and metamor145; Geochronology . 148; History of sedimentation and tectophism . 149; Ore deposits . 152 nism . Relationships between the Hamersley and Nabberu Basins . . . . . . . . . . . . . . . 152 153; Fold belts . 153; Depositional model for the iron forSedimentation . 154 mations . Middle Proterozoic Mt . Minnie and Bresnahan Basins . . . . . . . . . . . . . . . . . . . 156 Bresnahan Basin . 156; M t . Minnie Basin . 157; Trainor Inlier . 157; Scor157 pion Group . Middle Proterozoic Bangemall Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 158 Geochronology . 158; Western Bangemall Basin . 158; Eastern Bangemall 170; Relationships between the Yeneena and Bangemall Groups . Basin . 175; Regional correlations . 175; Ore deposits . 177 Minor sedimentary sequences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 177 Badgeradda Group and Nilling Beds . 177; Wenmillia Formation . 177; Yan177; Billerange Group and Dudawa Beds . 178; Moora danooka Group . 178; Wood178; Cardup Group . 178; “Stirling-Barren Series” . Group . 179; Officer Basin . 179 line Beds . Metamorphic complexes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 179 Albany-Fraser Province . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 179 Western Gneiss and Granite Zone . 181; Fraser Complex . 181; Eastern 182; Eastern Granite Zone . 182 Gneiss and Migmatite Zone . Gascoyne Province . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 183 184 Leeuwin Block . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 184 Northampton Block . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
XV Mullingarra Inlier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Rudall Inlier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Early Proterozoic dolerite dyke suites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Widgiemooltha Dyke Suite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Black Range Dyke Suite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Carbonatites and alkaline rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Proterozoic crustal structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chapter 4 .
185 185 186 186 188 188 189 193
THE PROTEROZOIC OF NORTHERN AUSTRALIA K.A. Plumb. G.M. Derrick. R.S. Needham and R.D. Shaw
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Isotopic ages . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Time subdivision . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tectonic terminology and definitions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Synthesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Principal tectonic units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Summary descriptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . West Australian Orogenic Province (Archaean basement inliers) . . . . . . . . . . . . North Australian Orogenic Province (Early Proterozoic orogenic domains) . . . . North Australian Platform Cover (Early to Mid-Proterozoic platform covers) . . . Central Australian Orogenic Province (Early to Mid-Proterozoic orogenic domains) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Central Australian Platform Cover (Late Proterozoic (Adelaidean) platform covers) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Regional structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mineral deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Early Proterozoic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mid-Proterozoic . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Regional Geology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Archaean basement inliers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Early Proterozoic orgenic domains . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . McClintock Subprovince . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Halls Creek Inlier - 221; Other domains - 226 Palmerston Subprovince . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pine Creek Inlier - 227; Other domains - 234 Early to Mid-Proterozoic platform covers . . . . . . . . . . . . . . . . . . . . . . . . . . . . Kimberley Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy - 235; Palaeogeographic evolution - 235; Hart Dolerite - 238; Structure - 238 McArthur Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tectonic setting - 239; Stratigraphy - 242; Age - 242; Tectonic elements 245; Palaeogeographic evolution - 245; Dolerite sills - 248; Structure 248; Mineral deposits - 249 Lawn Hill Platform . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Birrindudu Basin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Age - 251; Palaeogeography - 251; Structure - 251 Victoria River Basin and equivalents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphic relationships - 253; Age - 253; Palaeogeography - 255; Structure - 256 Early to Mid-Proterozoic orogenic domains . . . . . . . . . . . . . . . . . . . . . . . . . . .
205 206 208 208 209 209 210 210 211 213 214 216 218 219 219 219 220 220 221 221 227 235 235
239
250 250 251.
256
Mount Isa Orogen-Lawn Hi21 Platform . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tectonic setting . 260; Stratigraphy . 261; Age and correlations . 262; Tec263; Palaeogeographic evolution . 265; Early Proterozoic tonic elements . 265; Late Early Proterozoic-Carpentarian covers . 266; South basement . 269; Igneous, metamorphic and structural style . 269; Nicholson Basin . 270 Mineral deposits . Northeast Queensland Precambrian Province . . . . . . . . . . . . . . . . . . . . . . . . Arunta Inlier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 274; Stratigraphy . 276; Tectonic zones . 278; DescripTectonic setting . 279; Metamorphism . 281; Megastructure . 281; Tectonic tion of zones . models . 282 Late Proterozoic platform covers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Amadeus, Ngalia and Georgina Basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tectonic setting . 286; Stratigraphy . 286; Correlations . 286; Age . 287; 287 ; Structure . 288 Palaeogeography . Kimberley glacial successions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 289; Glaciated pavements and palaeotopography . 293; CorStratigraphy . relations . 293; Age of the sequences . 294; Palaeogeography . 294 , . . . . . . . . . Conclusion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chapter 5 .
273 274
283 283 289
295 297
THE PRECAMBRIAN O F SOUTH AUSTRALIA R.W.R. Rutland. A.J. Parker. G.M. Pitt, W.V. Preiss and B . Murrell
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Gawler Province (R.W.R. Rutland. A.J. Parker and G.M. Pitt) . . . . . . . . . . . . Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Archaean basement rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Early Proterozoic sequences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Deformation and metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Late tectonic history of the Gawler Province . . . . . . . . . . . . . . . . . . . . . . . . Post-1400 Ma evolution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Stuart Shelf and Adelaide Geosyncline (W.V. Preiss. R.W.R. Rutland and B . Murrell) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tectonic framework . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Principal problems . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphic systematics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . History of sedimentation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Willouran - 335; Torrensian - 336; Sturtian - 338; Marinoan - 340; Summary of history of sedimentation - 344 Tectono-stratigraphic discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Delamerian Orogeny . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Timing of deformation - 349; Style of deformation - 350 References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chapter 6.
256
309 309 309 314 316 322 324 326 327 327 330 331 335 345 349 354
AUSTRALIAN PRECAMBRIAN PALAEOBIOLOGY M.F. Glaessner and M.R. Walter
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Archaean (>2500 Ma) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Proterozoic (2500-570 Ma) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . A . 2500-1770 Ma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
361 362 363 363
XVII B . 1170-1400Ma ......................................... .......................................... C.1400-950Ma D.950-750Ma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . E.750-570Ma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
367 375 377 382 389 392
B. SOUTHERN AFRICA Chapter 7.
STRUCTURAL FRAMEWORK D.R. Hunter and D.A. Pretorius
A . Structural province (D.R. Hunter) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cratonic nuclei . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mobile belts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . R . Analysis of gravity data (D.A. Pretorius) . . . . . . . . . . . . . . . . . . . . . . . . . . . The regional gravity field over South Africa and South West Africa . . . . . . . . . The regional gravity field over Southern and Central Africa . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chapter 8.
397 397 405 411 411 417 419
THE GRANITIC-GNEISS GREENSTONE SHIELD C.R. Anhaeusser and J.F. Wilson
423 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 424 A . South Africa (C.R. Anhaeusser) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 424 The gneiss terrane . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 430 Greenstone belts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 432 Barberton Mountain Land . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy - 432; Structure - 435; Metamorphism - 436; Geochronol436; Mineralization . 437 ogy . 438 Murchison greenstone belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy . 438; Structure . 440; Metamorphism . 441; Geochronology . 441 ; Mineralization . 441 442 Pietersburg greenstone belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy . 442; Structure . 444; Metamorphism . 445; Mineralization . 445 445 Sutherland greenstone belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy . 447; Structure . 447; Metamorphism . 449; Mineralization . 449; Geochronology . 449 449 Other granite-reenstone occurrences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 451 Granitic events post-dating the greenstone belts . . . . . . . . . . . . . . . . . . . . . . . . B. Zimbabwe (J.F. Wilson) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 454 Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 454 457 The c . 3500 Ma terrane . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The granites and gneisses . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 457 460 The Sebakwian Group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 463 The c . 2900 Ma terrane . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Bulawayan Group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 463 Belingwe greenstone belt - 464 ; Lower Greenstone - regional considerations - 465; Upper Greenstones - regional considerations - 466 470 The Shamvaian Group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 471 Late granites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 472 Sesombi Suite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
XVIII Chilimanzi Suite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Layered intrusions and dyke swarms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mineralization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Problemsandmodels . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . C . Angola (C.R. Anhaeusser) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Early Precambrian (-3300-1700 Ma) . . . . . . . . . . . . . . . . . . . . . . . . . . . The Middle Precambrian (-1700-850 Ma) . . . . . . . . . . . . . . . . . . . . . . . . . . . The Late Precambrian (-850-500 Ma) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mineralization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Part A - 492; Part B - 495; Part C - 499 Chapter 9 .
473 473 475 479 480 482 488 488 488 490 490 491 492
THE CRATONIC ENVIRONMENT A . Button. D.A. Pretorius. H . Jansen. V . Stocklmayer. D.R. Hunter. J.F. Wilson. A.H. Wilson. C.F. Vermaak. C.A. Lee and J.G. Stagman
A . The older sedimentary-volcanic basins The Pongola Group (A . Button) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Distribution, thickness and areal extent . . . . . . . . . . . . . . . . . . . . . . . . . . . Age . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . (1) The pre-Insuzi basement - 505; (2) The Insuzi Subgroup - 505; The Mozaan Subgroup - 508 Post-depositional history . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mineral deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Witwatersrand Supergroup (D.A. Pretorius) . . . . . . . . . . . . . . . . . . . . . . . . Distribution. thickness and areal extent . . . . . . . . . . . . . . . . . . . . . . . . . . . Age . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . (1) Pre-Witwatersrand basement - 515; ( 2 ) General stratigraphy - 515; (3) Dominion Group-516; (4) West Rand Group-516; (5) Central Rand Group - 517 Post-depositional history . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mineral deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Ventersdorp Supergroup (A . Button) . . . . . . . . . . . . . . . . . . . . . . . . . . . . Distribution. thickness and areal extent . . . . . . . . . . . . . . . . . . . . . . . . . . . Age . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ( 1 ) The Klipriviersberg Group - 522; (2) The Platberg Group - 525; (3) The Pniel Group - 526 Post-depositional history . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mineral deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Transvaal Supergroup ( A.Button) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Distribution. thickness and areal extent . . . . . . . . . . . . . . . . . . . . . . . . . . . Age . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
501 501 501 503 503
510 510 511 511 513 513 515
518 519 520 520 520 522 522
527 527 527 527 529 529
XIX 531 Stratigraphy and conditions of deposition . . . . . . . . . . . . . . . . . . . . . . . . . . ( 1 ) Basal volcanic and elastic unit . 531; (2) Chemical sedimentary unit . 531; ( 3 ) Upper elastic unit . 535 536 Post-depositional history . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 536 Mineral deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . B . The younger sedimentary-volcanic basins 536 The Waterberg and Soutpansberg Groups (H. Jansen) . . . . . . . . . . . . . . . . . . . . 541 Stratigraphy of the Waterberg Group . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 543 Stratigraphy of the Soutpansberg Group . . . . . . . . . . . . . . . . . . . . . . . . . . . 543 The structural pattern of the basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 545 Volcanicity in the basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 547 Palaeogeography of the basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 555 Geodynamics of the basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 556 The Umkondo Group ( V. Stocklmayer) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 559 The Zimbabwe Facies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 559 The Moqambique Facies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 560 Deposition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 560 Structure and metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 561 Mineralization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . C . Plutonic events in the cratonic areas 562 Granitic events (D.R. Hunter) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 562 -2700 Ma granites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 564 -2250-2600 Ma granites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 565 Post-Waterberg granites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 565 Geochemistry and genesis of the granites . . . . . . . . . . . . . . . . . . . . . . . . . . . 567 Synthesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 568 Mafic layered complexes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 568 Usushwana Complex (D.R. Hunter) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 570 Mashaba Igneous Complex (J.F. Wilson) . . . . . . . . . . . . . . . . . . . . . . . . . . . 572 The Great “Dyke” (A.H. Wilson and J.F. Wilson) . . . . . . . . . . . . . . . . . . . . . 578 Kunene Anorthosite Complex (C.F. Vermaak) . . . . . . . . . . . . . . . . . . . . . . . Bushveld and Kindred Complexes (C.F. Vermaak and C.A. Lee) . . . . . . . . . . . 599 The Bushveld mafic sequence - 603; The Bushveld acid phase - 611; Kindred complexes - 612; Tectonic setting - 614; The pre-Bushveld history 614; The Bushveld disruptive cycle (2500-1300 Ma) - 616 D . Sedimentary sequences of uncertain correlation The Piriwiri, Deweras and Lomagundi Groups (J.G. Stagman) . . . . . . . . . . . . . . . 618 627 References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Part A - 627; Part B - 630; Part C - 631; Part D - 639 Chapter 1 0.
MOBILE BELTS I.D.M. Robertson, M.C. du Toit. P . Joubert. P.E. Matthews. N.H. Lockett. F . Mendelsohn. T.J. Broderick. K . Bloomfield and R . Mason
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A The Limpopo Belt (I.D.M. Robertson and M.C. du Toit) Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The North Marginal Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Craton margin and marginal granites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Amphibolite-granulite terrane . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mafic-ultramafic complexes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure and metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
641 642 642 647 649 650
xx The Tuli-Sabi Shear Belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Central Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ancient basement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Limpopo G r o u p . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Anorthositic complexes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Bulai “Granite” . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Singelele “Granite” . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure and metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ( a ) Pikwe . 659; ( b ) Beitbridge . 659; Mahalapye . 660 T h e South Marginal Zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Carton margin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Bandelierkop Formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure and metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Young unmetamorphosed cover and igneous complexes . . . . . . . . . . . . . . . . . . . Economic geology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Syngenetic deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Epigenetic deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Synthesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . B . ‘l’he Namaqualand Metamorphic Complex ( P. Joubert) Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Lithology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geological sequences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Namaqualand - 673; Eastern Bushmanland - 674; Vioolsdrif area - 675; Namibia - 677; Craton-mobile belt junction - 678 Intrusive rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Granites - 683; Charnockitic adamellites - 684; Mafic intrusives - 686; Dyke swarms - 686; Pegmatites - 687; Kimberlites and melilite basalts 688 Younger formations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . T h e 1200 Ma-old formations - 688; T h e Gariep and Damara Groups - 689; Tha Nama Group - 689; Karoo rocks - 690 Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structural events . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Metamorphism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Western Namaqualand . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Namibia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Bushmanland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Metamorphism associated with t h e 1000 Ma event . . . . . . . . . . . . . . . . . . . . Timing of deformational and metamorphic events . . . . . . . . . . . . . . . . . . . . . . . Mineral resources . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Summary of events . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . C . Eastern or Natal sector of the Namaqua-Natal mobile belt in Southern Africa (P.E. Matthews) Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Northern frontal zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Major geological features of t h e northern foreland . . . . . . . . . . . . . . . . . . . . Structural framework of t h e frontal zone . . . . . . . . . . . . . . . . . . . . . . . . . . Natal thrust-belt - 708; Natal nappe complex - 710 Metamorphism and tectonics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Migmatite and granite-gneiss zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Granitic zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Southern granulite zone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
650 652 654 656 657 658 658 659 660 661 661 661 663 663 663 664 664 671 672 672
683
688
690 690 694 694 695 696 696 699 702 703
705 707 707 708 711 712 712 713
XXI Disccusion and conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . D . The western part of t h e Zambezi Metamorphic Belt (N.H. Lockett) . . . . . . . . . The Dett-Kamativi inlier . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Malaputese Formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The linear belts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The granitic rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Economic geology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Wolframitc? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Copper . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mic~i. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Lead . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fluorspar . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geochronology and correlation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . E . Precambrian geology of Zaire and Zambia ( F. Mendelsohn) Shields . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Kibaran . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Katangan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Distrihution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mineral deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . F . The Zambezi Metamorphic Belt in Zimbabwe (T.J. Broderick) . . . . . . . . . . . . G . The Pan-African event in Malawi and eastern Zambia ( K. Bloomfield) Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Lithology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structural history . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Northern Malawi . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Zambia . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Central and southern Malawi . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Pan-African event . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . H . The Damara Mobile Belt in South West Africa/Namibia (R. Mason) Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The basement complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The southern margin inliers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Inliers in the central Damara Belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The northern inliers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Irumide Cycle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Sinclair-Helmeringhausen area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Nauchas/Rehoboth area . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Northwestern Botswana . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Damara Cycle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Nosib G r o u p . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Otavi-Swakop Groups . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Khomas Trough . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Mulden G r o u p . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Nama G r o u p . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Granite generation, regional metamorphism and deformation . . . . . . . . . . . . . . . Ore deposits in t h e Damara Belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The tectonic framework of the Damara Mobile Belt . . . . . . . . . . . . . . . . . . . . .
714 715 717 717 718 718 718 719 719 719 719 720 720 720 721 723 727 727 728 729 737 739 743 745 746 747 750 752 754 754 756 757 757 758 765 765 766 767 768 769 770 771 773 776 777 777 778 781 786
XXII References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 788; Part B . 791; Part C . 794; Part D . 795; Part E . 796; Part Part A . F. 798; Part G . 798; Part H . 799 Chapter 1 1 .
CARBONATITES. ALKALINE AND RELATED INTRUSIONS D.R. Hunter
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The -2000 Ma intrusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The -1300 Ma intrusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chapter I 2 .
788
803 803 807 810
PRECAMBRIAN PALAEONTOLOGY OF SOUTHERN AFRICA K.A. Eriksson and I.M. McGregor
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stromatolites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfossils. organic matter and carbon isotopes . . . . . . . . . . . . . . . . . . . . . . . . Body and trace fossils . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Broader implications of the Precambrian fossil record . . . . . . . . . . . . . . . . . . . . The age of biogenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Archaean-Proterozoic boundary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Proterozoic and metazoans . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
813 813 819 822 825 825 826 827 828
C. SOUTH AMERICA Chapter 13 .
STRUCTURAL FRAMEWORK E . Suszczynski
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Amazonian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Bacaja-Parti tectonic block . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Rondonides orogenic cycle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Guyana orogenic cycle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Rio Negro-Roraima orogenic cycle . . . . . . . . . . . . . . . . . . . . . . . . . . . The Brazilides tecto-orogenic cycle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chapter 14 .
THE GREENSTONE< E . Suszczynski
RANITIC GNEISS SHIELD
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Red schist belts. Goiiis and surrounding areas . . . . . . . . . . . . . . . . . . . . . . . . . . Greenstone belts adjacent to the Atlantic Coast . . . . . . . . . . . . . . . . . . . . . . . . Sialic plutonism. metamorphism and structure . . . . . . . . . . . . . . . . . . . . . . . . . Mineralization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chapter 15 .
835 841 841 842 842 838 843
845 847 848 848 849
THE CRATONIC ENVIRONMENT E . Suszczynski
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mineralization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
851 857 859
XXIII Chapter 16. Chapter 1 7.
ZONES OF REACTIVATION E. Suszczynski . . . . . . . . . . . . . . . .
863
KIMBERLITES, ALKALINE COMPLEXES AND CARBONATITES E. Suszczynski
Kirnberlites . . . . . . . . . Alkaline complexes . . . . Chapter 18.
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THE PALAEONTOLOGICAL RECORD E. Suszczynski . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
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.
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867 868 871
Precarnbrian-Cambrian boundary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8 7 2 References (Chapters 13-18) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 87 2 Subject Index
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875
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Section A
AUSTRALIA
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1
Chapter 1
STRUCTURAL FRAMEWORK OF THE AUSTRALIAN PRECAMBRIAN R.W.R. RUTLAND
INTRODUCTION
The fundamental units of the crust are the fold belts of which the crust is built, and the most important structures are the boundaries between them. However, the present exposure of the Precambrian fold belts of Australia is largely controlled by subsequent platform history, when new structural elements have been developed. Thus the present limits of orogenic provinces are generally not their original limits, and the term block has commonly been used for large units of the Australian basement (e.g., Trendall, 1975, p. 30). Various blocks are named but not defined on the Tectonic Map of Australia and New Guinea (Geological Society of Australia, 1971). These blocks are not fault blocks in the normal sense, nor d o they correspond to the original extent of a major orogenic province. Their boundaries are variously defined by younger orogenic or platformal activity. It will be shown below that the main blocks of Australia owe their form largely t o major reactivation of older basement during the mid-Proterozoic, with an important contribution also from reactivation in the later Precambrian. These deformations have in turn controlled the deposition and deformation of the later Proterozoic and Phanerozoic platform cover. In general the so-called blocks correspond to areas older than that cover (e.g., the Pilbara and Yilgarn Blocks, separated by the Middle Proterozoic Bangemall Basin). On the other hand the term Western Shield (e.g., Trendall, 1975) is not confined to areas of Precambrian basement rocks but includes both “blocks” and Precambrian folded basins in a large area limited by post-Carboniferous cover. Its present form depends on differential movements of the platform during the Mesozoic and also on the Cenozoic history of differential uplift and erosion. In the following sections a more systematic discussion is attempted both of the platform basement, which reveals the earlier orogenic history, and of the platform cover, for the cratonic history. Spatial and temporal relations between orogenic and platform history are also discussed in terms of the chelogenic cycle.
2 This discussion is largely in terms of tectonic units of platform cover and basement whose limits in time are not precisely defined. In many areas isotopic dates are inadequate in number, insufficiently precise, or subject t o varying interpretation when related to the rock record. Moreover, different principles of time-rock nomenclature are in use in different parts, of the continent. These have been summarized by Trendall (1975). In Western Australia a scheme has been evolved which is essentially “an arbitrarily fixed chronological framework . . . which may have no real tectonic significance for the Western Australian Precambrian” (Trendall, 1975, p. 26). The boundary between Archaean and Proterozoic was set in 1966 at 2400 Ma and the Proterozoic was divided into Lower, Middle and Upper with upper boundaries at 1640, 880 and 600 Ma, respectively. This followed Canadian practice (Stockwell, 1964) except that the division at 1280 Ma was not used. The Bureau of Mineral Resources on the other hand has used a time-rock system (McDougall et al., 1965; Dunn et al., 1966) in which rock units at the base of major platform sequences are taken to define the beginnings of Periods. Their Proterozoic Periods (=Systems) are: (1)Lower Proterozoic (=“Nullaginian”), beginning at the base of the Fortescue Group in Western Australia (c. 2300 Ma), (2) Carpentarian, beginning at the Cliffdale Volcanics below the Tawallah Group of the McArthur Basin (c. 1800 Ma); and (3) Adelaidean, beginning at the Paralana Quartzite of the Mt. Painter region in South Australia. The dates are subject to change by revision of the determined age of the rock standard, or by revision of stratigraphic concepts. Thus the most recent evidence has pushed back the age of the base of the Fortescue Group towards 2500 Ma (Trendall, 1978) which therefore becomes a more appropriate division between Archaean and Proterozoic. The corresponding tectonic sequence of platform cover is referred t o as the West Australian Platform Cover on the Tectonic Map of Australia and New Guinea (G.S.A., 1971). Acid volcanics similar t o the Cliffdale volcanics at the base of a major platform sequence are widespread across Northern Australia from the Kimberleys t o the Mt. Isa region. They are associated with post-tectonic granites emplaced in the Lower Proterozoic orogenic belts and have been assigned to a phase of transitional tectonism (G.S.A., 1971). In the west they appear to be significantly older (c. 1900 Ma) than the Cliffdale Volcanics although the dating lacks precision. The tectonic unit of platform cover, which overlies the acid volcanics with minor unconformity, therefore has a diachronous base, partly older and partly younger than 1800 Ma and is referred to as the North Australian Platform Cover (G.S.A. 1971; see also Chapter 4). The Adelaidean has been divided into Willouran, Torrensian, Sturtian and Marinoan Series. The date of 1400 Ma for the base of the Willouran Series in the Delamerian fold belt was based on the assumed stratigraphic correla-
3 tion of volcanics near the base of the sequence, with dated volcanics on the platform. Recently other volcanics on the platform have been tentatively dated at 700 Ma (Webb and Horr, 1978). These are pre-Sturtian on direct stratigraphic evidence but they offer a date for the base of the Adelaidean only if they can be correlated with Willouran Volcanics in the fold belt and if the date is primary (for fuller discussion see Preiss et al., Chapter 5, p. 327). A date from the basement to the Torrensian Series near Adelaide gives a maximum age for the Torrensian of 850 Ma (Cooper and Compston, 1971) while the basement to Adelaidean quartzites in the Musgrave Ranges has given an age of 1050 Ma. Thus, although the base of the Willouran is still imperfectly dated, it is clear that the main part of the Adelaidean System is younger than 1000 Ma and it may be desirable in the future not to apply the term t o sediments between 1400 and 1000 Ma old. Such sediments are included in the tectonic sequence known as the Central Australian Platform Cover on the Tectonic Map of Australia and New Guinea (G.S.A., 1971), especially in Western Australia (e.g., the Bangemall Basin). Correlation between subdivisions of these broadly defined chronostratigraphic units and the mobile belt sequences is of course impossible with any precision in the Precambrian. It seems that a purely geochronometric scale such as that discussed by Harland (1975) is likely to be most useful for intercontinental correlation. BASEMENT PROVINCES AND THEIR SUBDIVISION
Principles for recognition of provinces The discussion of Precambrian tectonics is usually based on the recognition of structural, orogenic or tectonic provinces. Thus in Canada, provinces correspond to discrete geographic entities of basement rocks distinguished initially by structural criteria and confirmed by isotopic dating characteristics. Kroner and Blignault (1976), writing of Southern Africa, propose to define a tectonic province as a “geographic region that is characterized by a combination of such parameters as lithology, structure, metamorphism, and predominant radiometric age, differing significantly from those of adjacent areas”. This non-genetic definition reflects the difficulty of applying Phanerozoic tectonic concepts to the Precambrian. Precambrian tectonic provinces are usually assumed t o have some analogy with Phanerozoic fold belts, but as defined in various regions, some are larger in area and time span than individual Phanerozoic fold belts, while others appear to represent fragments of fold belts, or the products of superposition of different fold belts (e.g., Stockwell, 1966; King, 1969).
4 Since provinces are defined in terms of orogenic characteristics, however, it is clear that province boundaries should be defined in terms of the change from one set of orogenic characteristics t o another in what is now cratonic basement. In the Phanerozoic, the principal unit of tectonic maps (e.g., King, 1969) is the fold belt usually defined in terms of the time span during which the principal rock structures were created. However, the fold belts can usually be distinguished in their histories of sedimentation and volcanism (stratotectonic evolution) as well as in their histories of deformation, metamorphism and plutonism (morphotectonic evolution). In the simplest plate-tectonic models successive fold belts are broadly parallel, and are regarded as largely ensimatic, so that each fold belt forms a new increment of continental crust. These generalizations require substantial qualification even for the Phanerozoic and in particular it may be that large parts of Phanerozoic fold belts are ensialic (e.g., Rutland, 1973a, b, 1976). Nevertheless, in spite of differences in interpretation, division of the Phanerozoic crust into distinct fold belts is quite readily achieved. In the Precambrian, however, division of the crust on this basis is complicated by two main factors. Firstly, Precambrian orogenies are evidently largely ensialic, since the superposition of orogenic belts is often observed. Furthermore, the trends of superposed belts are often at large angles to each other so that fragments of older belts may be found on both sides of younger belts (such relations have been interpreted in terms of collision orogenies following Phanerozoic models, or in terms of intracontinental orogeny for which there is no clear model in the Phanerozoic). For descriptive purposes the value of the term province becomes clear in this context. An orogenic province is a surviving portion of a fold belt, with a particular set of stratotectonic and/or morphotectonic characteristics, which is bounded by other orogenic provinces. Where an original fold belt has been cut by younger narrower fold belts, the older belt may now form several spatially distinct provinces. The Precambrian provinces so defined are sometimes larger in area and time span than individual Phanerozoic fold belts and this is best understood in terms of “gross tectonic cycles rather than single orogenies” (King, 1969, p. 33) or in terms of the chelogenic cycle (see below). The second factor complicating division of the Precambrian crust into fold belts is that, in their present expression, Precambrian fold belts have generally been much more affected by post-orogenic tectonic effects than their Phanerozoic counterparts. Precambrian platforms have been locally affected by rifting and shearing which have had substantial effects on the nature and thickness of platform sedimentation. The present exposure of the Precambrian fold belts is therefore largely dependent on features of subsequent platform evolution, quite unrelated to the earlier orogenic history. Consequently, only portions of provinces are exposed and these are limited by the outcrop of platform cover of various ages. The exposed basement areas are
5 described as orogenic domains on the Tectonic Map of Australia and New Guinea (G.S.A., 1971). Clearly the orogenic province, which now forms part of the platform basement, may be continuous under the platform cover. Thus it may be possible t o group together several orogenic domains in a single province if they show similar orogenic characteristics and if it can be assumed that they have continuity beneath the intervening platform cover. However, the province boundaries, which occur within the platform basement, are commonly concealed in Australia, by the platform cover. The continuity of the orogenic provinces beneath the platform cover is indicated by the trends of gravity anomalies (Wellman, 1976a). The anomalies generally reflect structures in the basement beneath the platform cover (Wellman, 1976a, b) and, therefore, can be used to define basement trends
Fig. 1.1. Gravity trends in Australia (after Wellman, 1976a). Boundaries between crustal blocks defined by gravity trends are shown solid where well defined, and shown broken elsewhere. Areas of Achaean, Proterozoic, and Phanerozoic stabilization have been differentiated. A revised map o f a more limited area is given in Wellman (1978).
6 over the whole continent (Fig. 1.1).Where the trends do in part reflect structures in thick platform cover it seems likely that the cover structures are controlled by the basement trends. The provinces defined in this way give good support to the groupings of orogenic domains made on geological grounds (G.S.A., 1971; Rutland, 1973b, 1976). In particular the gravity trend map strongly emphasizes a broad discordant belt trending between E-W and SE-NW across the continent. This belt was identified on geological grounds as the Amadeus Transverse Zone by Rutland (1973b, 1976). It contains both basement and cover elements which have suffered repeated tectonic activity during the Proterozoic and Palaeozoic. It separates Proterozoic provinces to north and south where tectonic activity terminated much earlier. On the Tectonic Map of Australia and New Guinea (G.S.A., 1971), orogenic provinces have been defined as groupings of domains on the basis of approximate contemporaneity of deformational events and their common relationship to overlying platform cover. This was generally successful and, following the explanation of that map (and excluding New Guinea), can be summarized as follows: (1)West Australian Orogenic Province - deformed Archaean geosynclines and metamorphic complexes, forming youngest basement to West Australian Platform Cover which dates from 2200 t o 1900 Ma. (2) North Australian Orogenic Province - deformed Early Proterozoic sedimentary sequences, forming youngest basement to North Australian Platform Cover which dates from approximately 1800 t o 1400 Ma. (3) Central Australian Orogenic Province - mainly metamorphic complexes forming youngest basement to Central Australian Platform Cover, which dates from beginning of Adelaidean sedimentation, then taken as 1400 Ma, to Permo-Triassic, but with widest extent in the Early Palaeozoic. ( 4)Late Precambrian Domains. ( 5) East Australian Orogenic Province - geosynclines deformed during Phanerozoic and forming youngest basement to Trans-Australian Platform Cover, which dates from uppermost Carboniferous to Present but with widest extent in Jurassic and Lower Cretaceous. The main weakness of this scheme, apart from the cumbersome terminology, is in the artificial grouping of domains with very different tectonic histories in the Central Australian Province. These domains not only show varying intrinsic characters but they are usually not closely related by transitional tectonics t o the Central Australian Platform Cover. Rutland (197313, 1976) distinguished two kinds of orogenic belts. The dominant type he regarded as broadly analogous to Phanerozoic orogenic belts developed between the arc-trench complex and the craton (miotectonic belts). Platform covers were distinguished both equivalent to and postdating these mobile belt sequences. The second type of orogenic belt (represented by the Fraser and Musgrave belts) was regarded as having no
7 close analogue in the Phanerozoic and was attributed to “high heat flow in narrow zones of lithosphere reactivation” rather than to continental collision. Rutland (1973b) also applied the concept of chelogenic cycles to the Australian Precambrian. He suggested that three chelogenic cycles could be recognized and, therefore, distinguished three major provinces (Fig. 1.2): (1)The Archaean Pilbara-Yilgurn Province -- corresponding to the West Australian Orogenic Province above. ( 2 j The Early and Middle Proterozoic Arunta-Gawler Province - including both the North Australian and Central Australian Orogenic Provinces above. ( 3 ) The Late Precambrian and Phanerozoic Tasman Province - including both the Late Precambrian Domains and the East Australian Province above. Within each of these major provinces the various orogenic domains can be clearly related to each other, especially in terms of their main plutonic and deformational events. On the other hand there are very strong contrasts in the age of basement rocks across the province boundaries. All of these provinces consist of orogenic domains in which climatic deformation, metamorphism and plutonism has been superimposed on preceding sedimentary and volcanic sequences. In addition a fourth province was distinguished which appears to represent a different kind of orogeny. This Musgrave-Fraser Province appears t o be the consequence of superposition of younger orogeny on the AruntaGawler orogenic complexes. In part it lies in a zone, the Amadeus Transverse Zone (Rutland, 1973b) which has been subject to intermittent tectonic mobility in both the Proterozoic and Phanerozoic and, perhaps, also in the Archaean. This general scheme for grouping orogenic domains into major provinces will be used here, but the two major Precambrian provinces previously distinguished (Pilbara-Yilgarn and Arunta-Gawler) are now regarded as superprovinces corresponding to the respective Archaean and Proterozoic chelogens. The term province is therefore now used as defined above. Consequently, parts of the same chelogen which are geographically separated by younger orogenic belts are now raised t o province status. Where differences in basement characteristics justify subdivision of the provinces, su bprovinces and zones can also be distinguished. The classification of provinces and subprovinces used here is as follows: (A) Archaean superprovince or chelogen (1)Pilbara Province ( 2 ) Yilgarn Province (a) Eastern Goldfields Subprovince (b) Murchison Subprovince (3) Kimberley Province (B) Early and Middle Proterozoic superprovince or chelogen
8
Fig. 1.2. Map showing the principal exposed orogenic domaim in Australia (after Rutland, 1976). The ornamentation of the domains indicates their allocation t o three major provinces, corresponding t o the three superprovinces or cheiogens of the present paper. The term paratectonic is used t o describe folding in basins within or on the margins of cratons, in contrast t o orthotectonic folding (see e.g., Dennis, 1970).
(4)North Australian Province (a) Halls Creek Subprovince (b) Palmerston Subprovince (c) Mt. Isa Subprovince (5) Gascoyne Province (a) Gascoyne Subprovince (b) Paterson Subprovince (6) Gawler Province ( C ) Middle Proterozoic reworking of Proterozoic chelogen ( 7 ) Musgrave-F’raser Province
9 Thus the hierarchy of terms, province, subprovince and zone describes the orogenic subdivisions of the platform basement as it exists today. The qualifying adjecth orogenic (G.S.A., 1971) is preferable to tectonic or structural but, as parts of the present platform basement, the provinces can also be described as basement provinces. The terms themselves are deliberately non-genetic. Interpretation of the evidence will show that each subdivision has passed through various genetic or evolutionary stages. The early evolution of a subprovince, for example, might be described in various stratotectonic terms such as mobile belt, geosyncline, miogeocline, arc-trench complex, or marginal basin depending on the nature of the interpretation. Subsequently it has been converted into a fold belt; and subse.quently again it has become part of a stable craton, subject t o platform deposition.
The chelogenic cycle and plate tectonics In Phanerozoic fold belts some link between the earlier stratotectonic history (of sedimentation and volcanism) and the later morphotectonic history (of deformation, metamorphism and plutonism) is usually inferred. Consequently the morphotectonic boundary between fold belt and platform is usually related to an earlier change in thickness and facies of sedimentation. In the simplest plate-tectonic scheme, fold belts are interpreted in terms of the “Wilson cycle” (Dewey and Burke, 1974) of development and destruction of ocean basins by ocean-floor spreading and subduction. The spatial associations, and histories, of fold belts strongly suggest, however, that long-term tectonic cycles, the chelogenic cycles (Sutton, 1963, 1967), must also be taken into account. The circum-Pacific tectonic evolution shows particularly clearly that a major tectonic contrast exists between the Precambrian shields and platforms, on the one hand, and the whole assemblage of Late Precambrian and Phanerozoic fold belts on the other. There are evidently close relationships in both space and time between the various fold belts which make up the latter assemblage and they can be grouped together as the products of a single chelogenic cycle (e.g., Rutland, 1973a, 1976). The oldest sedimentary rocks associated with these assemblages appear to have been developed on passive continental margins, usually assumed t o have been developed by rifting and drifting. The various fold belts, such as those of the Tasman Province in Australia, correspond with various episodes in the evolution of the active continental margins developed from the passive margins. All of the fold belts have been rejuvenated during the Late Cenozoic morphogenic phase, emphasizing the tectonic contrast between the assemblage as a whole and the adjacent craton. The sequence of rocks deformed within the whole assemblage of Palaeozoic fold belts corresponds with one major division of the platform cover of the adjacent craton. In Australia this platform cover consists of Upper Proterozoic (Adelaidean) and Palaeozoic rocks and has little or no relation
10 t o the earlier orogenic history of the cratonic basement. It is the Central Australian Platform Cover (G.S.A., 1971). Overlying the assemblage of Palaeozoic fold belts in the Tasman Province is another major platform unit, largely of Mesozoic age, named the TransAustralian Platform Cover. The cover is closely linked, via transitional tectonism including acid volcanism, with the orogenic history of the underlying orogenic belts. Thus it represents a typical post-orogenic stage. It also extends onto the Proterozoic provinces (Fig. 1.2) which formed the adjacent craton, where it has a different pattern of distribution from the underlying Central Australian Platform Cover. The younger platform cover is equivalent in age to fold belts formed during the main period of continental fragmentation and dispersal in the Mesozoic. Thus the chelogenic cycle can be divided into two main stages: (1)early platform cover stage (Central Australian Platform Cover) on the craton, corresponding in the mobile belt t o sequences related t o the active continental margin; (2) late platform cover stage (Trans-Australian Platform Cover; G.S.A., 1971) formed after formation and stabilization of fold belts from mobile belt sequences and extending across newly consolidated chelogen and previously existing craton. The main deformation, plutonism and metamorphism in the fold belts occurs in a somewhat diachronous interval between these two main sedimentary stages. In general, the platform cover of this interval, encompassing much of the later Palaeozoic, is less well developed on the craton so that there is a marked distinction between the two main elements of platform cover. This scheme can be successfully applied (Rutland, 1973b) to recognize a major Proterozoic chelogenic cycle in Australia (Fig. 1.4): (1) early platform cover stage (West Australian Platform Cover; G.S.A., 1971) developed on Archaean craton and corresponding t o Early Proterozoic mobile belt sequences (c. 2500-1800 Ma); (2) late platform cover stage (North Australian Platform Cover; G.S.A., 1971) formed after stabilization of the main Proterozoic fold belts (c. 1900-1400 Ma). The main deformation, plutonism and metamorphism in the Proterozoic fold belts occurs in the diachronous interval between these two stages (c. 1900-1700 Ma). Thus the existence of such a Precambrian chelogenic cycle largely explains why it has been possible to divide the Precambrian orogenic basement into major provinces exhibiting much larger contrasts in age than between individual Phanerozoic fold belts. The most important contrasts are those between provinces belonging to different chelogens. In Australia the Archaean chelogen is distinguished as those orogenic provinces which formed cratonic basement to the early platform cover of the Proterozoic chelogenic cycle. It is not possible t o identify platform sequences
11 corresponding to the Archaean mobile belt sequences within Australia. As already indicated the four main tectonic assemblages of platform cover in Australia fall into two pairs corresponding to the two post-Archaean chelogenic cycles. In space, chelogenic boundaries in Australia are defined where little deformed and unmetamorphosed early platform cover, resting on orogenic basement or platform cover of the earlier chelogenic cycle, passes laterally into deformed and metamorphosed mobile belt sequences of the later chelogenic cycle. Thus, for example, the Lower Proterozoic platform cover on the Archaean basement of the Pilbara province (c. 3100 Ma plutonism) passes laterally into a deformed and metamorphosed sequence in the Gascoyne Province of the Proterozoic (c. 1700 Ma plutonism). Similarly the Upper Proterozoic (Adelaidean) platform cover on the older Proterozoic basement of the Gawler Province (plutonism c. 1700 Ma) passes laterally into a deformed and metamorphosed sequence in the Kanmantoo subprovince (plutonism c. 500 Ma) of the Tasman Province (Fig. 1.2). Therefore, it is notable that the main concentrations of plutonic and metamorphic activity fall within the chelogenic cycle. In time the boundaries between the cycles are marked by major phases of basic dyke emplacement. One major phase of dyke emplacement affects the Archaean chelogen at about 2400 Ma and marks the beginning of the Proterozoic chelogenic cycle. A second major phase affects the Proterozoic chelogen at about 1000 Ma. In terms of plate-tectonic concepts it seems likely that the basic dyke phases mark major changes in the pattern of ocean-floor spreading and development; and that the main periods of plutonism during the chelogenic cycle are associated with the transition t o major phases of continental dispersal. In the Phanerozoic, widespread Late Palaeozoic plutonism in Hercyniantype belts precedes the transition while the main (Cretaceous) phase of circum-Pacific plutonism correlates with the maximum rate of ocean-floor spreading after the transition. The 1900-1700 Ma Proterozoic plutonic maximum in Australia is analogous t o the later Palaeozoic plutonism in Australia (Rutland, 1973b) and by inference, therefore, probably precedes a major phase of continental dispersal. The period between the termination of widespread deformation and plutonism and the main episode of basic dyke emplacement is the main period of megashear and rift development in the continental plates (14001000 Ma) and probably corresponds t o a period of continental dispersal and reassembly (cf. Sawkins, 1976). During this period plutonism was presumably largely confined to narrower zones at continental margins which are not now preserved in Australia. The chelogenic cycles as revealed in Australia show a broad correlation with those in other continents. The Pilbara and Yilgarn Provinces are here treated as representing two orogenic cycles within one chelogenic cycle but they could be regarded as representing two chelogenic cycles analogous to
12 the Kola and Shamvaian cycles of Sutton (1967). The Proterozoic chelogen in Australia clearly corresponds to Sutton’s Svecofennid chelogenic cycle. There can be little doubt that the similarity of timing of these major cycles in all the continents reflects long-term global thermal cycles.
The Archaean chelogen The principal exposed portions of the Archaean chelogen are the Pilbara and Yilgarn orogenic domains. They are here allocated t o two provinces since they are separated by younger orogenic basement of the Gascoyne Province (Fig. 1.3). Rather similar Lower Proterozoic platform sequences overlie the southern part of the Pilbara Province and the northern edge of the Yilgarn Province in the Hamersley and Nabberu basins respectively, although there is doubt as to how closely they can be correlated (Bunting et al., 1977). The Archaean province boundaries are reasonably well defined where these platform cover sequences pass laterally into strongly deformed and metamorphosed sequences in the Gascoyne Province with plutonism at c. 1700 Ma. It seems likely that the two Archaean provinces were originally joined in a single Archaean chelogen (e.g., Honvitz and Smith, 1978), but there are significant differences between them. The Pilbara region was probably already of cratonic character while the Yilgarn mobile belts were developing. The boundary zone between these two regions later became the site of the Gascoyne Orogenic Province and later again of the Bangemall Basin. The Pilbara domain is a granite-greenstone belt complex in which the principal granites are older than 3000 Ma and the greenstones are about 3450 Ma (Pidgeon, 1978). The greenstone formations are locally strongly deformed in northsouth-trending zones but in general they form irregular outcrops between the dome-like granitic bodies, so that they lack the regular trend characteristic of the Yilgarn domain. The northern and northwestern margins of the Pilbara Province are obscured by younger cover and the Indian Ocean, but, elsewhere around the margins, small inliers of Archaean rocks reveal younger Archaean dates like those of the Yilgarn domain. Thus it appears that the present form of the Pilbara Province was already outlined in the younger Archaean. DeLaeter et al. (1977) suggest, on isotopic grounds, that the eastern edge of the province corresponds approximately to the eastern edge of the older granitic crust i.e. the boundary is not simply due to younger Archaean reworking. East-west trends indicated by deformation of the late Archaean Mosquito Creek Group and by basic plutonism anticipate the later deformational trends both in the Gascoyne Province to the south and in the Amadeus Transverse Zone to the north. As recognized by Arriens (1971), and confirmed by subsequent workers, the principal granites of the Eastern Goldfields Subprovince of the Yilgarn
.-" 2L
n Concealed Archaean
Concealed Reworked Archaean Basement
n
300 k m
0 L
I
'5
--
Archaean
granitic
Archaean
greenstone
Older with
rocks
( A 1 Pilbara,
Archaean gneisses and younger granites ( A p )
Major
fault
zones
Major
dyke
trends
.Wia
A2 Yilgarn)
belts metasediments
(Al)
del. a.m.:. ' 7
Fig. 1.3. Principal tectonic elements of the Archaean chelogen. Note that the Pilbara and Yilgarn orogenic domains are the exposed portions of the more extensive Pilbara and Yilgarn provinces. The provinces include areas of concealed Archaean basement and are separated by areas of younger (reworked Archaean) basement.
14 Province were emplaced between 2650 and 2450 Ma (X87Rb= 1.42 lo-" a-'). The associated greenstones are probably younger than the high-grade gneisses of the Wheat Belt Subprovince which have yielded dates of 30302840 Ma. Thus the Wheat Belt Subprovince appears t o reflect the older Archaean orogeny of the Pilbara Province. However, it is notable that the metasediments involved appear t o represent a shelf facies rather than the greenstone facies which is seen in the Pilbara. The younger Yilgarn orogeny imparted the pronounced NNW trend to the greenstone belts, and a number of N- and NNW-trending major faults and shear zones were also developed. Subsequently (c. 2400 Ma) a suite of major basic dykes were emplaced with WSW-ENE trends approximately normal to the orogenic trend. These phenomena may be interpreted in terms of crustal shortening and relaxation with maximum principal stress in an ENE--WSW direction. O'Driscoll (1971), however, interprets the folding in terms of dextral simple shear on major lineaments trending NE-SW. There appears t o be no independent evidence of such a fundamental shear direction in the older Archaean rocks of the Wheat Belt Subprovince or the Pilbara Province. Several trends of basic dykes were emplaced in the Pilbara Province at about the same time as those in the Yilgarn but the best developed set trends NNE. This apparently reflects a different crustal stress system and does not imply that the provinces have been rotated with respect t o each other (e.g., Embleton, 1978). It has already been noted that the northern boundary of the Yilgarn Province is defined by the transition into the Gascoyne Province. In the southeast and south the boundaries are also well defined by the transition to younger orogenic provinces but in these cases no transition from platform to mobile belt sequences occurs. A transition zone carries granitic rocks of similar age to those in the Gascoyne Province but younger orogenic phases are also present (Bunting et al., 1976). A phase of granulite facies metamorphism associated with a major fault at about 1300 Ma is most important. Inliers of Precambrian metamorphic rocks in the Perth Basin also provide evidence of various phases of post-Archaean deformation and highgrade metamorphism (e.g., at c. lOO_O Ma, Northampton domain and 650 Ma, Naturaliste domain, Fig. 1.6). Thus it can be inferred that the Archaean chelogen once extended beyond the limits of the Yilgarn Province. The province boundaries were produced by the superposition of younger orogenic phases and have been subject to reactivation at various times during later Precambrian history. This phenomenon is evidently of great importance to theories of lithospheric and upper mantle evolution. Most recently, the western boundary of the province provided the locus for the Darling Fault during the rifting of Gondwanaland. A third Archaean orogenic province may be present in the Kimberley region, framed by the King Leopold and Halls Creek orogenic domains. How-
15
Fig. 1.4. Principal tectonic elements of the Early and Middle Proterozoic chelogen (2400--1400 Ma, excluding reactivation features, 1400-1000 Ma). The three Proterozoic provinces are three areas of Lower Proterozoic orogenic domains and their concealed extensions separated by areas of younger orogenic basement (Fig. 1.6). T h e earlier Archaean provinces formed stable cratons during the development of t h e Early a n d Middle Proterozoic chelogen. The oceans described as Carpentarian Platform Cover are more correctly described as North Australian Platform Cover (see pp. 2-3).
16 ever, the region is now completely concealed by the North Australian Platform Cover (Fig. 1.4). Much smaller inliers of Archaean rocks occur in a more or less reworked condition within the Proterozoic orogenic provinces. They indicate that the Archaean chelogen once extended over a much wider area. Indeed it seems likely that in general the Proterozoic provinces are underlain by reworked Archaean basement although the possibility that marginal basins, floored by oceanic crust, were developed in the Proterozoic cannot be ignored. However, it is clear that the Proterozoic provinces are not essentially due t o lateral continental accretion (Rutland, 1973b; cf. Kroner, 1977).
The Proterozoic chelogen Much of the Proterozoic chelogen is concealed by platform cover so that the relations between the various orogenic domains are often obscure. The overall unity of the chelogen is demonstrated by the presence in all domains of a major phase of plutonism, often closely related to acid volcanism immediately preceding the overlying platform cover sequence, in the range 19001700 Ma. The domains are here grouped into three various provinces which are thought to be separated by belts of younger orogenic basement formed during the Musgravian orogenic cycle (1400 -1000 Ma). The effects of the Musgravian orogenic cycle are best displayed in the Redbank and Woodroffe zones, north and south respectively, of the Amadeus Basin. These zones are marked by pairs of positive and negative gravity anomalies (“dipole” anomalies) which have been interpreted as defining the boundaries of crustal blocks of different mean densities in isostatic equilibrium (Wellman, 1978). Similar anomalies in Canada have been interpreted in terms of collision orogeny (Thomas and Gibb, 1977). The gravity maps (see for example, Anfiloff et al., 1976) suggest that the Musgrave positive anomaly may be continuous with an anomaly beneath the southern margin of the Canning Basin and the Paterson Subprovince (see below p. 23). A further dipole anomaly marks the Fraser orogenic domain of the Musgrave Province and this may extend northwards to join the Musgrave domain. Thus the zones of Musgravian tectonic activity apparently divide the older Proterozoic chelogen into three main provinces (Figs. 1.4 and 1.5). The domains north of the Amadeus Transverse Zone can be grouped together as the North Australian Province (see Chapter 4). Those south of the Amadeus Transverse Zone and east of the Fraser Zone can be grouped in the Gawler Province (Chapter 5) and the remainder can be allocated t o the Gascoyne Province (Chapter 3). The definition of the North Australian Orogenic Province (G.S.A., 1971) is here extended t o include the northern part of Arunta domain and the Mt. Isa, Georgetown and Coen orogenic domains. The Mt. Isa domain was
17 allocated to the Central Australian Orogenic Province (G.S.A., 1971) but its geological relationships demand a revision of that allocation, while the Arunta, Georgetown and Coen domains were unassigned. The North Australian Orogenic Province is here divided into three subprovinces on the evidence of the age of the sequences present and of the timing of the climactic deformations. However, plutonism dated at about 1800 Ma is present in all three subprovinces. The Halls Creek Subprovince includes the Halls Creek and King Leopold orogenic domains. Inliers of Archaean basement may be present in the belt (Ding Dong Downs Volcanics) but the main mobile belt sequence is Lower Proterozoic and was involved in major deformation and plutonism at c. 1960 Ma (Bofinger, 1967, ha7Rb= 1.3910-" a-'). The voluminous younger Bow River Granite and related rocks are dated at about 1850 Ma. Archaean inliers also occur in the Palmerston Subprovince" which includes the Pine Creek, Tennant Creek, Granites-Tanami and northern Arunta orogenic domains. The mobile belt sequences of low-grade metasediments in these domains may be of the same age as those in the Halls Creek Subprovince (e.g., Plumb and Derrick, 1975, pp. 220-221; Blake, 1978) but they may in part be younger. Their minimum age is given by post-tectonic granites (c. 1830 Ma) and the overlying Cliffdale Volcanics (1770 20 Ma) at the base of the overlying Carpentarian platform cover sequence. The basement of the Mt. Isa Subprovince is intruded by granite dated at c. 1860 Ma and overlying acid volcanics are dated at c. 1775 Ma, but the Carpentarian platform cover sequence itself becomes involved in strong folding and metamorphism with associated plutonism at c. 1670 Ma and 1450-1400 Ma (e.g., Page, 1976). The Lower Proterozoic flysch-type successions of the North Australian Province have been subjected to high-grade metamorphism in relatively localized zones. Over much of the exposed domains the metamorphic grade is relatively low. Thus the Palmerston Subprovince in particular shows broad similarities with the Lachlan Subprovince of the Tasman Province (Fig. 1.2; Rutland, 197313). The Gawler Province, however, displays significantly different characteristics even though it was developed during broadly the same time span. The Lower Proterozoic sequences are thinner and of a more platformal character (e.g., Glen et al., 1977). They are nevertheless strongly deformed and metamorphosed to high-grade. Gneissic granites are abundant, usually structurally near the base of the section but they are mainly younger than the metasediments. For example, Cooper et al. (1976) have dated such granites at c. 1820 Ma but they have also recognized relict areas of Archaean gneisses dated at c. 2600 Ma.
*
____
* Palmerston is a n earlier name for t h e city of Darwin.
18 Thus the North Australian Province contains fold belts developed from more “geosynclinal” sequences, whereas the Gawler Province contains fold belts developed from essentially platformal sequences. Therefore, the latter fold belts clearly involve the orogenic reworking of Archaean craton. The Amadeus Transverse Zone including the Musgrave Province now lies across the boundary between these two major Lower Proterozoic stratotectonic provinces. However, it seems likely that there was no abrupt change of sedimentary facies between the two provinces and the available data (Stewart and Warren, 1977) suggest that the Lower Proterozoic sediments of the Arunta domain have an intermediate character. Although the main metamorphism and plutonism in the Gawler Province is of about the same age as that in the Palmerston Subprovince, it was not followed by the development of a widespread platform cover like the Carpentarian cover. Instead the Gawler Province provides evidence of further deformation, plutonism and retrograde metamorphism culminating in the extrusion of the Gawler Range Volcanics at about 1500 Ma (Fig. 1.4). The Gascoyne Province (Gascoyne domain, Marymia dome and Rudall metamorphics in Fig. 1.4) contains a metamorphosed mobile belt sequence apparently equivalent to only the uppermost part of the Lower Proterozoic platform cover on the Pilbara Province. Daniels (1975, pp. 108-110) suggests that an Early Proterozoic phase of migmatization and deformation preceded deposition of the metasedimentary sequence which is itself intruded by granites dated at 1700-1800 Ma (De Laeter, 1976). Thus the mobile belt sequence is apparently equivalent in age to only the youngest part of the mobile belt sequences in the North Australian Province, but the significance of the Early Proterozoic (c. 2000 Ma) orogeny in both provinces requires further study in relation to the age of the mobile belt sequences. However, there is no doubt that deposition in all three provinces terminated prior to the ubiquitous 1800 f 100 Ma plutonism. Only in the Mt. Isa Subprovince was a succeeding platform cover involved in further orogeny although younger plutonic events are also found in the Gawler Province. All the Proterozoic provinces appear to be essentially ensialic although the presence of marginal basins floored by oceanic crust cannot be excluded. Since there is some evidence of east-west polarity in the provinces with younger tectonism in the east, it seems likely that the active continental margin during the Proterozoic lay east of the exposed provinces, in the region now occupied by the younger Tasman Province (Rutland, 1973b, 1976). Consequently it is inferred that the continental crust of the Proterozoic province is not the result of lateral accretion during the Proterozoic chelogenic cycle. It is the product of vertical accretion in the Proterozoic on, and within, a pre-existing Archaean proto-continental crust. Any lateral accretion must be attributed t o the Archaean.
19
The Musgrave-Fraser Province As indicated above, most of the Proterozoic chelogen was stabilized by about 1800 Ma before the deposition of the Carpentarian platform cover. Even where the Carpentarian cover was involved in younger orogeny in the east of the chelogen, plutonic activity and folding were completed before 1400 Ma. The younger tectonism of the Musgrave-Fraser Province was apparently superposed on these earlier Proterozoic provinces and it was not preceded by the deposition of younger mobile belt sedimentary or volcanic sequences. In the Musgrave and Redbank belts (Fig. 1.6) major crustal dislocations are associated with mylonite development, granulite facies metamorphism and both acid and basic plutonism (e.g., Marjoribanks and Black, 1974; Thomson, 1976; Moore and Goode, 1978). The main Musgravian metamorphism is dated at 1300-1400 Ma but it is superimposed on the earlier metamorphism at c. 1800 Ma (both in the Musgrave and the Fraser domains) and various intermediate events can also be identified. The granite plutons, which post-date the main mylonite development are dated at about 1100 Ma and the layered mafic-ultramafic complexes at about 1050 Ma. As already noted, the largest “dipole” gravity anomalies in Australia are associated with these major crustal dislocations which have the character of reversed faults or thrusts. On these dislocations the granulite facies rocks have been carried upwards and towards the present Amadeus Basin (which was developed later) from both north and south (compare Figs. 1.5 and 1.6). Rutland (1973b) suggested that the closest Phanerozoic analogue of this kind of intracontinental “orogeny” is in the Wyoming uplifts of the Cordilleran system of the U.S.A. This analogy has now been reinforced with the demonstration by deep seismic profiling that the Wind River Thrust is a major thrust extending at least to 24 km and possibly t o 36 km depth (Brewer et al., 1978). The Musgrave-Fraser Province here includes the zones of orogenic activation both north and south of the Amadeus Basin together with those of the Fraser and Albany domains (Fig. 1.6). It may also be appropriate t o include the inliers of granulite facies (Naturaliste and Northampton domains) west of Yilgarn Province (see below). It should be emphasized that these elements have very variable trends. Therefore, they cannot reasonably be regarded as lying on small circles about a point of rotation off the Australian part of the early continent (Davies and Windley, 1976). Rather they are associated with taphrogenic reworking of the craton, in which various shear trends occur in association (e.g., Glen et al., 1977). THE BLOCK STRUCTURE OF AUSTRALIA: BASEMENT REACTIVATION
The Musgrave, Redbank and Fraser zones are the most striking manifestations of the Musgravian orogeny and justify the distinction of a separate
20
province. However, the period between about 1400 Ma and 1000 Ma ago was also notable for widespread faulting and shearing throughout the Proterozoic chelogen (but not within the Archaean chelogen). This deformation is largely responsible for determining the present “block” structure of Australia. The less-deformed areas now stand as the main basement blocks of Australia, both Archaean and Proterozoic. The more-deformed zones have become the sites of thicker sedimentation, both during the Upper Proterozoic and in the Phanerozoic, so that younger basins now separate the basement “blocks”. The younger basins have themselves suffered deformation controlled by the underlying Middle Proterozoic basement structures; and the block boundary zones have also been the sites of subsequent metamorphism and minor plutonism, especially in the latest Precambrian.
The Archaean blocks The nature of the boundaries of the Pilbara and Yilgarn Provinces has been discussed above (pp. 12-14). The boundaries of the Pilbara Province were apparently partly outlined in the later Archaean. Both provinces, however, were probably largely outlined by encroachment of the Early Proterozoic orogenic provinces: the Middle Proterozoic Fraser belt (c. 1300 Ma) is evidently superimposed on earlier Proterozoic orogenic basement (16001700 Ma) and this is probably true for other areas of younger orogenic basement around the Archaean provinces. Middle Proterozoic faulting and shearing has also tended to follow the earlier Proterozoic orogenic activity. Daniels (1975) emphasizes the importance of this deformation and of its control of the subsequent deposition and renewed deformation of the Middle Proterozoic Bangemall Basin (older than 1000 Ma). Similarly, as will be indicated below, the western and northern margins of the Pilbara Province have been accentuated by Middle Proterozoic deformation. Evidently too, the Darling Fault, formed on the western edge of the Yilgarn Province during the separation of Western Australia from Gondwanaland, has also been controlled by Proterozoic tectonic trends. Consequently the Archaean orogenic provinces and the blocks defined by the subsequent platform history are essentially co-extensive.
The Amadeus Transverse Zone The most important zone of deformation is the Amadeus Transverse Zone which separates the two largest Proterozoic provinces. As already noted the coherence of this zone is indicated by the distinctive and consistent trend of gravity anomalies, but the geological evidence indicates that these trends have been developed or rejuvenated at different times in different places. Trends parallel to the zone are weakly developed in the early Archaean Pil-
21 bara Province and strongly developed in the Early Proterozoic King Leopold Zone of the Halls Creek Subprovince. The main basement expression of the Amadeus Transverse Zone is in the southern Arunta domain (the Redbank Zone) and in the Musgrave domain. The principal gravity anomalies of Australia are clearly related t o these Musgravian orogenic structures, which have therefore evidently been a major factor in defining the block structure of Australia (Figs. 1.5 and 1.6). The principal gravity anomalies, however, are confined t o the areas of basement domains. Their continuation, and therefore the continuation of the Middle Proterozoic structures with which they are associated, is conjectural east and west of the Musgrave and Arunta domains. The main positive anomaly associated with the Woodroffe Zone on the south of the Amadeus Basin is apparently continuous with a weaker anomaly beneath the southern margin of the Canning Basin and the Paterson Subprovince (Fig. 1.2). However, the significance of this is not clear since the geology of Paterson Subprovince (Williams et al., 1976) has little in common with the Woodroffe zone of the Musgrave Province. In particular there is no evidence of a similar granulite facies zone. The Rudall metamorphic complex of the Paterson Subprovince has yielded a date of about 1500 Ma and is tentatively allocated to the Gascoyne Province. However, similar dates have been found in the Arunta domain and in the King Leopold domain and may represent an early phase of reactivation in the Amadeus Transverse Zone. (In the Fraser domain too, the granulite facies zone dated at 1300 Ma is flanked by successive zones with plutonism at c. 1600 Ma and 1700 Ma.) Moreover, a strong foliation and mylonite development is present in Archaean granites on the western margin of the Paterson Subprovince and a mineral age of 1200 Ma which may correspond to this foliation has been obtained from whole rockmineral Rb/Sr data (De Laeter et al., 1976). The foliation may therefore be of similar age t o the main phase of mylonite development in the Musgrave Province. The Paterson Subprovince also contains post-tectonic granites dated at c. 600 Ma and these can be correlated with the Late Precambrian Petermann Ranges orogeny on the southern margin of the Amadeus Basin. Thus the southern margin of the Amadeus Transverse Zone, as indicated by the continuity of gravity anomalies, can be related t o a combination of Musgravian and younger tectonic events. Eastwards from the Woodroffe zone the positive anomaly and the flanking negative anomalies both weaken a t the edge of the exposed orogenic domain, where they are trending ENE. Possibly the Middle Proterozoic structures with which the anomalies are associated converge with those of the Redbank Zone north of the Amadeus Basin. Alternatively they may be related t o the positive anomalies and younger Proterozoic structures which mark the northern edge of the Gawler “Block” (see below).
22
Fig. 1.5. Pattern of free-air gravity anomalies in relation to tectonic elements from Fig. 1.2. (Anomalies simplified from Anfiloff et al., 1976).
North of the Amadeus Basin the strong dipole gravity anomalies terminate both to east and west within the exposed basement complex of the Arunta domain. Thus in the east they terminate in a NW-trending zone which has been called the Woolanga lineament (Anfiloff and Shaw, 1973). The main fault trends in the domain, and the main boundary between basement rock
23
Fig. 1.6. Features of t h e reactivation of t h e Proterozoic chelogen, 1400-1000 Ma, in relation t o t h e distribution of t h e Central Australian Platform Cover.
zones, are more nearly E-W like the gravity anomalies, so that, apart from some faults, the Woolanga lineament has little obvious expression in the basement geology. Granitic rocks north of, and granulite facies south of, the main zonal boundary extend across the Woolanga lineament (Stewart and Warren, 1977); but, while the main outcrops of granulite rocks west of the lineament correspond with the main positive gravity anomaly, there is no
24
corresponding anomaly related t o the granulites immediately east of the linement . Further to the east another positive anomaly on the southern margin of the Georgina Basin may have similar tectonic significance, and appears to have a further continuation to the southeast beneath the Eromanga Basin. Broadly therefore the break in the positive anomaly at the Woolanga lineament corresponds to the gap between the upper Proterozoic Ngalia and Georgina Basins (Fig. 1.6). West of the Arunta domain the anomaly pattern of the Amadeus Transverse Zone is interrupted by north-south trends but it is possible that the positive gravity anomalies associated with the margins of the Phanerozoic Fitzroy Trough represent basement continuations of the Middle Proterozoic zones of deformation of the Arunta domain (marked by the positive gravity anomalies north and south of the Ngalia Basin). Definite evidence of Musgravian structures is lacking but there is evidence of tectonic activity in the King Leopold domain on the northern margin of the Fitzroy Trough, provided by local metamorphism t o amphibolite facies of the Kimberley Basin platform rocks and isotopic events at about 1550 and 600 Ma (Plumb and Derrick, 1975, p. 231). Because of the repeated, largely taphrogenic, deformation within the King Leopold and Halls Creek domains both during the Proterozoic and to a lesser extent in the Palaeozoic, Plumb and Derrick (1975) have followed Traves (1955) in applying the terms Halls Creek and King Leopold Mobile Zones to these belts. Thus, in the Precambrian basement rocks, the Amadeus Transverse Zone is marked by a collection of structures and isotopic events of different ages. Taken together with the gravity evidence, they define a zone approximately 500 km wide separating more stable shield areas, with different trends, t o north and south. The most significant element in the development of the zone is the Musgravian Orogeny but it is not clear that this orogeny was of major importance throughout the zone. The Middle Proterozoic crustal dislocations probably varied greatly in the magnitude of their displacements so that granulite facies rocks became exposed at the surface by erosion only in those domains which axe now exposed and which display the large gravity anomalies. I t should be noted that, in the Woodroffe Zone the granulite facies rocks were already exposed to erosion in the Middle Proterozoic. They are overlain unconformably by the Tollu Volcanics, a bimodal volcanic suite dated at about 1050 Ma and related to the layered mafic and ultramafic intrusions of the Giles complex (Nesbitt et al., 1970). It can be inferred, therefore, that the Amadeus Transverse Zone already existed as a transcontinental structure in the Middle Proterozoic but at that stage it was also connected with zones of deformation now seen in the Fraser and Gascoyne domains. The further development of the Amadeus Transverse Zone took place
25
through its control on Upper Proterozoic and Phanerozoic sedimentation, and through the reactivation of its margins especially in the latest Precambrian. Thus the Late Precambrian Petermann Ranges “Orogeny ” and Devonian-Carboniferous Alice Springs “Orogeny ” reactivated the Middle Proterozoic structures on the southern and northern margins of the Amadeus Basin respectively. It is these deformations which define the present block boundaries. However, in Western Australia, the mid-Palaeozoic tectonism is marked by graben development along the old block boundaries (Fitzroy and Petrel grabens) rather than by compressional deformation as in the Amadeus Basin. The development of the Pedirka Basin and the continued development of the Canning Basin show that the Amadeus Transverse Zone continued t o exercise a control on sedimentation into the Mesozoic.
The Kimberley Block The Kimberley Block is defined by an area of little deformed North Australian Platform Cover (G.S.A., 1971) and is limited by the King Leopold and Halls Creek orogenic domains. I t has been suggested that the Kimberley Block is underlain by an Archaean orogenic province (Gellatly, 1971) and this may be the best explanation of the change of trend of the Lower Proterozoic folding between the two flanking orogenic domains. If this is so it is also possible, as suggested by Rutland (1977), that the two orogenic domains were initiated as aulacogens in a manner analogous t o the postulated Athapuscow and Bathurst aulacogens flanking the Slave Province in Canada (Hoffman, 1973). Such an interpretation would suggest that a Lower Proterozoic active continental margin lay northwest of the Kimberley Block. However, it should be noted that the block boundaries, as defined by the gravity dipole anomalies (Wellman, 1978), lie on the outer margins of the two orogenic domains. No major block boundary was inferred by Wellman on the junction between the Kimberley Block and the orogenic domains. Therefore, it could be inferred that the orogenic domains are the exposed upturned edge of a Lower Proterozoic crustal complex which also extends under the Kimberley Block (cf. Plumb and Gemuts, 1976). In any event the Kimberley Block has been defined by the younger episodes of deformation which affect the North Australian Platform Cover and younger platform sequences t o produce the Halls Creek and King Leopold Mobile Zones. As already noted isotopic events of approximately 1550 and 600 Ma have been tentatively identified in the King Leopold domain. During the Musgravian orogenic cycle the Halls Creek orogenic domain apparently formed a rift (the Halls Creek and Fitzmaurice Mobile Zones) in which up to 10,000 m of sediments (the Carr Boyd and Fitzmaurice Groups) were deposited. This was followed by strong NNE-trending folding and faulting, also localized in the Halls Creek domain and dated at about 1050 Ma (Plumb and Gemuts, 1976).
26
The Gawler Block
The Gawler Block is limited by thick accumulations of Adelaidean and younger platform cover related t o basement fault structures. The edges of the block are also indicated by significant gravity anomalies especially in the east and in the northwest (Fig. 1.5). The eastern boundary of the block is well known as the boundary (the Torrens Hinge Zone) between thin, little deformed, Adelaidean sediments of the Stuart Shelf and thick folded sediments of the Delamerian fold belt (Adelaide Geosyncline; Fig. 1.6). Study of the exposed basement inliers in the Delamerian fold belt shows that the folding is largely controlled by shear zones in the basement that were initiated at about 1500Ma (e.g., Glen-et al., 1977) and reactivated both during Adelaidean sedimentation and during the Delamerian orogeny (c. 500 Ma). Possibly the dipole gravity anomaly on the northwestern margin of the Gawler Block represents a similar intercratonic fold belt now obscured by younger cover of the Officer Basin (Fig. 1.6). The northeastern boundary of the Gawler Block is also marked by a WNW-trending zone of folding and faulting of Adelaidean sediments but in this case the zone may also represent an extension of the Musgravian orogenic basement. THE CENTRAL AUSTRALIAN PLATFORM COVER
The general pattern of orogenic and taphrogenic reactivation of the Proterozoic chelogen is shown in Fig. 1.6. It is notable that platform sequences belonging to the 1400-1000 Ma interval are preserved mainly on the North Australian and Gascoyne Provinces. The Roper and South Nicholson Groups (the youngest part of the North Australian Platform Cover, Fig. 1.4) can be regarded as post-orogenic sedimentation to the orogeny in the Mount Isa Subprovince and may be partly younger than 1400 Ma. The sediments of the Victoria River Basin and Bangemall Basin (on the Gascoyne Province), however, appear t o be broadly coeval with the reactivation of the chelogen along both orogenic (e.g., Musgrave) and taphrogenic (e.g., Halls Creek) zones. Similar platform sequences are not seen overlying the Gawler Province, which is largely covered by the Adelaidean ( < l o 0 0 Ma) sediments of the Central Australian Platform Cover. This sequence begins with mature blanket quartzites which unconformably overlie the Musgrave orogenic domain and which is therefore en tirely post-tectonic. Such Adelaidean sediments are preserved further north in association with the faulting of the Halls Creek and Arunta orogenic domains, but they are generally absent from the North Australian Province. Thus the Central Australian Platform Cover (G.S.A., 1971) contains two main elements. The older element (1400-1000 Ma) is coeval with basement
27
reactivation and is allocated t o the Proterozoic chelogenic cycle. Dolerite sills and dykes are widespread in the Bangemall and McArthur Basins and dolerite dyke swarms are also closely associated with emplacement of the Giles complex of mafic and ultramafic intrusions in the Musgrave domain. This major phase of mafic intrusion is taken t o mark a major phase of plate reorganization between two chelogenic cycles. The younger element of the Central Australian Platform Cover ( < l o 0 0 Ma) therefore, begins the new chelogenic cycle. Reactivation of the craton, affecting these platform sediments, occurred in the latest Precambrian and in the Palaeozoic during periods of orogenesis in the Tasman Province (Fig. 1.2). THE TRANS-AUSTRALIAN PLATFORM COVER
As a result of Mesozoic transgression, thin platform cover, developed after the stabilization of the Tasman Province, spread over much of the Precambrian shield. However, in general, the block boundaries within the shield were no longer active. The principal basins are those developed around the margins of the shield in association with the break-up and dispersal of Gondwanaland. Again however, the old block boundaries played a significant role in controlling the fragmentation. In general the rifts developed within the zones of Middle Proterozoic and younger mobility rather than in the stable blocks. SUMMARY
I t has been indicated above that the orogenic evolution of the Australian Precambrian can be discussed in terms of two chelogenic cycles. In each of the two major chelogens a number of discrete basement provinces, separated by younger basement can be distinguished. Platform cover related to the evolution of the Archaean chelogen is not present in Australia but two main stages of platform cover are related to the Proterozoic chelogen, one coeval with, and one younger than the main mobile belt sequences (Table 1.1). Two younger groups of platform cover are related to the Late PrecambrianPhanerozoic Tasman chelogenic cycle. Within the Pilbara and Yilgarn Provinces, which represent the surviving portions of the Archaean chelogen, Proterozoic events have had little effect. Shear zones were developed in the late stages of the Archaean chelogenic cycle. They are commonly NNW-trending in the Yilgarn Province while both N-S and E-W trends occur in the Pilbara. At the close of the chelogenic cycle major tensional fractures in both provinces were filled by mafic dykes. These trend ENE in the Yilgarn and mainly NNE in the Pilbara although other trends are also present. The West Australian Platform Cover rests on the Archaean chelogen but it is equivalent in age to mobile belt sequences in the Proterozoic chelogen.
28 TABLE 1.1 The Proterozoic chelogenic cyle in Australia ______ Approximate dates (Ma)
Principal tectonic stages
2400 2400
mafic magmatism and establishment of chelogenic pattern initiation of platform sedimentation (West Australian Platform Cover) and presumably of mobile belt sequences early deformation, metamorphism and plutonism in t h e Halls Creek Subprovince main deformation, metamorphism and plutonism extending across the chelogen late deformation, metamorphism and plutonism especially in the Mt. Isa Subprovince post-orogenic platform cover (North Australian) orogenic and taphrogenic reactivation of t h e chelogen especially along linear zones local platform cover especially in reactivated zones mafic magmatism and establishment of new chelogenic pattern
2000-1900
7 900-1700 1700-1 100
1900-1 4 0 0 1400-1 000 1400-1000 1050
1000
initiation of platform sedimentation of Tasman chelogenic cycle (Central Australian Platform Cover)
The Proterozoic chelogen in general appears to have been developed by reworking older Archaean crust. However, the predominant orogenic trends in the North Australian and Gawler Provinces are meridional or in the NE quadrant so that they cannot be said to mimic Archaean trends. The more nearly E-W trends of the Gascoyne Province anticipate those of the Musgrave-Fraser Province. The North Australian Platform Cover followed the stabilization of the Proterozoic chelogen and N-S-trending graben structures (e.g., Batten Trough, Figs. 1.2 and 1.4) are a notable feature. A further stage of orogenic reworking is represented by the Musgrave-Fraser Province. During this episode the principal E-W and WNW-ESE trends of the Amadeus Transverse Zone were established, but at the same time other block boundaries and various shear zone trends were established within the craton. Thus the period between 1400 and 1000 Ma v . 3 one of both orogenic and taphrogenic reworking of the pre-existing craton and a complete gradation between taphrogenic and intracontinental orogenic processes can be envisaged. At the end of the Proterozoic chelogenic cycle (c. 1100-1000 Ma) extensive basic dyke suites were again emplaced. In the Gawler Block for example a strong NW-SE trend is evident while in the Musgrave domain northerly and NE trends are prominent. The Middle Proterozoic structural trends largely controlled deposition of the main part of the Central Australian Platform Cover, younger than 1000
29 Ma, which belongs t o the Tasman chelogenic cycle. The Middle Proterozoic structural trends were reactivated during the deposition of this platform cover, especially in the latest Precambrian and at the end of deposition, in the Middle Palaeozoic. This latter period corresponds with the main orogenic activity in the Lachlan Subprovince of the Tasman chelogen. It is evidently commonplace for tectonic activity within the Precambrian craton to reactivate older tectonic elements. The tectonic pattern of the Proterozoic chelogen is largely due to reactivation, during the Middle Proterozoic and younger episodes, of earlier Proterozoic orogenic trends. This phenomenon gives some justification to concepts of a fundamental rhegmatic pattern in Australia (e.g., Hills, 1956; O’Driscoll, 1971). However, there is no obvious relationship between the Proterozoic trends and older Archaean trends. This is particularly evident for Proterozoic trends peripheral to the Pilbara and Yilgarn Provinces. Consequently, although Proterozoic orogeny is largely ensialic, the main Early Proterozoic orogenic trends can probably be attributed to the operation of a new plate system rather than an inherited tectonic pattern. It has been indicated that most Archaean and Proterozoic orogenies have analogy with Phanerozoic orogeny and can be interpreted in the context of chelogenic cycles and plate tectonics. Although orogenic belts were broad during major orogenic phases, and largely ensialic, they were probably developed in back-arc environments adjacent to continental margins (Rutland, 1973b). An exception is provided by the Musgrave-Fraser system which belongs t o a distinctive class of intracontinental reactivation orogeny (Rutland, 1973b, 1976). This reactivation orogeny can be regarded as the most extreme form of basement reactivation which elsewhere is expressed in taphrogenic activity, affecting the platform cover development. REFERENCES Anfilorf, W. and Shaw, R.D., 1973. T h e gravity effects of three large uplifted graiulite blocks in separate Australian shield areas. In: R.S. Methur a n d P.V. Angus-Leppan (Editors), Proc. Symp. Earth’s Gravitational Field and Secular Variations in Position. Svdnev. PP. 273-289. Anfiloff, W . , Barlow, B.C., Murray, A,, Denham, D. and Sandford, R., 1976. Compilation and production of the 1 9 7 6 1 : 5,000,000 Gravity Map of Australia. BMR J. Aust. Geol. Geophys., 1: 273-276. Arriens, P., 1971. T h e Archaean geochronology of Australia. Spec. Publ. Geol. SOC.Aust., 3: 11-23. Blake, D.H., 1978. T h e Proterozoic a n d Palaeozoic rocks of t h e Granites-Tanami region, Western Australia and Northern Territory and Interregional correlations. BMR J. Aust. Geol. Geophys., 3: 35-42. Bofinger, V.M., 1 9 6 7 . Geochronology in t h e East Kimberley Area of Western Australia. Ph.D. Thesis, Australian National University (unpubl.). Brakel, A.T. and Muhling, P.C., 1976. Stratigraphy sedimentation a n d structure i n t h e western and central part of t h e Bangemall Basin, Western Australia. West. Aust. Geol. Surv. Annu. Rep., 1 9 7 5 : 70-79.
30 Brewer, J.A., Brown, L.D., Cook, F.A., Kaufman, S., Oliver, J.E. and Smithson, S.B., 1978. The nature of the Wind River Thrust from Cocorp deep reflection data. Geol. SOC.Am., Abstr. with Programs, 1 0 (7): 371. Bunting, J.A., de Laeter, J.R. and Libby, W.G., 1976. Tectonic subdivisions and geochronology of the northeastern part of the Albany-Fraser province, Western Australia. West. Aust. Geol. Surv. Annu. Rep., 1975: l 17-1 26. Bunting, J.A., Commander, D.P. and Gee, R.D., 1977. Preliminary synthesis of Lower Proterozoic stratigraphy and structure adjacent t o the northern margin of the Yilgarn Block. West. Aust. Geol. Surv. Annu. Rep., 1976: 43-48. Cooper, J.A. and Compston, W., 1971. Rb-Sr dating within the Houghton Inlier, South Australia. J. Geol. SOC.Aust., 1 7 ( 2 ) : 213-219. Cooper, J.A , Fanning, C.M., Flook, M.M. and Oliver, R.L., 1976. Archaean and Proterozoic metamorphic rocks on southern Eyre Peninsula, South Australia. J. Geol. SOC. Aust., 23: 287-292. Daniels, J.L., 1975. Gascoyne province. In: Geology of Western Australia. West. Aust. Geol. Surv. Mem., 2: 107-114. Davies, F.B. and Windley, B.F., 1976. Significance of major Proterozoic high grade linear belts in continental evolution. Nature, 263: 383-385. Dennis. J.G., 1967. International Tectonic Dictionary. Am. Assoc. Pet. Geol. Mem., 7: 196 pp. Dewey, J.F. and Burke, K., 1974. Hot spots and continental break-up: Implications for collisional orogeny. Geology, 2: 57-60. Dooley, J.C., 1976. Variation of crustal mass over the Australian region. BMR J. Aust. Geol. Geophys., 1: 291-296. Dooley, J.C., 1977. Implications of Australian seismic and gravity measurements f o r the structure and composition of the upper mantle. BMR J. Aust. Geol. Geophys., 2: 1-5. Dunn, P.R., Plumb, K.A. and Roberts, H.G., 1966. A proposal for time-stratigraphic subdivision of the Australian Precambrian. J. Geol. Soc. Aust., 1 3 : 5 9 3 4 0 8 . Embleton, B.J.J., 1978. The palaeomagnetism of 2400 Ma rocks from the Australian Pilbara craton and its relation to Archaean-Proterozoic tectonics. Precambrian Res., 6: 275-291. Gee, R.D., de Laeter, J.R. and Drake, J.R., 1976. Geology and geochronology of altered rhyolite from the lower part of the Bangemall Group near Tangadee, Western Australia. West. Aust. Geol. Surv. Annu. Rep., 1975: 112-117. Gellatly, D.C., 1971. Possible Archaean rocks a t the Kimberley region, Western Australia. Spec. Publ. Geol. SOC.Aust., 3: 93-101. Geological Society of Australia, 1971. Tectonic Map of Australia and New Guinea 1 : 5,000,000. Sydney. Glen, R.A., Laing, W.P., Parker, A.J. and Rutland, R.W.R., 1977. Tectonic relationships hetween the Proterozoic Gawler and Willyama orogenic domains, Australia. J. Geol. SOC.Aust., 24: 125-150. Harland, W.B., 1975. The two geological time scales. Nature, 253: 505-507. Hills, E.S., 1956. The tectonic style of Australia. Geotekt. Symp. zu Ehren von H. Stille. Dtsch. Geol. Ges., Hannover, pp. 336-346. Hoffman, P., 1973. Evolution of an early Proterozoic continental margin: the Coronation geosyncline and associated aulacogens of the northwestern Canadian Shield. Philos. Trans. R. SOC.London, Ser. A, 273: 547-581. Horwitz, R.C. and Smith, R.E., 1978. Bridging the Yilgarn and Pilbara Blocks, Western Australia. Precambrian Res., 6: 293-322. De Laeter, J.R., 1976. Rb-Sr whole rock and mineral ages from the Gascoyne province. West. Aust. Geol. Surv. Annu. Rep., 1975: 126-130. De Laeter, J.R., Hickman, A., Trendall, A.G. and Lewis, J.D., 1977. Geochronological
31 data concerning the eastern extent of the Pilbara Block. West. Aust. Geol. Surv. Annu. Rep., 1976: 56-62. King, P.B., 1969. The tectonics of North America - A discussion to accompany the Tectonic Map of North America, Scale 1 : 5,000,000. U.S. Geol. Sum. Prof. Pap. 628, 9-1 pp. Kriiner, A , , 1977. The Precambrian geotectonic evolution of Africa: plate accretion versus plate destruction. Precambrian Res., 4:163-213. Krliner, A. and Blignault, H.J., 1976. Towards a definition of some tectonic and igneous provinces in western South Africa and southern South West Africa. Trans. Geol. SOC. S. Afr., 79: 232--238. Marjoribanks, R.W. and Black, L.P., 1974. Geology and geochronology of the Arunta Complex, north of Ormiston Gorge, Central Australia. J. Geol. SOC.Aust., 21: 291-300. Mathur, S.P., 1976. Relation of Bouguer anomalies to crustal structure in central and southwestern Australia. BMR J. Aust. Geol. Geophys., 1: 277~-286. McDougall, J., Dunn, P.R., Compston, W., Webb, A.W., Richards, J.R. and Bofinger, V.M., 1965. Isotope age determination on Precambrian rocks of the Carpentaria Region, Northern Territory, Australia. J. Geol. SOC.Aust., 1 2 : 67-90. Moore, A.C. and Goode, A.D.T., 1978. Petrography and origin of granulite-facies rocks in the Western Musgrave Block, Central Australia. J. Geol. SOC.Aust., 25: 341-358. Nesbitt, R.W., Goode, A.D.T., Moore, A.C. and Hopwood, T.P., 1970. The Giles Complex, Central Australia: a stratified sequence of mafic and ultramafic intrusions. Geol. SOC.S. Afr. Spec. Publ., 1: 547-564. O’Driscoll, E.S.T., 1971. Deformational concepts in relation t o some ultramafic rocks in Western Australia. Spec. Publ. Geol. SOC.Aust., 3 : 351-366. Page, R.W., 1976. Response of U-Pb zircon and Rb-Sr total-rock systems to low-grade regional metamorphism in Proterozoic igneous rocks, Mount Isa, Australia. Annu. Rep. Director Geophys. Lab., Carnegie Inst. Washington, Yearb., pp. 813-821. Pidgeon, R.T., 1978. 3450 m.y. old volcanics in Archaean layered greenstone succession of the Pilbara Block, Western Australia. Earth Planet. Sci. Lett., 37: 421-428. Plumb, K.A., 1976. The Precambrian shields and platforms of northern Australia. Bur. Miner. Resour. Aust. Rec. 1976/72 (unpubl.). Plumb, K.A. and Derrick, G.M., 1975. Geology of the Proterozoic rocks of the Kimberley to Mt. Isa region. In: C.L. Knight (Editor), Economic Geology of Australian and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5: 217-252. Plumb, K.A. and Gemuts, I., 1976. Precambrian geology of the Kimberley region, Western Australia. 25th Int. Geol. Congr., Excursion Guide No. 44C, Sydney, 69 pp. Rutland, R.W.R., 1973a. On the interpretation of Cordilleran orogenic belts. Am. J. Sci., 273: 811---849. Rutland, R.W.R., 1973b. Tectonic evolution of the continental crust of Australia. In: D.H. Tarling and S.K. Runcorn (Editors), Continental Drift, Sea Floor Spreading and Plate Tectonics: Implications to the Earth Sciences. Academic Press, London, pp. 1003-1025. Rutland, R.W.R., 1976. Orogenic evolution of Australia. Earth-Sci. Rev., 1 2 : 161-196. Rutland, R.W.R., 1977. Proterozoic Platforms and Mobile Belts (Abstr.). 48th ANZAAS Congr., Melbourne. Sawkins, F.J., 1976. Widespread continental rifting. Geology, 4 : 427-430. Stewart, A.J. and Warren, R.G., 1977. The mineral potential of the Arunta Block, central Australia. BMR J. Aust. Geol. Geophys., 2: 21-34. Stockwell, C.H., 1964. Fourth report on structural provinces, orogenies, and time-classification of rocks of the Canadian Precambrian Shield. Can. Geol. Surv. Pap. 64-17, part. 11, pp. 1-21.
32 Stockwell, C.H., 1966. Notes on the tectonic map of the Canadian Shield. In: Scientific communications read to the Commission for the Geological Map of the World. 22nd Int. Geol. Congr. India, 1964, pp. 33-40. Sutton, J., 1963. Long-term cycles in the evolution of the continents. Nature, 198: 731735. Sutton, J., 1967. The extension of the geological record into the Precambrian. Proc. Geol. ASSOC.,78: 493-534. Thomas, M.D. and Gibb, R.A., 1977. Gravity anomalies and deep structure of the Cape Smith foldbelt, northern Ungava, Quebec. Geology, 5: 169-172. Thomson, B.P., 1976. The Musgrave Block-Regional Geology. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5: 451-454. Traves, D.M., 1955. The geology of the Ord-Victoria region, northern Australia. Bur. Miner. Resour. Aust. Bull. 27. Trendall, A.R., 1975. Introduction. In: Geology of Western Australia. West. Austr. Geol. Surv., Mem., 2: 25-32. Trendall, A.R., 1978. Basins of Iron-formation deposition (Abstr.). Metallogeny of the Baltic Shield, Helsinki Symposium 1978. Int. Geol. Correlation Progr. Project No. 91. Warren, R.G., 1978. Delny-Mount Sainthill Fault System, eastern Arunta Block, Central Australia. BMR J. Aust. Geol. Geophys., 3: 76-79. Webb, A.W. and Horr, G., 1978. The Rb-Sr age and petrology of a flow from the Beda Volcanics. South Aust. Geol. Surv., Q. Geol. Notes, No. 66. Wellman, P., 1976a. Gravity trends and the growth of Australia: a tentative correlation. J. Geol. SOC.Aust., 23: 11-14. Wellman, P., 197613. The gravity field of the Australian basement. BMR J. Aust. Geol. Geophys., 1: 287-290. Wellman, P., 1976c. Regional variation of gravity and isostatic equilibrium of the Australian crust. BMR J. Aust. Geol. Geophys., 1: 297-302. Wellman, P., 1978. Gravity evidence for abrupt changes in mean crustal density a t the junction of Australian crustal blocks. BMR J. Aust. Geol. Geophys., 3: 153-162. Williams, T.R., Brakel, A.T., Chin, R.J. and Williams, S.J., 1976. The stratigraphy of the eastern Bangemall Basin and the Paterson Province. West. Aust. Geol. Surv. Annu. Rep., 1975: 79-83.
Note added in p r o o f The text of this chapter was largely written in 1978. Since then significant general papers on the structural framework of the Australian Precambrian have been published including the following: Gee, R.D., 1979. Structure and tectonic style of the Western Australian Shield. Tectonophysics, 58: 327-369. Plumb, K.A., 1979. Structure and tectonic style of the Precambrian shields and platforms of northern Australia. Tectonophysics, 58: 291-325. Plumb, K.A., 1979. The tectonic evolution of Australia. Earth-Sci. Rev., 1 4 : 205-249.
33
Chapter 2
ARCHAEAN GRANITE-GREENSTONE TERRANES O F WESTERN AUSTRALIA J.A. HALLBERG and A.Y. GLIKSON
INTRODUCTION
The principal concentration of rocks of Archaean * age within Australia is found in the Western Australian Shield (or Western Shield), as defined by Trendall (1975), where such rocks occupy an area in excess of 700,000 km2. Within this shield two Archaean cratonic nuclei are recognized - the Yilgarn and Pilbara Blocks (Fig. 2.1). The Yilgarn Block covers a rectangular area of about 650,000 km2 within the coordinates 25”-35” S and 115”---124”E- roughly the southwestern part of the Australian continent. Most of this region is characterized by low rainfall (200-250 mm annually), poorly developed ephemeral drainages, sinuous systems of elongate, salt-encrusted dry “lakes”, and large areas of sand plain and laterite. The latter is a remnant of a Jurassic or Tertiary erosional surface - the Old Plateau (Jutson, 1934). In the area of exterior drainage to the southwest of the Yilgarn Block, rainfall may exceed 1000 mm per annum and permanent rivers (Murray, Swan, Blackwood) present greater dissection. The main population centres in the interior of the Yilgarn Block are either present-day mining centres (Kambalda, Kalgoorlie, Norseman, Laverton) or relics of earlier gold mining days which have remained as pastoral centres (Cue, Yalgoo, Leonora, Meekatharra). The more densely populated southwest corner of the Yilgarn Block contains many well-established agricultural centres. Geologically, the Yilgarn Block is bounded on the west by the Darling Fault which separates it from the Perth Basin - a graben-like structure filled with some 15 km of Palaeozoic to Tertiary sediments. To the east the Yil-
* The Achaean-Proterozoic boundary, as referred t o in this paper, is defined at 2.6 Ga - an age representing a global thermal rise (Dearnley, 1966), rather than on the onset of continental platform-type volcanism and sedimentation.
34
WESTERN PRECAMBRIAN SHIELD AND MARGINS
Fig. 2.1. Map of part of Western Australia showing t h e location of t h e Yilgarn and Pilbara Blocks.
garn Block is covered by Permian t o Mesozoic sediments of the Officer Basin. The north and northeast portions of the Yilgarn Block are unconformably overlain by Proterozoic sediments of the Nabberu Basin. The
35 south-southeastern and northwestern comers of the Yilgarn Block are bounded by mobile belts of Proterozoic age. The Albany-Fraser Province (or mobile zone) lies t o the south-southeast and the Gascoyne Province t o the northwest. The latter may have been connected with high-grade metamorphic Proterozoic inliers in the Perth Basin. The Pilbara Block lies to the north of the Yilgarn Block and occupies an elongate triangular area of 56,000 km2 within the coordinates 20"-22"s and 116"30'-120"30'E (Fig. 2.1). The terrain is bounded by the northern edge of the Chichester Range to the south, the Gregory and Throssell Ranges to the east, the Canning Desert t o the northeast and the Indian Ocean to the northwest. The limits are also broadly defined by the Fortescue River t o the south and the Oakover River and De Grey River to the east and north, respectively. Most of the principal drainages within the Archaean terrane flow north into the De Grey River in the eastern Pilbara, and northwest into the Indian Ocean in the western Pilbara. None of these is perennial, the climate being arid with an average annual rainfall of 300 mm, concentrated in the summer months when the region is affected by cyclones. Principal population centres are Port Hedland, Dampier (iron-ore ports), Mount Goldsworthy, Shay Gap (iron-ore mining centres), Marble Bar and Nullagine (old gold mining centres dating to the last century), Karratha, Wickham and Roebourne. The region contains numerous small mining camps and sheep and/or cattle stations. Geologically the Archaean craton is bounded by outcrops of the Lower Proterozoic Mount Bruce Supergroup to the south, Middle Proterozoic deposits of the Patterson Province to the east, and Jurassic to Cretaceous sediments of the Canning Basin to the north. The Archaean terrane consists of: (1) large oval granitoid batholiths exposed as bouldery outcrops and tors over extensive plains; (2) intervening hilly belts and rugged narrow ridges between 20@-500 m above sea level formed by Archaean volcanic rocks and interbedded jaspilites respectively; ( 3 ) massive, rounded, relatively elevated hills formed by quartzites and banded iron-formation; and (4) dissected plateaus and low table-top hills formed by flat-lying volcanic and sedimentary outliers of Lower Proterozoic rocks. Apart from the Yilgarn and Pilbara Blocks the only other area in which rocks older than 2.6 Ga have been isotopically dated is the Gawler Block, South Australia (Cooper et al., 1976a). It is likely, however, that earliest Proterozoic ages yielded by gneisses in other parts of the Gawler Block (Webb and Thomson, 1977), Alligator River-Rum Jungle area, Northern Territory (Page, 1976) and the Gascoyne Province (De Laeter, 1975) are imprinted on rocks of Archaean age. Cooper et al. (1976a) recorded whole-rock Rb-Sr isochron ages of 2643 134 Ma on augen and layered gneisses from the Cape Carnot area, southern Eyre Peninsula, South Australia, which rocks also include biotite-garnet gneiss and basic granulite. A poorly defined age of 2837 5 134 Ma was ob-
*
36 tained from leuco-gneiss from this area. Gneisses, yielding earliest Proterozoic ages from the southwestern part of the Eyre Peninsula (Whidbey Granite, 2407 k 50 Ma, initial 87Sr/86Srratio ( R i ) of 0.7029) and the Lake Harris-Tarcoola area of the Gawler Block (Glenloth Granite, 2350 f 33 Ma, Ri of 0.7028) (Webb and Thomson, 1977), contain bodies of amphibolite which, together with low R i values, suggest possible derivation from Archaean granitegreenstone associations. In the Rum Jungle-Alligator River area, granitic gneiss and migmatites, which form a basement t o sediments of the Pine Creek Geosyncline, have yielded Rb-Sr and U-Pb zircon ages older than 2400 Ma (Richards et al., 1966; Page, 1976). Samples collected from the southern part of the Gascoyne Province in Western Australia have yielded a Rb-Sr isochron age of c. 2208 Ma with a high Ri of 0.7136, suggesting a derivation from older Archaean basement (De Laeter, 1975). Rocks of presumed Archaean age are also exposed in the cores of anticlines in Proterozoic sedimentary basins in Western Australia (Sylvania, Malmac and Rocklea Domes; Wyloo, Rat Hill and Milli Milli Anticlines) (Honvitz and Smith, 1978; Fig. 2.1). The above occurrences, while highly significant in determining the original extent of the Archaean crust, have t o date yielded little information on the original nature of such a crust and will not be further discussed in this chapter. Geological studies in Archaean terranes of Western Australia were greatly stimulated by a “mining boom” in the late 1960’s-early 1970’s. Further stimuli were provided by the Symposium on Archaean Rocks held in Perth in 1970 and by the regional mapping program of the Geological Survey of Western Australia. Such studies culminated in the presentation of an abundance of critical data in the period 1971-1976. The present period is one of reassessment and of gathering further fundamental data relevant to enable a complete understanding of the Archaean areas and their relationships. This review outlines the development of concepts on the stratigraphy and structure of the Archaean granite greenstone terranes and attempts t o emphasize recent interpretations, correlations and information provided by regional studies by various research institutes, Commonwealth and State geological organizations and mining companies. Y ILGARN BLOCK
Subdivision A gross subdivision of the Yilgarn Block based primarily on tectonic style and lithological association (Trendall, 1975) recognizes three major provinces: Southwestern, Murchison and Eastern Goldfields (Fig. 2.2A). Boundaries between the provinces do not, in some instances, appear to represent chronological or tectonic discontinuities. Greenstone belts of the Murchison and Eastern Goldfields Provinces are essentially similar in lithol-
37
ogy and age, and subdivision into these two provinces is based on a change in dominant structural trend from N to NNW in the Eastern Goldfields Province t o NE in the Murchison Province, and the arcuate discontinuous nature of greenstone belts in the Murchison Province as opposed to the more continuous, linear belts of the Eastern Goldfields Province. The Southwestern Province is distinct in containing high-grade metamorphic supracrustal sequences and in yielding the oldest dates obtained from the Yilgarn Block. The Eastern Goldfields Province has been further subdivided into three subprovinces on the basis of structural discontinuities and rock type distribution (Williams, 1974). These subdivisions - the Southern Cross, Kalgoorlie and Laverton Subprovinces - are shown in Fig. 2.2A. Comprehensive source references and general background for the Yilgarn Block are provided by two recent reviews by Gee (1975) and Binns and Marston (1976). Stratigraphy and rock associations The widespread occurrence of surficial cover in the Southwestern Province and the lack of any published recent comprehensive regional study of the area preclude any detailed discussion in this chapter. However, it should be noted that mapping of the area is being presently carried out by the Geological Survey of Western Australia (Wilde, 1976). Briefly, supracrustal sequences in the Southwestern Province occur as high-grade metamorphic remnants of mafic to ultramafic igneous rocks, and clastic and chemical sedimentary rocks within ortho- and para-gneisses. Such remnants comprise both small, diffuse enclaves and more widespread schist-gneiss belts. Metamorphic grade ranges from middle amphibolite t o granulite facies. Continental and epicontinental sediments are widely represented in the sedimentary relics (Gee, 1975). As further discussion will centre on granite-greenstone sequences within the remainder of the Yilgarn Block, readers are referred to the reviews by Gee (1975) and Binns and Marston (1976) for further information on, and source references to, the high-grade Southwestern Province. Early mapping in the Yilgarn Block distinguished some sedimentary and volcanic series and their possible relationships (cf. Ellis, 1939) and emphasized the presence of pre- and post-folding mafic intrusions (the “older” and “younger” greenstones of Prider, 1948, 1965). Local stratigraphic successions in the vicinity of the major mining centres of Norseman (Hall and Bekker, 1965), Kalgoorlie (Woodall, 1965), Kambalda (Ross and Hopkins, 1975) and Mt. Magnet (Forman, 1960) have been established. Detailed studies of larger areas by McCall (1969), Glikson (1971a, b) and Hallberg et al. (1976b) have yielded stratigraphic information. Stratigraphic terminology has been applied within various 1 : 250,000 map sheets in both the Eastern Goldfields and Murchison Provinces by the Geological Survey of Western
38
N
Granitoid
,TJ Kurrawang Group
--
a * o a *
Kalpini Group Gundockerta Group Major felric centres in Gindalbie Group
w-
,c--’
Morelands Gindalble and Mulgabbie Groups
Fault I lineament
KILOMETRES
50
OF -PI
THE AREA
(1976)
,-’ Fault I lineament KILOMETR ES
3o
STRATIGRAPHY OF THE COOLGARDIE.NORSEMAN AREA AFTER GEMUTS AND THERON 11975)
Fig. 2.2. Subdivision and stratigraphy in the Yilgarn Block. A. Distribution of greenstone belts and subdivision after Trendall (1975) and Williams (1974). B. Stratigraphic interpretation of the Edjudina-Kurnalpi area after Williams (1976). C. Stratigraphic interpretation of the Kalgoorlie-Norseman area after Gemuts and Theron (1975). Lithological descriptions of stratigraphic units in B and C are given in Table 2.1.
TABLE 2.1 Comparison of various stratigraphic sequences from the Eastern Goldfields and Murchison Provinces. Modified from Gemuts and Theron (1975) Gemuts and Therm ( 1 9 7 5 ) CoolgardieNorseman area
Williams ( 1976) Southeastern portion of Eastern Goldfieds
Mc Ca ll(l969) Lake Lefroy-Kambalda area
Glikson (1971a) CoolgardieMungari area
Sequence 8 polymictic conglomerate and pebbly greywacke
Kurrawang Group epiclastic sediments mainly conglomerate some felsic intrusive rocks
Kurrawang Conglomerate
Kurrawang Beds
volcaniclastic sediments, acidintermediate volcanic rocks, conglomerate, greywacke, black shale
Sequence 7 acid tuffaceous rocks and acid volcanic breccia, some acid extrusive rocks
Kalpini Group basic to intermediate extrusive and intrusive rocks, ultramafic intrusive rocks, minor clastic rocks and chert
Mungari Beds Black Flag Metasediments
tholeiitic basalts, high -Mg basalt , podiform ultrabasic bodies, black shale, acid tuff
Sequence 6 high-Mg basalt, ultramafic rocks, minor chert, black slate and tholeiitic basalt Sequence 5 conglomerate, arkosic greywacke and argillite with minor tholeiitic basalt, basal conglomerate
Gundockerta Group turbidite sequence, clastic sequence, conglomerate and sandstone shallow and deep water facies
Yilmia Ophiolite belt Cave Rocks Ophiolite belt
Red Lake Ophiolites
Lake Dam Beds Wanda Wanda Beds Cave Rock Beds Mandilla Beds Triangle Islands Beds Merougil Creek Beds
Brown Lake Metasediments Mt. Robinson Ophiolites
Hall and Bekker (19 6 5) Norseman region
Hallberg ( 1976) Murchison F’rovince
shale, siltstone, basic sediments, podiform ultrabasic bodies, banded iron-formation, quartzsericite schist, acid tuff, quartzite
+ 0
TABLE 2.1 (continued)
Sequence 4 t o p of sequence gradational t o sequence 5. Greywacke, minor chert and black slate, acid extrusive rocks and feldspar-porphyry intrusives, minor tholeiitic basalt
acid volcanic complexes
Causeway Beds
Gunga Meta-argillites
Mt. Thirsty Beds grey wacke, slate, acid tuff tholeiitic basalt quartzite, chert Abbotshall Beds chert with interbedded basalt, acid volcanic and sedimentary rocks
Sequence 3 black chert marker, tholeiitic basalt, 4 horizons of ultramafic rocks, intercalated slate and chert bands, minor high-Mg basalt
Mulgabbie Group basic intrusive and extrusive rocks, ultramafic rocks. intermediate and acid extrusive rocks and minor cherts
Coolgardie Ophiolites Woolyeenyer Group tholeiitic pillow lavas, slate, minor high-Mg basalt
Sequence 2 equivalent and exactly the same as the Noganyer Group
Gindalbie Group clastic sequence, acid volcanic complexes
Noganyer Group banded iron-formation, greywacke, shale, conglomerate sandstone
Sequence I equivalent and exactly the same as Penneshaw Beds
Morelands Group tholeiitic and ultramafic lavas
Penneshaw Beds greywacke and acid lithic tuff, tholeiitic basalt
ortho-amphibolite, para-amphibolite, ultramafic schists, banded iron-formation, quartzsericite schist, shale, siltstone, acid tuff
A
w
42
Fig. 2.3. Landsat image of Archaean granite-greenstone terrane about Laverton, Yilgarn Block. Granite, greenstone belts, ephemeral drainages and salt lakes are clearly shown.
43 Australia (Williams, 1970; De la Hunty, 1973; Muhling and Low, 1973). Regional stratigraphic syntheses centred on the Kalgoorlie area have been put forward by Williams (1970,1975) and Gemuts and Theron (1975) and a preliminary stratigraphy has been outlined for the Murchison Province by Hallberg (1976). A stratigraphic interpretation of the entire southeastern part of the Yilgarn Block is given by Williams (1976). A summary of the principal regional stratigraphic interpretations proposed t o date is given in Table 2.1 and two examples are shown in Fig. 2.2B, C. Both Williams (1969, 1975) and Gemuts (1976) have recognized cyclic repetition of the sequence (mafic/ultramafic volcanic rocks)-(felsic volcanic rocks/volcaniclastic sedimentary rocks)-(clastic sedimentary rocks) in the Kalgoorlie area. Williams documented three complete cycles and Gemuts two complete cycles and possible portions of a third. The regional distribution of, and correlation between, such cycles is not known, but it is likely that they are not represented in some areas. The repetitive cycles appear t o be more characteristic of the Kalgoorlie Subprovince than of adjacent areas. It is evident from Table 2.1 that some elements of the regional stratigraphies are broadly compatible. Some authors have suggested that regional correlation between supracrustal sequences may be possible and Gemuts (1976) has proposed the concept of spatial and temporal stratigraphic continuity throughout the Yilgarn Block. However, it is most difficult to correlate individual sequences between widely spaced greenstone belts from existing information. Furthermore, tectonic slicing combined with very poor outcrop (Fig. 2.3) and limited understanding of regional structure often inhibit correlation even within individual belts. Despite these problems, the development of stratigraphic concepts is essential, as it is through such concepts that the continuity and distribution of various rock associations both in space and in time can be investigated. It is felt that a stratigraphic framework offers the best means of presenting relevant data on rock type distribution in greenstone terranes of the Murchison and Eastern Goldfields Provinces. Only the broadest elements of stratigraphy, as shown in Table 2.1, will be used in the ensuing discussion. The limitations mentioned above and those imposed by the scarcity of published information from areas outside the Kalgoorlie Subprovince must be borne in mind.
Lo wer maf ic/ul t rama fic asso cia tio ns The lower portions of recognized stratigraphic successions in both the Murchison and Eastern Goldfields Provinces are dominated by piles of mafic/ ultramafic lavas and associated comagmatic hypabyssal to shallow-level intrusive rocks. Felsic volcanic/volcaniclastic sedimentary and sedimentary units are invariably present and may be locally important. In no recorded instance have the lower members of stratigraphic successions been found t o unconformably overlie older granitoid rocks. Mafic volcanic rocks are found in lower stratigraphic positions in both the
44 Murchison and Eastern Goldfields Provinces. Stratigraphically low units of mafic volcanic rocks often contain abundant units of pillow basalts; amygdaloidal, variolitic and hyaloclastic textures are often preserved. A majority of these basalts can be classified as low-K, quartz-normative tholeiites (Glikson, 1 9 7 1 ~ ;Hallberg, 1972a, 1976; Hallberg and Williams, 1972). Lower sequences of mafic lavas in the Kalgoorlie Subprovince are significantly TABLE 2.11 Analysis of mafic volcanic rocks from lower mafic/ultramafic association
4
3
2
1
5
~~ ~
n
SiOz A1203
Fez03 FeO MgO CaO NazO KZO H20t C02 TiOz p205
MnO Ba Cr cu Nb Ni Rb Sr V Y Zn
Zr
429
133
103
66
51.2 (2.3) 14.5 (1.4) 2.2 (0.9) 8.8 (1.5) 6.9 (1.7) 10.8 (1.6) 2.5 (0.8) 0.25(0.19) 1.8 (0.8) 0.1 (0.4) 0.86(0.32) 0.1 2( 0.06) O.Zl(0.05)
51.9 (2.2) 14.1 (1.6) 2.0 (0.7) 8.6 (1.8) 7.4 (1.6) 10.4 (1.6) 2.3 (0.8) 0.20(0.22) 1.6 (0.7) 0.2 (0.7) 0.87( 0.36) 0.13(0.09) 0.21(0.14)
51.8 (2.0) 14.0 (1.1) 2.5 (0.8) 8.1 (1.5) 7.3 (1.5) 10.5 (1.7) 2.4 (0.8) 0.22(0.21) 1.9 (0.7) 0.1 (0.5) 0.82(0.35) O.ll(0.06) 0.20(0.04)
52.2 (1.6) 13.7 (1.2) 2.2 (0.6) 8.2 (1.2) 7.8 (1.4) 10.7 (1.3) 2.2 (0.8) 0.17(0.18) 1.5 (0.5) 0.1 (0.1) 0.79(0.28) O.lO(0.05) 0.20(0.04)
103(122) * 321(214) 107 (75) 4 (5)* 157 (48) 10 (11) 134 (71) 289 (79) 21 (7) 9 3 (27) 62 (27)
105(173) 312( 173) 69 (52) 4 (5) 1 2 9 (45) 6 (10) 143 (14) 277 (53) 22 (18) 8 0 (27) 62 (33)
116(137) 292(217) 7 5 (58) 2 (3) 1 2 6 (42) 7 (9) 1 2 3 (50) 281 ( 6 0 ) 22 (9) 72 (22) 6 1 (39)
116( 188) 315(179) 64 ( 4 8 ) 3 (5) 125 (45) 5 (7) 142(117) 282 (44) 2 3 (20) 7 5 (26) 62 (35)
195 51.3 (1.6) 14.7 (1.1) 1.6 (0.5) 9.0 (1.2) 6.9 (0.9) 10.6 (1.0) 2.6 (0.5) 0.18(0.05) 1.4 (0.8) 0.1 (0.2) 0.91( 0.18) 0.13(0.05) 0.21( 0.02) 398( 102) 1 0 5 (82) 163 10 121 301 20 98 61
(25) (12) (32) (38) (4) (19) (11)
1 = average and standard deviation for lower to middle greenschist facies mafic/volcanic rocks (Murchison and Eastern Goldfields Provinces); 2 = average and standard deviation for upper greenschist facies t o amphibolite facies mafic/volcanic rocks (Murchison and Eastern Goldfields Provinces); 3 = average and standard deviation for lower t o middle greenschist facies mafic/volcanic rocks from the Murchison Province and Southern Cross Subprovince; 4 = average and standard deviation f o r upper greenschist t o amphibolite facies mafic/ volcanic rocks from t h e Murchison Province and Southern Cross Subprovince; 5 = average and standard deviation for lower to middle greenschist facies mafic/volcanic rocks from t h e Kalgoorlie Subprovince. * Nh and Ba analyses o n only 200 samples.
45
thicker and more continuous than those in the rest of the Yilgarn Block. The thicker basalt piles of the Kalgoorlie Subprovince, such as the Norseman basalts, display a monotonous chemistry characterized by high Cr, Ni and Fe/Mg and low Ti and Al. In many respects they are similar in composition to ocean floor basalts (Glikson, 1 9 7 1 ~Hallberg, ; 1972a; Pearce, 1976; Winchester and Floyd, 1976). The thinner, lower mafic volcanic sequences in the Southern Cross and Laverton Subprovinces and Murchison Province are much more variable in composition and contain a higher proportion of more evolved tholeiitic basalts enriched in Fe, Ti and Zr and depleted in Cr and Ni (Hallberg, 1976). Individual units of basalt in the Murchison Province and Southern Cross Subprovince can be highly differentiated (Hallberg et al., 1976a). Analysis of basalts from lower stratigraphic sequences from various portions of the Yilgarn Block are presented in Table 2.11. Ultramafic volcanic rocks and related dykes and sills are found in low stratigraphic positions in both the Murchison and Eastern Goldfields Provinces. In the Kalgoorlie Subprovince, well-preserved ultramafic rocks range from peridotite flows (cf. Barnes et al., 1974) through magnesian basalts t o more evolved basaltic members whose genetic relationships with associated tholeiitic basalts are little known at this time. The extrusive ultramafic rocks characteristically display platy, skeletal and arborescent textures which have been attributed t o quenching (Lewis, 1971; Nesbitt, 1971). Pillow structures may be developed in the more mafic lavas, and interflow hyaloclastic breccias are associated with sequences of peridotite flows. Spilitized high-Mg basalts have been reported from the Mt. Hunt area, 1 9 km south of Kalgoorlie (Hallberg, 197213). Within the Kalgoorlie Subprovince, ultramafic volcanic rocks are found as horizons within mafic sequences and as thick accumulations of greater lateral extent which may possibly represent volcanic centres associated with fundamental fracture systems. In the Murchison Province and Southern Cross Subprovince, stratigraphically low ultramafic units have often suffered penetrative deformation and textural obliteration. The bulk of published chemical data on ultramafic rocks of the Yilgarn Block pertains t o lower stratigraphic sequences in the Kalgoorlie Subprovince. The chemical diversity of these rocks and uncertainties in their field relationships render difficult an elucidation of genetic affinities between the different varieties even within a given locality. Relationships of the ultramafic volcanic rocks t o associated mafic tholeiites are likewise unclear. However, some data suggest that the two groups are genetically unrelated. High-Mg basalts generally display a greater compositional diversity than associated tholeiitic basalts both within and between localities (Hallberg and Williams, 1972). Their compositions may show evolved characteristics not associated with sequences of tholeiitic basalt (Hallberg and Williams, 1972; Hallberg, 1976). Nesbitt and Sun (1976) on the basis of chemical composition (V, Y, TiOz, MgO), textures and field association suggest that the voluminous tholeiites are primary partial melts and
46 could not be differentiates of the high-Mg basalt suite. Fine-grained chemical and clastic sedimentary rocks are an integral component of mafic/ultramafic volcanic associations in lower stratigraphic positions. Whilst banded iron-formation is conspicuous in mafic/ultramafic sequences in most of the Murchison and Eastern Goldfields Provinces, it is virtually absent from the Kalgoorlie Subprovince where only minor amounts are present at Norseman (Williams, 1974). Black graphitic shales are common in all areas. This distribution may be interpreted in terms of a mafic lateral variation from an environment of black shale-chert-thick sequences (deep water?) t o one of thin, more variable mafic sequences with banded iron-formation and black shales (stable basins?). Alternatively, banded iron-formation-containing sequences in the Murchison Province and Southern Cross and Laverton Subprovinces may be older than sequences in the Kalgoorlie Subprovince. In such an instance, banded iron-formation units could be expected to occur beneath presently exposed sequences of the Kalgoorlie Subprovince. Banded iron-formation is represented at the surface by ferruginous caps, siliceous “cherts” or laminated magnetite and/or hematite rocks which form topographic and aeromagnetic “highs”. A classification of Archaean banded iron-formation based mainly on surficial exposures in the Murchison Province (Muhling and Low, 1973) is reproduced after Baxter (1975) in Table 2.111.
TABLE 2.111 Classification of Archaean banded iron-formation (after Baxter, 1975)
______
Type of banded iron forma- Associated rock types Lion (after Muhling and and Low, 1973)
Interpreted unoxidized rock type
hematite/magnetitequartz-mafic mineral
mafic volcanic rocks or cyclic succession of ultramafic t o felsic rocks
magnetite-quartz-mafic mineral iron-formation
hematite/magnetitequartz
fine-grained sedimentary and felsic volcanic rocks o r mafic volcanic rocks
hematite/magnetite-quartz iron-formation
goe thite-quartz
fine-grained sedimentary rocks
magnetite-carbonate ironformation
jasp i I i t e
fine-grained sedimentary or volcanic rocks
magnetite-quartz-mafic mineral iron-formation o r black shale
handed chert
fine-grained sedimentary rocks o r felsic volcanic rocks
possibly chert o r finegrained orthoquartzite
47
The first two varieties are prevalent in stratigraphically low mafic/ultramafic assemblages. Early documentation from gold mines by the Geological Survey o f Western Australia and more recent exploratory diamond drilling suggest that, at depth, unweathered banded iron-formation often consists of thinly laminated magnetite silicate rock quite distinct from surface appearance. Silicate bands in such banded iron-formation include felsic (quartz, quartz-albite) and/or mafic (amphibole, chlorite, epidote, pyroxene etc.) varieties, which may co-exist in individual samples. Clastic sedimentary textures have been observed (Baxter, 1975) as well as deformational textures which are probably related to penecontemporaneous rather than later tectonic deformation. Mesobanding, microbanding, pods and macules similar to those found in Proterozoic banded iron-formation have been recognized (Gole, 1976). Phosphate distribution in the magnetitesilicate banded ironformation is similar to that reported for Proterozoic banded iron-formation from the Hamersley Basin (Morris, 1973) as shown in Fig. 2.4 and the phosphate distribution is apparently unaffected by metamorphism as high as upper amphibolite facies (R.C. Morris, pers. commun., 1977). Although sulphide-rich units are common (mainly pyrite- and/or pyrrhotite-bearing), there is no suggestion of a systematic regional transition between sulphide, carbonate and oxide facies, such as has been observed in Canadian greenstone belts (Goodwin and Shklanka, 1967). Units of banded iron-formation in the Murchison Province and Southern Cross Subprovince are commonly intercalated with a variety of ultramafic schists, black shales, tuffs and paraamphibolites (mafic tuffs?) (Hallberg, 1976). Black shales are a common associate of lower mafic/ultramafic units. The shales are often carbonaceous and invariably contain disseminated to clotty sulphides (pyrite, pyrrhotite, sphalerite and often chalcopyrite). Tuffaceous textures may be found. It is possible that the shales have acted as scavengers, taking metals from their local aqueous environment. Further local enrichment may have occurred during diagenesis and subsequent hydrothermal activity. Sulphur, carbon and oxygen isotope studies of black shales from lower stratigraphic sequences indicate that sulphides, carbon and carbonate are of probable magmatic rather than biological origin (Donnelly et al., 1978). Sulphides from the shales have 634S values near O%O suggesting the incorporation of magmatic sulphur; the 6I3C and 6l8O values of introduced carbonates from stratigraphically low shales are also consistent with a magmatic origin. The reduced carbon (kerogen) in the shales appears t o be syngenetic, having formed wholly from hydrothermal exhalations containing reduced carbon or possibly from biogenic sources with loss of light carbon through subsequent thermal activity. Felsic volcanic rocks and associated volcaniclastic and sedimentary rocks are present in lower mafic/ultramafic associations. In the Murchison Province such units are represented by quartzsericite schists, felsic tuffs and siltstones which are laterally discontinuous. In the Kalgoorlie Subprovince
48
Fig. 2.4. Archaean magnetitesilicate banded iron-formation from the Murchison Province ( A ) showing phosphate distribution ( B )and streak print (C). Bar is 2 cm in length.
49 similar sequences are represented at Norseman, while thicker felsic volcanic complexes are found in the Kurnalpi area (Williams’ Gindalbie Group, Table 2.1). As felsic volcanic centres are more characteristically developed at lower levels of overlying sedimentary associations, they will be discussed in the next section. Bimodal basalt-rhyolite volcanic associations in the Kookynie-Melita area of the Kalgoorlie Subprovince may occupy a low stratigraphic position (Fig. 2.2B). Here, tholeiitic basalts (in part pillowed) are interbedded with felsic flows and thick sequences of rhyolitic crystal and vitric tuffs. The felsic flows tend to be amygdaloidal and felsic tuffs display a variety of devitrification textures (spheruloids, spherulites, micropegmatite) suggesting a high glass component. Lenses and pods of agglomerate, autobreccias and intrusion breccias are abundant. Sedimentary units are virtually absent. Both potassic and sodic felsic rocks are present. Although the sodic varieties are dominant, there appears t o be little systematic distribution of chemically distinct felsic rocks. Age dates from stratigraphically low basalts (Turek, 1966; J.C. Roddick, pers. commun., in Ross and Hopkins, 1975), sediments (Turek, 1966) and acid volcanic rocks and porphyries (Turek, 1966; Arriens, 1971) fall within the range of 2600-2700 Ma. Such dates, although representing minimum ages for the development of the supracrustals, are metamorphic ages and true eruptional ages are unknown.
Lower sedimentary association Stratigraphically overlying the lower mafic/ultramafic units in most areas of the Murchison and Eastern Goldfields Provinces is an essentially sedimentary association with felsic volcanic/sedimentary complexes common at basal stratigraphic levels. In the Murchison Province and Laverton Subprovince the lower sedimentary association is dominated by thick sequences of finely bedded shales and siltstones with locally abundant banded iron-formation and less commonly greywacke, mature, quartz and chert-rich sandstone and oligomictic conglomerate. Sedimentary structures developed primarily in shalesiltstone sequences include graded bedding and rare cross-bedding. In the Southern Cross Subprovince, the lower sedimentary association contains a greater abundance of quartz-rich greywackes and sandstones as well as shale-siltstone sequences. Although outcrop is poor, it is probable that the sedimentary succession is much thinner than in the Murchison Province. In the Kalgoorlie Subprovince, the lower sedimentary association is represented by thick beds of greywacke, arkose, siltstone, shale and polymictic and oligomictic conglomerates. Lateral facies changes in the lower sedimentary succession with an eastern deep-water facies and a western shallowwater facies have been documented near Kalgoorlie. Deep-water greywacke-
50 shale sequences containing abundant turbidite structures (graded bedding, washouts, scour channels, convolute lamination) are described from an area east of Kalgoorlie by Dunbar and McCall (1971). The western shallow-water facies is well exposed along the western margin of Lake Lefroy. Conglomerates become important in the western shallow water facies which also includes greywackes and shales with cross-bedding as the dominant sedimentary structure (McCall et al., 1970). However, it is unclear as to whether these changes occurred in a continuous basin or represent deposition in separate troughs. Furthermore, sediments of apparent deep-water origins do occur in the west (the Brown Lake metasediments near Coolgardie; Glikson, 1971b). .4rare earth element (REE) study of shales and greywackes from the lower sedimentary association is presented by Nance and Taylor (1977). Two groups of sediments were recognized on the basis of their REE patterns. One group displays heavy REE depletion and members were apparently derived from adjacent Na-rich granitoids or felsic volcanic rocks. A second group has patterns identical to recent calc-alkaline volcanic rocks and associated greywackes. Such patterns could be derived either from sediments shed from calc-alkaline volcanic rocks or by an appropriate mechanical mixture of debris derived from tholeiitic basalts and sodic felsic igneous rocks. The latter alternative is preferred since calc-alkaline volcanic activity is absent from the Kalgoorlie Subprovince, and occurs as discrete high-level centres in adjacent areas. In general, the sedimentary REE patterns are similar t o those of Archaean sediments elsewhere, showing greater variability, lower total REE and Eu enrichment as compared t o recent shales (Condie et al., 1970; Wildeman and Condie, 1973). Pods of agglomeratic and tuffaceous sodic felsic volcanic rocks occur sporadically within sedimentary sequences of the Murchison Province. There is little detailed information on these units and their relationships t o enclosing sediments are often unclear. Conglomerates similar to those associated with acid volcanic centres in the Kalgoorlie Subprovince (see below) are not known. Felsic volcanic rocks have not been reported from lower sedimentary units in the Southern Cross Subprovince. Small lenses to larger centres of predominantly fragmental sodic rhyolite are found at basal stratigraphic levels in the lower sedimentary association in the Kalgoorlie Subprovince (Fig. 2.2B). Agglomerates, autobreccias, and lapilli, crystal, crystal-lithic and lithic tuffs predominate, and rare flows are present. Poorly bedded oligomictic conglomerates containing locally derived, volcanic debris are a common associate. Although these conglomerates were probably shed from adjacent felsic volcanic centres, exotic clasts (granitoid, chert, mafic volcanic rocks, and at Kanowna, ultramafic rocks) may be present (Williams, 1970; Gee et al., 1976). The oligomictic conglomerates pass laterally into finer-grained sediments. Whether the felsic volcanic centres have also contributed material to these distal sediments is unclear in most cases, however, the previously cited REE work indicates that in some instances this contribution may be appreciable.
51 Upper mafichltramafic association An upper sequence of mafic/ultramafic extrusive rocks is locally preserved in regional synclines in the Murchison and Kalgoorlie Provinces. These sequences consist of well-preserved, often pillowed, tholeiitic basalts and interbedded or discrete units of high-Mg basalts. Podiform peridotites and peridotitic flows are uncommon, although transgressive units of dunite/ peridotite are locally abundant. Limited chemical data from these higher level sequences (Hallberg, 1976 and unpublished data; Hallberg et al., 1976b) suggest that the predominant tholeiite is relatively enriched in Ti and Zr
TABLE 2.IV Average analysis of representative mafic volcanic rocks from upper mafic/ultramafic sequences 1
2
3
4
5
II
8
6
1
3
16
SiOz A1203
47.4 (1.2) 14.3 (0.3) 2.5 (1.0) 10.7 (0.9) 5.9 (0.5) 10.1 (1.3) 3.1 (0.5) 0.04( 0.04) 3.5 (0.4) 0.3 (0.3) 1.32(0.1) 0.16(0.01) 0.25(0.02)
51.5 (2.5) 14.2 (0.7) 2.5 (0.5) 8.0 (1.0) 5.7 (0.5) 9.3 (1.6) 3.1 (1.3) 0.29(0.17) 2.4 (1.3) 1.0 (1.6) 1.18(0.3) 0.14(0.04) 0.22(0.02)
53.26 12.73 2.94 10.59 3.89 6.31 4.48 0.43 1.67 1.87 1.39 0.19 0.26
48.5 (0.3) 13.7 (0.6) 4.7 (0.5) 10.2 (0.2) 5.3 (0.9) 10.8 (0.5) 2.2 (0.03) 0.28(0.03) 1.3 (0.3) 0.1 (0.2) 2.03( 0.7) 0.29(0.1) 0.24( 0.02)
51.4 (0.9) 13.5 (1.3) 2.0 (0.5) 7.9 (2.2) 8.6 (0.3) 11.0 (1.3) 2.0 (0.5) 0.14( 0.1 2) 1.7 (0.3) 0.1 (0.1) 0.75( 0.2) 0.16(0.05) 0.1 9( 0.02)
32(20) 116 ( 9 ) 106 ( 7 ) <3 90 ( 8 ) <2 184(68) 343( 27) 37 ( 4 ) 108(11) 82 (4)
345(130) 70 (11) 151 (51) <3 94 ( 3 3 ) 9 (7) 112 ( 5 1 ) 343 ( 5 9 ) 23 ( 6 ) 87 (23) 8 5 (13)
310 10 61 <3 72 9 49 390 26 110 109
84 (82) 116 (42) 71 (26) <3 104 ( 5 0 ) 4 (2) 1 2 6 (23) 459 ( 7 0 ) 55 ( 1 6 ) 142 (58) 191(111)
59(20) 527(90) 53(21) <3 145(34) 5 (5) 133(65) 262(25) 14 (2) 59(14) 47(16)
O3
FeO MgO CaO NazO
KZ0 H20t COZ Ti02 PZOS MnO Ba Cr
cu Nb Ni Rh Sr V Y Zn Zr
I = pillowed mafic lavas (prehnite-pumpellyite facies), Laverton Subprovince; 2 = pillowed mafic lavas (prehnite-pumpellyite facies), Polelle (Hallberg e t al., 1976b); 3 = mafic pillow lava (hydrous greenschist facies), west of Abbotts, Murchison Province; 4 = mafic lavas (lower greenschist facies), Gullewa, Murchison Province; 5 = mafic pillow lavas (lower-middle greenschist facies), Mt. Singleton, Murchison Province.
52 and depleted in Cr compared t o the compositionally monotonous tholeiites of lower sequences in the Kalgoorlie Subprovince (Table 2.IV). However, at Mt. Singleton a sequence of pillow lavas and hyaloclastites, resting unconformably on a sequence of shalesiltstone-conglomerate, represents a pile of invariant Cr-rich, Ti-poor basalts (Hallberg, 1976). Ultramafic rocks in the Murchison Province and Southern Cross and Laverton Subprovinces include a small proportion of peridotitic flows and a preponderance of clinopyroxene-bearing, high-Mg basalts containing 1018% (weight) MgO. More evolved members possibly related t o the high-Mg basalts may also be present. The latter consist of platy clinopyroxene, plagioclase and quartz, and may contain as little as 4% MgO. Black shales and fine-grained felsic tuffs form intercalations in the mafic/ ultramafic sequences. Banded iron-formation is generally absent though thin units are found in basal sequences at Mt. Singleton, and clastic sediments are rare. In the Meekatharra area, there are indications that the mafic volcanism may be in part contemporaneous with the early phases of calc-alkaline volcanism (Hallberg et al., 1976b).
Upper felsic volcanicsedimentary association The highest stratigraphic units in some areas consist of unconformable sandstone, arkose and oligomictic or polymictic conglomerate. The polymictic Kurrawang conglomerate near Kalgoorlie contains rounded clasts of felsic: porphyry, basalt, granitoids (trondhjemite), and banded iron-formation (Glikson, 1971b). Cross-bedding and imbricate pebble structures are common. Granitoid pebbles from this unit probably fall within the time range 2670-2880 Ma, with a maximum possible age of 2885 f 160 Ma (Compston and Turek, 1973). Oligomictic conglomerates containing granitoid clasts have been reported from several areas. At Jones Creek, south of Mt. Keith (Durney, 1972), a sequence of arkoses and oligomictic conglomerates containing granitoid clasts unconformably overlies a granodiorite intrusive into mafic volcanic rocks. Durney further suggests that the conglomerate is overlain by a younger mafic volcanic sequence, but this view is disputed by Marston and Travis (1976) who suggest that extensive penetrative deformation and marginal shearing render relationships with the supposedly younger mafic sequence impossible to ascertain, and note that it is probable that the conglomerate is the highest unit present in the area. Marston (1976) further notes that the conglomerate is both chemically and texturally immature and suggests that it was formed by the rapid deposition of locally derived material in an alluvial fan. A maximum age for deposition of the Jones Creek conglomerate is provided by the date of 2689 1 7 Ma (Ri of 0.7015) obtained for the underlying granodiorite (Roddick et al., 1976). A similar deformed, unconformable conglomerate near Diemal’s north of Southern Cross (Binns and Marston, 1976), also contains granitoid-derived detritus. This conglomerate is intruded by granites. In general, the sedimen-
*
53 tary features of the Jones Creek conglomerate apply t o many of the highlevel coarse clastic sequences. Most are immature, represent locally derived debris and may have been deposited in high-energy, rapidly subsiding troughs (Glikson, 1971b). Felsic volcanic activity is evident in high-level sedimentary sequences in the Kalgoorlie Subprovince (for instance in the Binduli area, west of Kalgoorlie), but has not been reported from other areas. Laterally restricted high-level, largely subaerial calc-alkaline volcanic centres and associated sedimentary shed are found in the Murchison Province and Southern Cross and Laverton Subprovinces, but are unknown from the Kalgoorlie Subprovince. Such complexes are often preserved in regional synclines, lack penetrative deformation and are of low metamorphic grade (prehnite--pumpellyite t o lower greenschist facies). The calc-alkaline complexes are composed of flows of massive, porphyritic and amygdaloidal andesite with associated andesitic tuffs and agglomerates. Lack of sedimentary intercalations and pillow structures, and the extreme vesicularity of many of the flows suggest a subaerial origin. With the exception of the Marda complex north of Southern Cross (Hallberg et al., 1976a) which contains a sequence of subaerial rhyolitic ignimbrites, dacitic, and particularly rhyolitic volcanic rocks are rare. The Marda complex has major and trace element similarities with modern continental margin calc-alkaline suites. REE data on rocks from this complex suggest derivation by partial melting in lower crust of mafic composition (Taylor and Hallberg, 1977). The Marda complex has yielded an Rb-Sr wholerock age of 2 6 3 5 2 80 Ma (Ri of 0.703), which is believed to be primary (Hallberg et al., 1976a). The calc-alkaline complexes appear to occupy different tectonic settings from more modern analogues in continental margin and island arc environments. Calc-alkaline complexes in the Murchison and Eastern Goldfields Provinces have been deposited on sedimentary successions. The Marda complex was formed in a restricted basin as indicated by local derivation of underlying sediments and lateral facies changes. Andesitic rocks from the calc-alkaline complexes invariably show high Cr and low A1201 as compared to Phanerozoic andesites. The calc-alkaline complexes are often associated with coarse, immature sediments which represent rapidly derived, poorly sorted shed from the volcanic centres. These sediments range from coarse andesitic lithic wackes to greywackes t o more distal shale/siltstone sequences, and are composed essentially of andesite-dacite lithic fragments, clastic feldspar, carbonate and chlorite; quartz is nowhere prominent as a clastic mineral. Rare exotic fragments of chert, basalt and banded iron-formation are present. The dominant sedimentary structure is graded bedding, with few instances of crossbedding and intraformational breccias in shaly units. The very coarse immature lithic wackes attest t o the rapid stripping of the volcanic centres. They have many features in common with the high-level conglomerates just described, and were probably deposited in a similar high-energy environment.
54 Greywackeshale sequences further removed from the volcanic centres are conimonly interbedded with pillowed tholeiitic basalts unrelated to the calcalkaline complexes, black shales associated with distal felsic greywackes in the Laverton area contain carbonate of possible marine sedimentary origin (Donnelly et al., 1978). The shale-greywacke units attest to a far different environment of deposition than the coarser sediments and suggest rapid lateral changes in moving away from the volcanic centres. These changes are in accord with those suggested by Walker (1976) who indicated that “the gradation from relatively high-relief land areas to deep ocean basins was apparently abrupt” in Archaean terranes. High-level oligomictic and polymictic conglomerate sequences and calcalkaline volcanic centres and associated sediments are only sporadically distributed in the Murchison and Eastern Goldfields Provinces. It is not known whether this distribution reflects a trend towards more localized volcanic episodes and more restricted sedimentary provenance in younger stratigraphic levels, or whether it is an artifact produced by structure and exposure. Paucity of outcrop, thickening by intrusive rocks, lateral thickness changes and structural complications preclude the determination of accurate regional stratigraphic thicknesses. For these reasons, thickness estimates of 30,000 ni (McCall, 1969), 20,000 m (Glikson, 1970), 28,300 m (Doepel, 1973), 32,000 m (Gee, 1975), 27,000 m (Williams, 1975) and 9100 m (De la Hunty, 1973) could be unrealistic. On the other hand, because of tectonic attenuation, the measured thicknesses of well-exposed, structurally coherent units may represent minimum values.
Intrusive rocks Associated with mafic/ultramafic sequences and emplaced into the lower sedimentary sequences is a variety of mafic and ultramafic intrusive rocks. Mafic intrusive rocks include undifferentiated dolerite and gabbro dykes and sills similar in major and trace element composition t o the tholeiitic basalts (Hallberg, 1972a), differentiated gabbro dykes and sills with granophyric tops and noritic bases, porphyritic gabbros with accumulations of large calcic plagioclase phenocrysts (Gower, 1971), and layered mafic sills of probable tholeiitic parentage (Jaques, 1976). These intrusive rocks have been deformed with, and are believed to be comagmatic with associated or underlying mafic volcanic rocks, in contrast to the “older” and “younger” greenstone concept of Prider (1948, 1965). Discordant and concordant stringers and pods of intrusive dunite/peridotite are common both within mafic/ultramafic sequences and sedimentary sequences in many parts of the Murchison and Eastern Goldfields Provinces. Poor outcrop and a tendency toward semi-conformable relations leads to confusion between extrusive and intrusive varieties even at a local scale. The intrusive podiform dunite/peridotite bodies often show a marginal develop-
55 ment of talc-carbonate mantles, and may also display well-developed marginal quench-textured zones. These bodies generally consist of equant t o elongate, euhedrally terminated, olivine crystals (often serpentinized), with Interstitial bladed clinopyroxene crystals and/or chlorite-serpentine aggregates with minor chromite. Such bodies have been interpreted by Williams I 1972) as subvolcanic sills. Laterally extensive massive dunite t o peridotite dykes composed of closely packed, equant olivine crystals are found in the northern part of the Kalgoorlie Subprovince and in portions of the Southern ('ross Subprovince. They appear t o concentrate in zones of linear disruption bounded by major tectonic lineaments, and may represent ultramafic f t ~ d i-hannels ~ r (Hudson, 1972). Semi-conformable layered sills showing Stillwater-type igneous layering occur ubiquitously in both the Murchison and Eastern Goldfields Provinces. These sills show the differentiation sequence dunite - harzburgite - orthopyroxenite -norite - gabbro-(granophyre) with cumulus textures, phase and cryptic layering. Clinopyroxene tamulates are found in similar sills in the Murchison Province. These sills have bulk compositions with 15-1896 MgO and are genetically related t o high-Mg basalts; they probably represent differentiated subvolcanic intrusions (Williams and Hallberg, 1973). Large differentiated sill-form bodies of rhythmically layered gabbros and anorthosites with minor basal ultramafic cumulates and zones of cumulus vanadiferous titanomagnetite are common in the Murchison Province and northern portion of the Southern Cross Subprovince. Amongst these bodies is tht. Windimurra complex east of Mt. Magnet which occupies an area of some 11,000 km2. The Windimurra complex is intruded by granitoids and post-Archaean dykes and is presumably Archaean in age. These layered complexes may have affinities with anorthositic gabbros or high-A1 basalt (Binns and Marston, 1976). Discordant quartz, quartz-albite, and less commonly quartz-K-feldspar porphyries abound in portions of supracrustal sequences. In some areas they are associated with felsic volcanic centres; in others they occur as dense swarms, often near granitoid contacts. In the Marda area, porphyry dykes show discordant relationships with all stratigraphic members including the high-level calc-alkaline volcanic complex with which they have no chemical affinity. The presence of rhyolitic t o dacitic porphyry dyke swarms at some granitegreenstone contacts, and compositional similarities t o associated granodioriteadamellite and granite plutons (O'Beirne, 1968; Larking, 1969) suggest that some of the felsic dykes may represent the initial phases of granitoid generation. Others may be feeders t o felsic volcanic activity.
Structure and metamorphism St.ructura1 trends in the Murchison Province are NE t o NW with NNE t o NE trends dominating the northern portion of the area. The NNW trend
56 which characterizes the Eastern Goldfields Province reflects the orientation of both granitoid plutons and greenstone belts, regional and local fold patterns, major faults and lineaments and metamorphic foliation. Regional folds have steeply dipping limbs and vertical t o subvertical axial surfaces. The steep isoclinal folds and associated major strike faults of Kalgoorlie (Woodall, 1965) are typical. Mild E-W to ENE cross-folds may be superimposed on the dominant NNW trend (Prider, 1948; McMath, 1953; Sofoulis, 1966; Glikson, 1971a). Glikson (1971a) suggests that the crossfolding in the area between Coolgardie and Kalgoorlie post-dates the NNW isoclinal folding. Gemuts and Theron (1975) suggest that the cross-folds were derived from the synchronous intrusion of granitoids into anticlinal crests, O’Driscoll (1971) recognized NE and NW quadrantal lineaments on airborne magnetic maps of the Yilgarn Block which were believed t o be conjugate shear systems. Major NW-NNW lineaments and strike faults extending for distances of up to 500 km have been recognized (Williams, 1974; Gee and Williams, 1975). Gee (1976) notes that the disposition of sediments and granitoids about these lineaments suggests a long history of activity. Gower (1973) suggests that thick conglomerate sequences spatially associated with some of the major lineaments in the Laverton area may indicate vertical tectonic activity along the zones in late Archaean times. Evidence for polyphase deformation comes from the Kurnalpi-Edjudina area in the Eastern Goldfields Province (Williams, 1974; Williams et al., 1976). Williams recognizes an older north-trending fold system on which is imposed a younger NNW fold system. Many granitoid batholiths in the Kurnalpi area (Williams, 1974) are concordant with the NNW system, but truncate the older northerly trend. Regional refolding in the Mt. Clifford area northwest of Leonora is proposed by Thom and Barnes (1974) and further evidence from folded cleavage, superposed cleavages and superposed mesoscopic folds is given in Gee et al. (1976). Metamorphic grades in greenstone belts in both the Murchison and Eastern Goldfields Provinces range from prehnite-pumpellyite facies to upper amphibolite facies. The preponderance of sedimentary and volcanic assemblages documented lie in the lower t o middle greenschist facies. McCall (1969) placed the well-exposed volcanic and sedimentary rocks on the western shore of Lake Lefroy in the greenschist facies. Glikson (1971a) established that volcanic and sedimentary assemblages between Coolgardie and Norseman had undergone an Abukuma-type low-pressure greenschist facies metamorphism. Such assemblages predominate in the thick pillowed tholeiitic basalt piles of the Norseman area (Hallberg, 1972a). Within mafic sequences, the extended stability field of calcic plagioclase under such conditions generally results in greenschist facies assemblages of primary calcic plagioclase-actinolite with negligible amounts of carbonate, epidote, chlorite and albite. Anhydrous, low-pressure greenschist assemblages dominate mafic sequences in the Kalgoorlie Subprovince. Although similar assem-
57 blages are widespread in mafic volcanic rocks in the Murchison Province and Southern Cross and Laverton Subprovinces, these areas (the Laverton Subprovince in particular) also contain extensive prograde mafic greenschist assemblages of epidote-actinolite-albite-chlorite f sericite, carbonate and clinozoisite (Hallberg, 1976; Hallberg, unpublished data). These variations indicate that, on a regional basis, changes in style of greenschist metamorphism were present, and were perhaps related to differences in the availability of water during metamorphism. Glikson (1971a) presents evidence for at least two metamorphic/structural events in the Coolgardie-Kalgoorlie area. The earlier event is represented by a NNW flow-cleavage set and the younger event by an E- -Wstriking intersecting strain-slip cleavage set. Both events were accompanied by the growth of parallel nematoblastic phases. It is possible that the earlier phase is related t o the intrusion of concordant granodiorite batholiths and the later phase to the thermal rise accompanying late Mungari-type anatectic granites. Williams (1970, p. 23) notes that in the Kurnalpi area, “metamorphic grade is slightly less for each successive volcanic cycle”, with the uppermost units being nearly fresh. This situation is repeated in many areas in both the Murchison and Eastern Goldfields Provinces, in that the stratigraphically youngest volcanic rocks represented by calc-alkaline volcanic centres are of generally lower metamorphic grade than underlying volcanic-sedimentary assemblages. Binns et al. (1976) recognize two regional metamorphic domains in the Eastern Goldfields Province. In static style domains, penetrative deformation is minimal, primary rock textures are well-preserved, and grades range from prehnite-pumpellyite to middle greenschist facies. Dynamic style domains are characterized by extensive penetrative deformation, textural obliteration, and middle to upper amphibolite facies metamorphism. Dynamic style domains are restricted t o the margins of supracrustal sequences and screens between granite intrusions. Binns et al. (1976) suggest that the entire regional metamorphic pattern of the Eastern Goldfield Province developed during a single episode coincident with basement reactivation and generation of granitoid plutons and a marked change in the thermal regime in the supracrustal sequences. Both static and dynamic style domains are believed to have been contemporaneous and were imposed on previously deformed supracrustal sequences. The support of Binns et al. (1976) for the “single metamorphism” hypothesis derives from lack of evidence for repeated metamorphism and from relic mineral distribution which suggests that metamorphism was imprinted on previously unmetamorphosed rock sequences. The division into dynamic and static style domains by Binns et al. (1976) is not consistent with some published data. Glikson (1971a) cites strain “rates” of up t o 50% in “static” style domains as shown by pebble analysis in the area between Kalgoorlie and Coolgardie, although Gee et al. (1976)
58
suggest that the bulk flattening of these pebbles may be considerably less than 50%. The use by Binns et al. (1976) of calcic plagioclase as a prograde metamorphic indicator without data on the alumina content of associated amphiboles is questionable. The transition in mafic rocks from greenschist facies to amphibolite facies is represented by a change from actinolitelabradorite t o andesine-aluminous hornblende assemblages (Glikson, 1972b; Hallberg, 1976). Using calcic plagioclase alone as a prograde indicator would result in the niisassignment of metastable greenschist assemblages t o a higher metamorphic grade. The “single metamorphism” hypothesis of Binns e t al. ( 1976) does not account for previously mentioned evidence for polyphase metamorphism and, for advocates of a pre-greenstone sialic crust, necessitates the development of dynamic and static style metamorphic zones prior to acid plutonic activity. Areas of very low-grade, static-style metamorphism as shown by Binns et al. ( 1976) are broadly coincident with younger stratigraphic sequences exposed in regional synclines, further indicating that metamorphic grade is related t o relative stratigraphic level and to spatial relationships with granitoids. It is probable that granitoids were both the cause for the folding and a heat source for metamorphism, which explains why synclinal cores distant from the granitoids are both least deformed and least metamorphosed. G ra n ito ids
Extensive surficial cover over granitic terrane has hindered studies of granitoids on a regional basis. However, regional studies in recent years have greatly advanced our understanding of these rocks in the Yilgarn Block. Sofoulis (1963) recognized “external” and “internal” granitoids in an area west of Coolgardie, the former including the large masses of granite separating greenstone belts, and the latter distinct granitoid plutons emplaced within the greenstone belts themselves. Age dating by Turek (1966) showed that the “internal” and “external” granitoids were broadly contemporaneous. Mapping in the eastern portion of the Eastern Goldfields Province (Williams, 1970; Williams et al., 1976) led t o the recognition of early granodioritic t o tonalitic, heterogeneous, strong foliated gneisses and migmatites, later composite, concordant t o discordant intrusive batholiths of granodiorite to adamellite and granite, and still younger massive discordant plutons of granite t o syenite. The oldest gneisses and migmatites are associated with extensive thermal aureoles in surrounding volcanic and sedimentary rocks, and the younger varieties with less extensive thermal aureoles. The granodiorite-adamellite plutons were often found t o occupy major anticlinal structures. Williams (1975) viewed the felsic plutonic activity as a continual process, coincident with regional metamorphism and tectonism. Williams also stressed the point that the granitoids formed a series, becoming less conformable and more differentiated with progressively younger ages.
59 On the basis of the limited chemical information available a t the time, Glikson and Sheraton (1972) suggested a three-fold classification of acid igneous rocks in the Kalgoorlie region as follows: (1) pre-geosynclinal Na-rich plutons (represented by trondhjemitic pebbles in the Kurrawang Conglomerate); ( 2 ) syn-geosynclinal rhyolitic t o dacitic porphyries and small stocks of Na-rich granite (O’Beirne, 1968); and ( 3 ) late-kinematic granites, exemplified by the Mungari adamellite. Glikson and Sheraton emphasized a secular decrease in Na/K and K/Rb with time, in agreement with the previously c*ited field evidence of Williams. Glikson and Sheraton also note that the older sodic granitoids have compositional features (high Ni and Cr) suggestive of an origin by partial melting of, or contamination by, pre-existing niafic rocks. They suggest a high-temperature emplacement for the early sodic granites, which may account for the extensive thermal aureoles associated witah such bodies. Glikson and Sheraton also suggested that the younger granitoids may have been derived from the older granitoids by anatexis. The older granodiorites t o tonalites appear t o intrude only stratigraphically low supracrustal sequences whereas younger adamellites t o syenites intrude all stratigraphic levels (Glikson and Lambert, 1976). In a recent classification of granitoids based on texture and structural relationships with greenstone belts, Archibald and Bettenay (1977) recognized four types of granitoid in the Eastern Goldfields Province: synkinematic plutons of tonalite t o syenite, homogeneous postkinematic plutons of porphyritic adamellite, nebulitic and banded migmatites, and ortho- and para-gneisses. The younger postkinematic batholiths have followed the major period of deformation, and are locally discordant to structural trends in host rocks; thermal aureoles are restricted and boundaries linear and faulted. The older, elliptical, often composite synkinematic plutons are found a t regular intervals within greenstone belt margins. They are invariably restricted t o dynamic-style metamorphic domains and lineations within them parallel foliations within host rocks. Archibald and Bettenay (1977) suggested a solid-state emplacement for these plutons. The latter suggestion contradicted the previously cited observation of contact metamorphic aureoles around older sodic granitoids (Williams, 1970, 1976). More likely the dynamic-style metamorphic domains associated with such granitoids reflect contact metamorphic aureoles, shearing and folding associated with their emplacement, as well as solid-state deformation of progressively chilled mantles of the plutons, rather than pre-existing metamorphic zones. Archibald and Bettenay noted that banded gneisses and migmatites occur as xenoliths and enclaves in both synkinematic and postkinematic granitoid plutons and may predate the supracrustal sequences. Systematic relations between gneiss and massive granitoid plutons, as observed in the Pilbara Block (this chapter), may be obscured by poor exposure. The classification of Archibald and Bettenay (1977) includes neither chemical nor isotopic parameters. With poor exposure, and unknown compositional affinities,
TABLE 2.V Isotopic ages of granites in the Eastern Goldfields Province, Murchison Province and Pilbara craton, Western Australia (only the oldest ages of each suite are shown; younger reset ages of these rocks are not listed) Unit
Locality
Isotopic age (Ma)
Method
Ri
(1)Pilbara craton foliated migmatitic granite Mt. Edgar batholith
Marble Bar area, Pilbara
3125 2 366
Rh-Sr whole rock
0.7016
migmatite, granite and mafic bands
Tambourah. Pilbara
3 0 7 0 2 1 2 or Pb-Pb mineral total rock isochron 2 9 3 8 2 33
granitic rocks
east and west Pilbara
3050 2 180
Rb-Sr whole rock
foliated granite and pegmatite vein
204 km south of Port Hedland, Pilbara
2961 2
Pb-Pb mineral total rock isochron
foliated migmatitic granite Shaw batholith
Spear Hill area, Pilbara
2 9 5 1 2 83
Rb-Sr whole rock
plagioclase-quartzmicrocline-biotite granite
112 km south of Port Hedland, Pilbara
2 8 8 2 2 60
Pb-Pb mineral total rock isochron
0.93(3)
Oversby, 1976
gneissic granite and pegmatite
Woodstock, Pilbara
2769f
Pb-Pb mineral total rock isochron
0.99(4)
Oversby, 1976
7
13
Average Average KzOl RblSr NazO (number of samples shown in brackets)
0.31(10)
Reference
De Laeter and Blockley, 1972 Overshy, 1976
0.7 5(4 )
Compston and Arriens, 1968 Oversby, 1976
0.86(4)
O.lO(10)
0.702
De Laeter e t al., 1975
Moolyella adamellite Marble Bar area, Pi1bara
2 6 7 0 2 95
Rh-Sr whole rock
0.7397
9.2(6)
De Laeter and Blockley, 1972
Cooglegong adamellite
2602 2 132
Rb-Sr whole rock
0.7310
8.58(8)
De Laeter e t al., 1975
Shaw batholith, Pilbara
( 2 ) Eastern G o l d f i e l d s Province
Kambalda granodiorite
Kambalda, Eastern Goldfields
2760 -C
70
Pb-Ph mineral whole rock isochron
0.42(4)
Karonie granodiorite
Karonie, 112 km east of Kalgoorlie
2699 5 75
Pb-Pb mineral whole rock isochron
0.67(4)
Mt. Keith granodiorite (and derived pebbles)
west of Jones Creek
2689 i 1 7
Rb-Sr whole rock
0.70149
0.72(10)
0.23(22)
Roddick e t al.. 1975
trondhjemite pebbles Kurrawang Conglomerate
west of Kalgoorlie
2670 k 255 (minium age)
Rb-Sr whole rock
0.7025
0.066(7)
0.052(11)
Compston and Turek, 1 9 7 3
Stennet Rocks veined gneiss
4 8 km SSW of Norseman
2671 k 79
Pb-Pb mineral whole rock isochron
0.60( 2 )
Oversby, 1 9 7 5
Buldania Rocks
east of Norseman
2655 k 3 5
Pb-Pb mineral whole rock isochron
0.87(3)
Oversby, 1 9 7 5
Wheeler Rock: biotite adamellite and porphyritic granite
Lake Johnson, west of Norseman
2632 f. 28
Pb-Pb mineral whole rock isochron
0.9-1.0 (west) 1.4-2.7 (centre)
Oversby, 1 9 7 5
Mungari-Karramindie Soak, adamellite
Kalgoorliee-coolgardie area
2640
Pb-Pb mineral whole rock isochron
1.28(4)
4.6
Oversby, 1 9 7 5
Lawlers tonalite
Agnew area
2632 2 35
Rb-Sr whole rock
0.52
0.104
Cooper e t al., 1976b
Mungari Granite
east of Coolgardie
2620 k 20
Rb-Sr (assumed Ri value)
1.2
4.6
Compston and Turek, 1 9 7 3
Lawlers Leucotonalite
Agnew area
2549 ? 18
Rb-Sr whole rock
0.70223
0.57
0.5
Cooper e t al., 197613
2605 f 51
Rb-Sr whole rock
0.7031
0.44
Muhling and De Laeter, 1 9 7 1
2590 k 23
Rb-Sr whole rock
0.7028
3.99
Muhling and De Laeter, 1 9 7 1
( 3 )Murchison Province Cue “outer” granitic Cue area rocks (granodiorites)
Cue “inner” granitic Cue area rocks (adamellites)
0.70145
0.11
Oversby, 1 9 7 5 Overshy, 1 9 7 5
6,
F
62 granitoids identified as postkinematic could represent the little-deformed cores of gneiss-mantled plutons. Furthermore, a classification based on contact relationships with greenstones is difficult to extend for the identification of individual plutons in the continuous granitoid terrane between greenstone be 1ts . Granitoid ages are summarized in Table 2.V, the average major and trace elernent content of granitoid types are presented in Table 2.VI and various chemical parameters illustrated in Figs. 2.5 and 2.6. Pooled regional Rb-Sr isochrons, total-rock and mineral Rb-Sr isochrons and Pb-Pb systematics yielded granit,oid ages generally within the 2600-2700 Ma range with some dates to 2550 Ma. Similar as well as older ages are reflected in the Southwestern Province. Initial 87Sr/86Srratios are generally low, suggesting that rc.cyc.1e.d sialic crust was not involved in granitoid generation. Binns and ilarston (1976) and Glikson and Lambert (1976) noted, however, that a rltajurity of the granitoid ages have been derived from the postkinematic variety, leaving wide scope for further revision. Multistage Pb-Pb models for granitoids near Norseman (Oversby, 1975) were interpreted in terms o f tvolutionary history involving parental adamellite at least 3300 Ma old. High initial 87Sr/R6Sr ratios in a pegmatite from the Lawlers area northwest o f 1,eonora (Cooper et al., 1976b) also suggest the involvement of more a nc*icntsialic material. fikon om ic geology One of the more significant aspects of the economic geology of the Yilgarn Block is the concentration of mineralization in lower stratigraphic levels. The ensuing discussion will proceed in terms of this stratigraphic asso(,;ation. The West Australian “nickel province” is largely confined to ultramafic rock sequences of the Kalgoorlie Subprovince, although some deposits occur in the Laverton and Southern Cross Subprovinces. Palladium, platinum, ruthenium, gold, copper, silver and cobalt are recovered from the nickel wlphide ores as by-products. Massive, matrix and disseminated nickel sultdirdes occur near footwall contacts in volcanic flows and subvolcanic sills of peridotitic composition. The sulphides are believed to be of magmatic origin (cf. Ewers and Hudson, 1972), with sulphur isotopes in the magmatic range (Donnelly et al., 1978). Disseminated nickel sulphides and, less commonly, massive and matrix sulphides, also occur in intrusive bodies of dunite in the northern portion of the Kalgoorlie Subprovince and southern part of the Southern Cross Subprovince (Mt. Keith, Forrestania). These bodies may represent more deep seated subvolcanic equivalents of the concordant hypabyssal and volcanic varieties (Hudson, 1972). Economic concentrations of nickel and copper sulphides have also been found in breccia pipes in a layered complex (Purvis et al., 1972). Gemuts and Theron (1975) note that
63 TABLE 2.VI Average and representative granite analysis f r o m t h e Kalgoorlie and Pilhara regions 1 -__
~~~
s10,
TiO, 3
Fez07
FeO MnO MRO CaO Ne20 K*O HZOt P205
Total
Ba Sr Rh Pb Zr Th Y
c 11
74.60 0.15 13.96 1.02 * ~
0.03 0.31 1.20 7.40 0.49 1.19 0.03 100.58 37 1 298 12 17 100 13 17 7 -
LI NI C1. V
9 6 18
Zn
-
__-
2
3 71.28 0.33 15.94 0.97 9.84
4 72.53 0.27 15.70 0.26 1.04
74.38 0.15 14.39 0.04 0.90
-
-
-
0.70 1.09 5.68 2.28 0.74
1.54 0.48 5.43 2.19 0.58
0.02 0.56 4.05 4.93 0.23
-
-
100.54
100.02
-
882 67 36
-
787 89 25
-
99.69 -
91 465 95
5
6
69.8 0.4 14.8 1.2 1.8 0.05 0.8 2.7 4.4 2.6 1.0 0.10 99.65 390 135 -
73.4 0.35 13.4 0.8 1.78 0.07 0.4 1.63 3.89 3.6 0.63 0.06 100.02
7
600 120 240 25 120
74.9 0.1 13.4 0.4 0.8 0.05 0.1 0.8 4.1 4.6 0.7 0.03 99.98 -
40 520 -
-
-
-
-
-
-
-
-
-
-
-
-
-
-
-
-
38 20 27 27 46
33 161 14 16
-
10
-
120
-
-
10 70
190 -
~
32 14 19 19 -
-
-
-
-
-
-
-
30
-
1 = average of 7 trondhjemite pebbles from t h e Kurrawang conglomerate, Kalgoorlie area (Glikson and Sheraton, 1972); 2 = average albite porphyry from t h e Kalgoorlie-Norseman area (major elements - 50 analyses, trace elements - 17 analyses) from O’Beirne (1968); 3 = sodic granite with zoned plagioclase, Kambalda (O’Beirne, 1968); 4 = hlungari granite, Coolgardie (O’Beirne, 1968); 5 = average o f 9 biotite granodiorites from migmatites of t h e Shaw and Mt. Edgar hatholiths, eastern Pilhara (Blockley, 1 9 7 0 ) ; 6 = Mondana granite, Corunna Downs batholith, eastern Pilhara (Hickman and Lipple, 1975); 7 = average of 11 analyses of t h e Moolyella adamellite, Mt. Edgar batholith, eastern Pilbara (Hickman and Lipple, 1975). * Total iron as F e 2 0 3 .
most of the nickel mineralization in the Kalgoorlie-Norseman area occurs in Sequence 3 of their stratigraphy (Table 2.1). Auriferous quartz lodes and reefs are found within all major rock types present in the Yilgarn Block. Arsenic, antimony and silver have been
64 __
740
~
* . 0
.
A
t
+
710
.700
i
i_
4
6
+;+*
%
*,. 3.5
.+*
-. 3
2.5
2
1
I .5
I
0.5
B .Y.
Fig. 2.5. Initial 87Sr/86Srratio versus Rb-Sr isochron age plot for acid igneous rocks and acid orthogneisses from the Pilbara and Yilgarn Blocks. Dots - granitic rocks; triangles acid volcanic rocks; diamonds - orthogneisses; crosses - granites from the Canadian Shield. Sources of data are listed in Glikson (1980). Principal sources listed in this chapter are: Arriens (1971); Compston and Arriens (1968); Bennett et al. (1975); De Laeter and Blockley (1972); De Laeter and Trendall (1974); De Laeter (1977).
recovered from some gold mining districts. Woodall (1975) stresses the importance of mafic host rocks (dolerite, basalt), but in some areas major deposits occur in clastic and chemical sediments and para-amphibolites (Hallberg,
65
‘
’ I
00
-
+
K20/Na20
0 50
+
+
+
+
0
*H 0
0
-
0
B.Y.
Fig. 2.6. K20/Na20 versus isotopic age plot for granites from the Eastern Goldfields Region and the Pilbara Block, based o n data from Table 2.VI. Solid circles - Pb-Pb ages for Eastern Goldfields granites; open circles - Rb-Sr ages for Eastern Goldfields granites; KP - pebbles from the Kurrawang Conglomerate; crosses - Pb-Pb ages for Pilbara granites.
1976). Despite the major role that structure has played in localizing gold mineralization, it is apparent that a gross stratigraphic control has also been exerted. Gold deposits in the Murchison Province occur almost exclusively in lower mafic/ultramafic and lower sedimentary associations (Hallberg, 1976), and are absent from stratigraphically higher sequences. A similar situation exists in the southern part of the Kalgoorlie Subprovince, where the majority of the gold production comes from the lower mafic/ultramafic associations (Gemuts and Theron, 197 5). High-level calc-alkaline complexes and elastic sedimentary sequences have produced little or no gold. Banded iron-formation units, which are generally restricted t o lower stratigraphic positions have yielded supergene-enriched iron ore, lateritic ochre and manganese. However, iron ore production is of minor importance compared to that from the Proterozoic formations of the Hamersley Basin. Application of the so-called “Noranda model” (Sangster, 1972) has failed to date to disclose significant deposits of c o p p e r u i n c (ZeadsiZuer) sulphides in the Yilgarn Block. The high-level calc-alkaline complexes are devoid of sulphide mineralization. Recently discovered c o p p e m i n c sulphides at Golden Grove, southeast of Yalgoo (Murchison Province), appear to be in part stratabound, and occur in a stratigraphically low sequence of finegrained sediments and felsic tuffs (K.M. Frater, pers. commun., 1977). Copper and zinc sulphides of uneconomic quantities are also found at phase
66 interfaces in some layered sills and differentiated gabbros and in fractures, joints, and shears in host rocks of the layered and differentiated basic bodies. Weathering and lateritization of black shale sequences in places has led to uneconomic near-surface concentrations of copper, zinc and rarely silver. At depth, such shale sequences may contain pods and lenses of massive pyrite which are often laterally discontinuous and show erratic distribution of copper and zinc. Rich supergene and oxidized copper ores have been worked at Ravensthorpe, Eulaminna and Gabanintha in the early 19OO’s, but further exploration has failed to delimit more extensive primary sulphides at depth in these areas. Archaean granitoids and pegmatites yielding ages in the 2550-2700 Ma range have been worked for beryl, bismuth, lithium, molybdenum, niobium, tungsten and tin. Surficial weathering and lateritization of Tertiary age over portions of the Yilgarn Block have yielded deposits of magnesite, bauxite, tin and nickel laterite.
Evolution of the Yilgarn Block Because of difficulties inherent in the unravelling of the granite-greenstone terrane of the Yilgarn Block, no unified evolutionary model exists which is consistent with all available data and acceptable to all workers. Rather than attempt such a presentation, we propose to discuss the nature of the difficulties, present a review of the development of evolutionary concepts, and indicate some genetic constraints that may be applicable. Because outcrops in most areas of the Yilgarn Block are limited to salt lake margins, creek beds and low rubbly hills, relatively less resistant lithologies, faults and unconformities are seldom exposed. Direct observation of their nature often depends on the availability of mine and drill-core information. Inferences based on structural data, geochemistry and isotopic geology have therefore been widely used in models of the evolution of the granite-greenstone system. However, as no consensus exists regarding the unique applicability of any single method or criterion and because models based on different criteria cannot always be reconciled, conflicts arise. Viable models should attempt t o explain a maximum amount of available data in terms of a minimum number of variables. Comparisons between Archaean and modern tectonic environments, while a necessary and useful approach, must not assume any strictly uniformitarian principle or correlation, and the significance of differences between Archaean and recent systems must be fully considered. Extrapolations from well-exposed Archaean systems, such as the Pilbara or the Barberton Mountain Land (Transvaal), though instructive, cannot be used as the sole basis for genetic models because of the possibility that significant differences existed between the mode of evolution of granite-greenstone systems of different ages or in differen t places.
67 Genetic concepts have been largely based on information from the Kalgoorlie Subprovince, which has provided a majority of the gold and nickel produced from the Yilgarn Block and is the most extensively studied area therein. Early geological reports, following the discovery of gold at Coolgardie in 1892, interpreted the terrane as the deeply eroded roots of an orogenic zone (Chewings, 1896). Greenstones were regarded as “dioritic dykes intruded into dioritic country rocks” (Goczel, cited in Honman, 1916), a view mirrored by the later classification of these rocks into “older greenstones” (volcanic rocks) and “younger greenstones” (intrusive sills) by Simpson and Gibson (1912). Early geologists were impressed by the massive conglomerate beds between Coolgardie and Kalgoorlie (Kurrawang Conglomerate), whose sedimentary origin was identified (Chewings, 1896; Gibson, 1911); however, until much later, the arenaceous and argillaceous metasediments of the Eastern Goldfields Province were described by many workers as sheared porphyries and porphyroid rocks - an understandable mistake arising from the derivation of many of these rocks from acid volcanic centres. McKinstry (1945) commented on the resemblance between the terrane of the Eastern Goldfields and Canadian greenstone belts. Prider (1952) suggested that the deposition of volcanic and sedimentary sequences in the Jimperding (Wheat Belt), Yilgarn (Southern Cross) and Kalgoorlie geosynclines - which he did not consider as necessarily contemporaneous - was succeeded by isoclinal folding, granitization and granite intrusion representing an orogenic cycle. Intrusive mafic bodies were regarded as younger than extrusive and sedimentary rocks of the supracrustal sequences. Subsequently Prider (1965) formalized the supracrustal stratigraphy as the Kalgoorlie-Yilgarn Succession, which he regarded as the oldest unit of the Yilgarn Block. Wilson (1958) and Prider (1965) proposed a broad subdivision of the granitic rocks of the Yilgarn Block into “older granites” and “younger granites”: the former were thought to represent granitized synkinematic equivalents of the Kalgoorlie-Yilgarn Succession, and the latter postkinematic, locally discordant, magmatic intrusions. Wilson (1958) regarded the older granites as representing in part, or comprising remnants of, a pre-greenstone sialic basement, t o which the high-grade ultramafic relicts of the Southwestern Province were thought to be welded. Wilson (in Horwitz and Sofoulis, 1965) suggested that the Kalgoorlie-Norseman supracrustals evolved within a rifted intrasialic structure, and that remnants of the older granitic basement may be found west of Coolgardie - a concept revived at a later stage by Binns and Marston (1976) and Archibald and Bettenay (1977). During 1963-1970 detailed studies in parts of the Kalgoorlie-Norseman area by staff and students of the University of Western Australia gave rise t o new concepts on the evolution of this terrane. The stratigraphy of this area, originally divided into two major sequences separated by an unconformity (Honvitz and Sofoulis, 1965) has been elucidated in areas west of Lake Lefroy (McCall, 1969; McCall et al., 1970), northeastern Lake Lefroy (Dun-
68 bar and McCall, 1971) and the Coolgardie-Kurrawang area (Glikson, 1970, 1971a). An analogy with Alpine-type geosynclinal belts was generally favoured, as based on: (1) a transition with time from a eugeosynclinal ophiolite-like assemblage into a turbidite flysch-like assemblage and finally a conglomeratic molasse-like assemblage (Glikson, 1970); (2) the occurrence of spilite-keratophyre-like associations; and ( 3 ) a possible division into a western eugeosynclinal zone and an eastern miogeosynclinal zone (Dunbar and McCall, 1971). Comprehensive geochemical studies of the metavolcanic rocks, notably by O’Beirne (1968), Hallberg (1972a) and Hallberg and Williams (1972), documented the Na-rich composition of the dacitic t o rhyolitic volcanics, the LIL-element poor composition of the tholeiitic basalts, and the unique high-Mg to peridotitic composition of the ultramafic volcanic to hypabyssal rocks. These results furnished strong support t o models advocating ensimatic origin of the greenstones, i.e. oceanic crust and/or island arc type regimes (Glikson, 1970; White et al., 1971; Hallberg, 1972a). At the same time, they emphasized the temporally unique affinities of the volcanics, particularly the quench-textured high-Mg basaltic to peridotitic volcanic rocks. Another result of these studies was the chemical correlation of extrusive and hypabyssal volcanic rocks, now regarded by most workers as coeval analogues (McCall and Leishman, 1971; Glikson, 1972b; Hallberg and Williams, 1972; McCall, 1973). O’Beirne (1968), Hallberg (1972) and Hallberg and Williams (1972) emphasized compositional distinctions between acid porphyries, tholeiitic basalts and high-Mg volcanic rocks which they regarded as genetically unrelated. These detailed geochemical and petrological studies suffered from the lack of a regional stratigraphic framework and geochronological data to which they could be related, but during the last ten years this gap has been progressively bridged by regional mapping and stratigraphic synthesis. Despite the limitations imposed by poor outcrop and complex structure, compatible stratigraphic concepts have been developed. Inherent in these concepts is multiple volcanicsedimentary cyclicity. Gemuts and Theron (1975) noted the occurrence of felsic volcanic and sedimentary rocks at low stratigraphic levels in the Norseman area, and favoured deposition of the supracrustal successions above granitic basement although no evidence for the latter was presented. In Williams’ (1975) stratigraphic model for a portion of the Eastern Goldfields Province, the preservation of thicker sequences, larger numbers of greenstone cycles, and great abundances of ultramafic rocks in the Kalgoorlie Subprovince were taken as evidence for a downfaulted position and greater mobility in relation to the Southern Cross and Laverton Subprovinces. Williams (1975) and Gee (1976) supported an ensimatic origin for the greenstones, the first author interpreting the differences between the various sectors of the Yilgarn Block in terms of different levels of outcrop. Geophysical (seismic) evidence (Mathur, 1976) has been interpreted in terms of an eastward tilt to the Yilgarn Block (Glikson and Lambert, 1976) in sup-
69 port of such a concept. However, the occurrence of a high-level calc-alkaline volcanic complex at Marda in the Southern Cross Subprovince (Hallberg et al., 1976a) and of high-level felsic volcanicsedimentary associations in the Laverton Subprovince (this chapter) cast some doubt on this model, unless downfaulting and/or eastward crustal tilting and erosion occurred prior t o the formation of these late-stage volcanic rocks. In contrast to the ensimatic models reviewed above, Oversby (1976), on the basis of multistage lead model calculations, proposed that some of the granodiorites in the Norseman area were derived from adamellites at least 3300 Ma old - deducing a differentiated sialic basement in the area. Further, an ensialic trough model (Wilson, 1958) has been supported by Binns and coworkers who, on the basis of aspects of structural and metamorphic geology of the Kalgoorlie-Norseman-Southern Cross area, attempted t o demonstrate the pre-existence of a continuous granitic basement beneath the greenstones (Binns et al., 1976; Binns and Marston, 1976; Archibald and Bettenay, 1977). In this model, anatexis and remobilization of basement underneath the supracrustal cover resulted in folding and metamorphism of the latter according t o two contrasting schemes: (1)dynamic-style metamorphism along trough-margin loci, associated with emplacement of solid-state synkinematic granites: and (2) static-style metamorphism toward central portions of the trough. These developments are regarded as parts of a single tectono-thermal episode, culminating in extensive invasion by postkinematic granites. The validity of stratigraphic schemes based on regional mapping and the significance of the isotopic data are questioned by this school. Occurrences of banded gneiss at the margin of the Pioneer granite dome (Fig. 2.2) have been regarded as derivatives from an underlying granitic basement (Archibald and Bettenay, 1977). Felsic volcanic rocks incorporated in the greenstone belts and units of banded iron-formation have been cited as evidence for an underlying sialic basement and for shallow-water, tectonically stable conditions (Binns et al., 1976). The ensialic trough model has been criticized by Glikson (1979) on the basis of: (1)A preferred view of gneissic margins around diapiric granitoids as dynamic structures associated with plutonic emplacement, rather than as xenoliths from an older basement. (2) The low-K and low-& composition of the synkinematic granodiorites and of the dacitic to rhyolitic volcanics, which renders unlikely their derivation by anatexis of granitic rocks and favours their origin by partial melting of greenstones. ( 3 ) The tectonically unstable environment of deposition, i.e., evidence for rapid subsidence of the greenstone troughs. (4)A criticism of assumptions underlying Oversby’s (1976) deductions. (5) The significance of lithological and geochemical distinctions between ensialic Proterozoic systems and Archaean granite-greenstone terrane.
70 The question of the nature of the basement of the Kalgoorlie supracrustals is intrinsically related t o the relations between the latter and high-grade gneiss-ranulite suites such as those of the Southwestern Province. As mentioned above, Prider (1952) regarded the greenstones as the older components throughout the Yilgarn Block, whereas Wilson (1958) considered at least. some of the gneisses as antedating the greenstones. McCall (1972) and Glikson and Lambert (1976) preferred an interpretation of the gneiss-granulite suite as the equivalent coeval root zones of the granite-greenstone terrane, the latter authors suggesting that c. 3.0 Ga-old gneisses such as occur in the Southwestern Province (Arriens, 1971) may be represented by littleexposed, yet undated, tonalites in the Eastern Goldfields. Gee (1976) pointed out the importance of clastic metasediments in the high-grade Chittering and Jumperding belts in the Southwestern Province, and Binns and Marston (1976) remarked on a paucity of high-grade greenstone equivalents in the Southwestern Province - both constituting arguments for a temporally distinct, rather than coeval, nature of the high-grade suite in relation to the granite-reenstone system. Gee (1976) suggested that the Southwestern Province represents the site of an early sialic nucleus - the rest of the Yilgarn having formed subsequently through the cratonization of an originally simatic crust. The age of the metasedimentary components of the high-grade suite, however, is not clear. One possibility is that they predate the c. 3.0-Ga-old gneisses dated by Arriens (1971). Another is that they represent equivalents of the high-level, late stage, felsic volcanic-sedimentary association of the greenstone belts. In the latter case the uplift and erosion of the granite-greenstone system in the Southwestern Province would have preceded both the termination of greenstone belt evolution and the regional metamorphism. Further insight into these questions must await detailed mapping in the southwestern and central sectors of the Yilgarn Block and determination of age relations between the high-grade and the low-grade terranes. It appears that, to date, no single model has been presented which can satisfactorily account for all observed features, nor, in the absence of detailed studies of the granitic terranes, can such a model be advanced here. Nevertheless, currently available data provide numerous constraints on models for the evolution of the granite-greenstone terranes. Some suggestions bearing on the genesis and development of the Archaean crust of the Yilgarn Block are listed below: (1)Some stratigraphic successions in the Eastern Goldfields and Murchison provinces comprise at least two ultramafic-mafic t o felsic volcanicsedimentary cycles, with the possibility of a third cycle in some areas. Whereas the lower cycle predated the older granodiorites, it is possible, though unproven, that the upper cycle was either contemporaneous with or postdated these plutons. (2) In the Kalgoorlie Subprovince the lower cycle has a preponderance of
71 thick mafic-ultramafic sequences containing tholeiitic basalts depleted in incompatible elements. The upper cycle has thinner mafic-ultramafic sequences containing tholeiitic basalts less depleted in some incompatible elements. Late-stage volcanic centres in the Southern Cross Subprovince contain Andean-type calc-alkaline assemblages. These secular variations point t o a progressive development from ensimatic to ensialic domains. ( 3 ) The LIL-element-depleted composition of the majority of volcanic rocks in the greestone belts cannot in itself be considered as proof for ensimatic crustal environment, because such rocks can occur as a minor component of continental traps and volcanic rocks of Proterozoic ensialic domains. However, as the two latter groups are dominated by high-LIL rock types, which are rare in Archaean terranes, the low-LIL content of Archaean volcanic rocks is more consistent with models advocating ensimatic to partly cratonized environments. (4)The occurrence of thick sequences of pillowed volcanic rocks accompanied by only a few sedimentary intercalations in the Kalgoorlie area may indicate rapid subsidence, because the scarcity of sediments suggests continuous extrusion, and the consistently pillowed nature of the lavas indicated persistence of subaqueous conditions. This observation, and the abundance at higher stratigraphic levels of immature turbidites and poorly sorted greywacke-conglomerates, suggest tectonically unstable basins dominated by strong vertical movements. In adjacent areas (Murchison Province, Southern Cross and Laverton Subprovinces) conditions may have generally been more stable. (5) N o basement is observed beneath the greenstone belts, and therefore the nature of the crust on which the supracrustal sequences were deposited is unknown. However, had this crust been dominated by felsic rocks, its subsidence and anatexis beneath the thick volcanic-sedimentary column should have resulted in the anatectic generation of differentiated LIL-enriched magmas (such as abound in Proterozoic ensialic domains) at early stages in greenstone belt evolution. No such volcanic rocks have so far been observed, and early dacite and Na-rhyolite observed in greenstone belts are compositionally indicative of derivation from basic source rocks or magmas. (6) The low Ri and low LIL-element levels of the older granodiorites intruded into the greenstones (Tables 2.V, 2.VI; Figs. 2.4 and 2.5) place limits on the composition and crustal history of their precursors. The low initial ratios limit the time span during which acid precursors could exist after their derivation from source materials with mantle-type R i values to no more, and possibly less, than 200 million years. The low LIL-element values make derivation of the granodiorites by anatexis of compositionally similar materials or of more differentiated parents unlikely, unless this was not accompanied by fractional crystallization and production of tonalitic residues and of highly-fractionated melts, a feature for which there is no explanation. Derivation of the granodiorites, and of possible coeval dacitic vol-
72 canic rocks within the greenstone belts, from greenstone-type assemblages by partial melting is consistent with geochemical and isotopic data. The transition with time from Na-rich t o K-rich intrusive rocks is consistent with progression from a simatic t o a sialic crustal environment. (7) Because the lowermost stratigraphic units in the greenstone belts contain felsic volcanic rocks, coeval hypabyssal and intrusive felsic dykes and derivative sediments it is evident that felsic igneous activity took place at least as early as the age of the oldest supracrustal rocks for which there is evidence. At the same time, the Na-rich dacitic to rhyolitic composition of these rocks (O’Beirne, 1968) suggests that they have themselves been derived from parental mafic rocks and/or magmas. We suggest that, throughout the history of the greenstone belts, episodes of melting in the mantle and the overlying roots of greenstone belts proceeded simultaneously - reflecting results of the same thermal events over a range of depths. (8) The ascent of the older granodiorites (c. 2.7 Ga old) was a major factor inducing isoclinal t o open folding and contact to semiregional low-pressure Abukuma-type metamorphism in the greenstone belts. This accounts for a decrease in metamorphic grade from anticlinal zones, juxtaposed with the intrusions, t o open-folded synclinal zones. The ascent of the younger, late to postkinematic adamellites of the Mungari-type reflects ensialic anatexis and was conceivably associated with the development of cross-folding and strain-slip cleavage. The growth of biotite along this cleavage suggests a late thermal elevation. There is thus evidence for polymetamorphism in part of the Kalgoorlie terrane. (9) The evolution of the Yilgarn greenstone belts spans at least 120 million years, i.e. between 2760 Ma and 2640 Ma, the ages of the Kambalda granodiorite and the Mungari granite, respectively (Oversby, 1976). N o lower isotopic age limits are available for lower mafic/ultramafic and sedimentary associations in the Kambalda-Kalgoorlie area and stratigraphic equivalents in other areas. (10) Whereas the lower mafic/ultramafic volcanic and sedimentary sequences are interpreted as the folded and metamorphosed relict of a onceextensive supracrustal layer, or simatic crust, it is likely that this layer/crust has evolved in the vicinity of previously cratonized sialic nuclei, such as the Southwestern Province (Gee, 1976), the Pilbara Block and possibly other smaller nuclei for which, however, no evidence is at hand. Such relations are consistent with concepts advocating progressive, but diachronous, transformation of simatic into sialic crust. These observations d o not appear to favour actualistic models. To date no evidence for lateral age and geochemical zonation, extensive thrusting, olistostromes and high-pressure glaucophane schist metamorphics - features diagnostic of plate collision sutures - has been recorded, casting doubt on the early Alpine geosynclinal model of McCall (1969) and Glikson (1970) and
73 the island arc-trench subduction model of White et al. (1971). The majority of Archaean structural data, with a few exceptions (Stowe, 1968) are indicative of vertical rat’ ,r than horizontal tectonic movements, the main sources of lateral stress and consequent folding being gravitative subsidence (Anhaeusser et al., 1969) and the ascent of anatectic granites. While important analogies between Archaean and Phanerozoic systems remain, the field is open for non-uniformitarian models consistent with the accumulated data and with modern petrological and geochemical experimental work. The occurrence of major impacts on the Moon and the terrestrial planets about 4.1-3.8 aeons ago (Schmitt, 1975) has given rise to speculations on possible genetic relations between early Archaean ultramafic-mafic volcanism and such impacts (Green, 1972; Glikson, 1976a, 1977). Fundamental differences between Archaean and younger mantle-crust relations are a distinct possibility. However, these subjects are outside the scope of this chapter. We conclude this section by stating that, in our view, no single method or criterion can alone form a basis for unravelling the complex granite-greenstone terrane, whose study demands the utmost coordination between field work and laboratory investigations. PILBARA BLOCK
Stratigraphy of the greenstone belts Early and current stratigraphic classifications of the layered Archaean succession of the Pilbara Block are summarized in Table 2.VII and some are shown in Figs. 2.7 and 2.8. Following the discovery of gold at Pilbara (110 km SSW of Port Hedland) in 1888, the geology in the vicinity of mining centres was mapped. Early concepts on regional stratigraphy were either derived from local observations or adopted from the Eastern Goldfields Province. Thus, in analogy with the division of the Kalgoorlie sequence into “greenstones” and “whitestones” (Honman, 1916; Forman, 1937), the Pilbara sequence was divided into: (1)Warrawoona Beds, which included the bulk of the volcanic rocks; and (2) Mosquito Creek Beds, including an arenaceous to argillaceous sequence east of Nullagine (Maitland, 1904, 1908; Woodward, 1911; David, 1932; Prider, 1948; Finucane, 1953). Early regional geological maps of the eastern Pilbara, based on work by the Geological Survey of Western Australia (Noldart and Wyatt, 1962), outlined the distribution of the Warrawoona Succession, the Mosquito Creek Succession, granites, granitized rocks, and dykes. The Mosquito Creek Succession was divided by these workers into five units, and a sequence of quartzite, conglomerate and banded iron-formation identified in the northern part of the Pilbara which had been correlated with the Mosquito Creek Succession was designated the Gorge Creek Formation (Table 2.VII). This unit was sub-
TABLE 2.VII Proposed stratigraphic schemes for and correlations of Archaean supracrustal sequence in the western and eastern parts of the Pilbara Goldfield, Western Australia ~ _ _ _ ____ .___ __ . 41 Western Pilbara Eastern Pilbara +b
______
Noldart and Wyatt (1962)
Ingram ( 1 9 7 2 )
Hickman and Lipple (1975)
Ryan ( 1 9 6 5 )
_____--
Fitton e t al. ( 1 9 7 5 ) ~~
Nullagine Succession
Nullagine Succession
Fortescue Group (Mt. Roe Basalt) Millindinna Complex Negri Volcanics
Fortescue Group (Mt. Roe Basalt)
Mosquito Creek Formation Budjan Formation Middle Creek Formation Eastern Creek Formation Dromedary Conglomerate
Mosquito Creek Formation
2 ~
b 2
Lalla Rookh Sandtone and Budjan Creek Formation
__
.
Gorge Creek Formation
2 Mallina Formation
e;
C3 Honeyeater Basalt
1 Paddy Mar-
01
-? l Gorge Creek Formation
Zleaverville Formation
1 1
ville
0
C3
I
Subgroup/
Gorge Creek Group Corboy For-
1 mation IWyman Formation
~WymanFormation
i Panorama and
I
Regal Formation
Kelly Forma-
Warrawoona Succession
B
Isharks Formation Chocolate Hill Formation
Igroup/
a/Empress Formation
Marble Bar Chert
2 Duffer Formation
*tFalga Talga Subgroup 1
Nickol River Formation [Williams, 1 9 6 8 )
Friendly Creek
1 Formation
-
75 sequently mapped by Low (1965) in the Port Hedland 1 : 250,000 sheet area, and similar units were recognized in the Mount Goldworthy and Soanesville areas. Maitland (1904, 1908) observed a major unconformity between the Mosquito Creek Beds and Nullagine Beds - the latter now dated as Lower Proterozoic (Compston and Arriens, 1968). Forman (1937) and Finucane (1953) identified an unconformity between the Warrawoona and Mosquito Creek successions. Noldart and Wyatt (1962, p. 105) considered the Gorge Creek Formation as “disconformably separated from the main (Warrawoona) succession by a short time break”. However, contrasting concepts concerning the relations between the main volcanic and sedimentary sequences have arisen as a result of regional mapping in the western part of the Pilbara Goldfield. Ryan (1965, 1966), Ryan and Kriewaldt (1964) and Kriewaldt and Ryan (1967) regarded these sequences as time equivalents, representing contemporaneous volcanic shelf and sedimentary trough facies, respectively (Fig. 2.8). In the Roebourne 1 : 250,000 sheet area the entire layered Archaean succession was designated Roebourne Group, which was subdivided into: (1) Regal Formation (mafic and felsic volcanic rocks); and (2) Cleaverville Formation (banded iron-formation and pelite). The synclinal sedimentary belt in the Mallina-Croydon area (Fig. 2.7) was thought t o be a lateral equivalent of the Regal-Cleaverville succession in the Roebourne and Whim Creek areas t o the north and west and the Pilbara Mining Centre to the south (Ryan, 1964, 1966). The Cleaverville Formation and sediments in the Dampier 1 : 250,000 sheet area to the west (Kriewaldt, 1964) were correlated with the Gorge Creek Formation. Williams (1968) redefined lower Regal Formation equivalents in the Yarraloola 1 : 250,000 sheet area as Nickol River Formation. Ingram (1972) mapped the Marble Bar, and parts of the Nullagine, Yarrie and Port Hedland 1 : 250,000 sheet areas, and divided the layered Archaean succession into five units (Table 2.V111), comprising: (1)Chocolate Hill Formation: 60-2500 m of picritic basalt, pyroxenite, peridotite and serpentinite occurring along granite margins and as roof pendants in the granites. (2) Sharks Formation: 300-1300 m of basalt, minor chert and banded iron-formation, minor felsic volcanic rocks and sediments; ultramafic flows and sills. (3) Miralga Formation: over 600 m of predominantly basaltic lavas with interflow chert, locally minor t o major felsic volcanic cycles and ultramafic intrusions, minor sedimentary beds; forms basalt-rhyolite cycles which may be repeated several times. (4) Gorge Creek Formation: chert, banded iron-formation, conglomerate, sandstone, greywacke, shale, talc-carbonate schist, minor interbedded basic and acid volcanic units.
76
'
0
KILOMETRES
I---+-_
100
I
-~~
' *- 1 &2 1 fLL1,I +
2 6 84 E 5 E z [--:'I ~
2j
~
~~-
~
-
~
0,
1 Protaroza,c cover
r
z
I
SIMPLIFIED GEOLOGICAL MAP OF THE PILBARA REGION
-_tl
Grsnitoidr Gorge C r e k Group Wyman Formation
j SjlgashSubgroup
STRATIGRAPHIC INTERPRETATION OF A PORTION OF THE EASTERN PILBARA AFTER HICKMAN A N D LlPPLE 11975) A N D HICKMAN (19756)
Duffer Formation
;]:
Talga Talga Subgroup
Fig. 2 . 7 . Stratigraphy in the Pilbara Block. A. Generalized geological sketch map of the Pilbara Block. B. Stratigraphic interpretation of a portion of the eastern Pilbara after Hickman and Lipple ( 1 9 7 5 ) and Hickman (1975b).
Fig. 2.7. Stratigraphy in the Pilbara Block. C. Stratigraphic interpretation of a portion of the western Pilbara after Fitton et al. (1975). Lithological descriptions of stratigraphic units in B and C are given in the text. 4 4
78
Fig. 2.8. Diagrammatic representation of facies changes in the Roebourne Group, western Pilbara, after Ryan (1965). This concept is not upheld, but is illustrated for the purpose ot‘ recognition of the development of ideas o n Pilbara stratigraphy. R = basement; s = sandstone; g = greywacke and shale; a = acid volcanics; cg = conglomerate; j = chert, jaspilite and carbonates; mu = basic volcanic rocks; c = Cleaverville Formation.
(5) Mosquito Creek Succession, divided (after Noldart and Wyatt, 1962) into: (a) Dromedary Conglomerate (including banded iron-formation pebbles). (11) Eastern Creek Formation (breccia, conglomerate, slate, quartzite). ( c ) Middle Creek Formation (shale, conglomerate, micaceous sandy beds); (d) Budjan Creek Formation (sandstone, shale). Ingram’s (1972) map delineates the distribution of these subunits of the Warrawoona Group over most of the eastern Pilbara. With reference t o the western Pilbara, Ingram has basically accepted Ryan’s (1964) concept of lateral volcanic-sedimentary facies transitions. The Mallina-Croydon trough sediments in the west are regarded as possible equivalents of the Mosquito Creek Succession in the east. Between 1972 and 1975 the Geological Survey of Western Australia undertook remapping of the Marble Bar, Nullagine, Yarrie and Port Hedland 1 : 250,000 sheet areas (Hickman, 1975a, b, 1977a; Hickman and Lipple, 1975; Hickman and Chin, 1977). For the first time, lithological rather than composite stratigraphical rock units were delineated - displaying the heterogeneous composition and internal complexity of the Warrawoona and Gorge Creek Groups. The stratigraphic classification emerging from this work (Lipple, 1975) comprises the following units (from base t o top): (1)‘I’alga-Talga Subgroup: 5-8 km of pillowed basalt with subordinate felsic, ultramafic and chert units.
79
(2) Duffer Formation: 5-8 km thick lenses of dacitic agglomerate, tuff and flows. ( 3 ) Salgash Subgroup: 5-9 km of pillowed basalt and chert intercalations, ultramafic rocks, including a prominant and stratigraphically consistent basal chert unit designated Marble Bar Chert, and stratigraphically intermediate felsic volcanic lenses defined as Panorama Formation and Kelly Formation. (4) Wyman Formation: 1 km of porphyritic rhyolite displaying columnar jointing. (5) Soimesville Subgroup: consisting of Corboy Formation (1-2 km of quartzite, sandstone and psammopelitic sediments) and Paddy Market Formation (1km of banded iron-formation and ferruginous clastic sediments). (6) Honeyeater Basalt: 0.5 km of pillowed basalt. ( 7 ) Lalla Rookh Sandstone: 2-3 km of arkosic sandstone and polymictic conglomcrate . llnits 1-4 are parts of the Warrawoona Group and 5-7 of the Gorge Creek iroup. Fhckman and Lipple (1975) traced the stratigraphically lowermost volcanic unit - the Talga-Talga Subgroup - around, as well as xenoliths within, the major granitic batholiths (Fig. 2.7). Hickman (1975a) utilized these xenoliths in subdividing the felsic plutonic complexes into individual granitic domes. Their maps also show that the dacitic agglomerates of the Duffer Formation form large lenticular bodies which are missing in several areas (Fig. 2.7B). I t will be suggested later that the Duffer Formation and the intrusive granodiorites could be comagnetic expressions of the same acid igneous episode. One of the most significant contributions made by Hickman and his coworkers is the delineation of the.Marble Bar Chert over most of the eastern Pilbara region. This conspicuous and stratigraphically consistent unit enables the correlation of underlying, associated, and overlying volcanic units and enables palaeogeographic reconstructions, and localization of stratigraphically controlled mineralization. I. Martin (pers. commun., 1975) has observed the divergence, in places, of two or three chert members of the Marble Bar Chert, outlining pyroclastic lenses which represent “fossil” volcanic cones contemporaneous with the chert. Equivalents of the Marble Bar Chert were also identified in the western Pilbara, where the unit is denoted Hong Kong Chert (Fitton et al., 1975). The extensive occurrence of the Marble Bar Chert and underlying units over most of the Pilbara region -an area about 500 km long and 100-200 km wide -demonstrates that their deposition was not confined to the linear greenstone belts (Hickman, 1975a). Only at a later stage were greenstone belts structurally defined. The appearance of granite-derived sediments in the Lalla Rookh Sandstone of the Gorge Creek Group signifies uplift and erosion of granites associated with the formation of inter-batholithic depressions. Hickman and Lipple (1975) regarded the units of the Warrawoona Group
80
as essentially internally conformable, but note local unconformities at the base of the Gorge Creek Group. In contrast, Ingram (1972) and Fitton et al. (1975) suggested that the banded iron-formation of the Gorge Creek Group represent a chemical sedimentary facies continuous with, and marking the termination of, the Warrawoona volcanic megacycle. Fitton e t al. (1975) recognized a regional unconformity above the banded iron-formations (Paddy Market Formation of Hickman and Lipple, 1975). In the southeastern Pilbara this unconformity is defined at the basal contact of the Mosquito Creek Group. In the northern Pilbara it correlates with granite- and jaspilite-derived polymictic conglomerate and arkose at the base of the Lalla Rookh Sandstone. In the western Pilbara the unconformity is commonly intruded by gabbro and dolerite sills of the Millindinna Complex (Fitton et al., 1975). Overlying this unconformity in the western Pilbara is the Whim Creek Group (Fig. 2.7C) which includes: (1)Warrambie Basalt: 200 m of amygdaloidal t o vesicular basalt and andesite. (2) Mons Cupri Volcanics: 500 m of felsic lavas, agglomerate, tuff, intercalated volcanic sediments, some intrusive felsic rocks. (3) Constantine Sandstone: 500 m of coarse-grained quartzo-feldspathic sandstone, pebbly arkosic grit and coarse-grained greywacke. (4) Mallina Formation: 2500 m of shale, greywacke and turbidite. The Whim Creek Group is intruded by granite, and both are unconformably overlain by the Negri Volcanics - a sequence of continental basalts and acid lavas, including spinifex textured rocks, fine-grained sediments and local conglomerates toward the top. The Negri Volcanics could represent extrusive equivalents of the Millindinna Complex - - a suite of late Archaean or early Proterozoic ultramafic t o gabbroic and granophyric sills and pods (Fitton et al., 1975). The field relations demonstrated by Fitton et al. (1975) in the western Pilbara conflict with the earlier view of lateral transition between the volcanic Teichmans Group (Warrawoona Group equivalent) and sedimentary sequences (Gorge Creek Group, Whim Creek Group and Mosquito Creek Group) as conceived by Ryan (1964, 1966), Ingram (1972) and Miller (1975). The predominance of Constantine Sandstone and Mallina Formation in the Mallina-Croydon belt and Mosquito Creek belt reflect the stratigraphically high position of these units within the synclinal structure, and in the lack of evidence for lateral interfingering between greenstones and clastic sedimentary sequences there appears to be no reason to infer a primary absence of Teichmans or Warrawoona Group volcanics beneath these synclines. Relations between the Mons Cupri Volcanics and the Constantine Sandstone-Mallina Formation sequence within the Whim Creek Group are less clear. Fitton et al. (1975) consider the sediments as either stratigraphically higher than, or laterally equivalent with, the felsic volcanic piles. In con-
81 trast, Hickman (1977b) considers the felsic volcanic rocks as younger than the Lalla Rookh Sandstone. The problem is complicated by faulting and by outliers of Negri Volcanics south of Whim Creek. Another point of view concerning stratigraphic correlation between the eastern and western Pilbara, cited by Button (1976), considers the acid Mons Cupri Volcanics as equivalents of the Duffer Formation, and the Negri Volcanics as equivalents of the Salgash Subgroup. However, this concept is inconsistent with the young isotopic age of the Mons Cupri Volcanics (R.C. Horwitz, pers. commun., 1976). Stratigraphic correlations are thus hindered by the apparent absence in the eastern Pilbara of volcanic counterparts of the Whim Creek Group - a feature which suggests a lenticular configuration of the felsic volcanic rocks. Other problems involved in the correlation of eastern and western sequences are local changes in the abundance of chert and banded iron-formation in the Gorge Creek Group and variations in the thickness of the Marble Bar Chert, rendering their positive identification uncertain in places.
Structure and metamorphism Noldart and Wyatt (1962) mapped the characteristic circular and arcuate granite dome-greenstone belt pattern in the Marble Bar and Nullagine 1 : 250,000 sheet areas (Figs. 2.7 and 2.9), which they interpreted in terms of cross-folding of the Warrawoona Succession followed by intrusion of granites into domal positions in crests of the cross-folds - a view based on a recognition of early folds truncated by granites. Thus, tight assymetrical fold limbs and part-domal structures were regarded as segments of larger pre-granite structures. The tectonic role of the batholiths was thought t o be limited t o the intensification of the early structures. According t o this concept, Noldart and Wyatt outlined axes of hypothetical NNE- and superposed WNW-trending fold systems, interpreted in terms of a compressive southeasterly stress field. Notwithstanding this interpretation, Noldart and Wyatt (1962, p. 126) made the observation that “no overall regional trend is apparent within the (Warrawoona) succession alone, the greater majority of the internal structural trends and fold patterns following marginally on the plutons, and reflecting the outlines of the plutons. This relationship is consistent throughout the succession. A marked wrap-around pattern is present in the jaspilites fringing the larger plutons, and noticeable domal structures are developed around the smaller stocks. Distant from the plutons the bedding trends indicate large synclinorial trough structures with subsidiary anticlinal and synclinal structures developed on the flanks. The central portion of cores of most of these structures are infilled with remnants of younger successions. ” In contrast to the pronounced NNW strike of both greenstones and elongate granite axes in the Eastern Goldfields Province of the Yilgarn Block, an original tectonic trend is difficult t o decipher in the Pilbara. Hickman (1975a) has shown that the structure of the greenstones is fully explicable in
82
Fig. 2.3. ERTS image of granite-reenstone stone relationships are clearly shown.
terrane in the eastern Pilbara. Granite-green-
83 terms of a vertical emplacement of the batholiths, and that no evidence for cross-folding is at hand. Thus, no fold interference structures were observed in areas where linear greenstone belts merge, forming triple or quadrupole junctions (Figs. 2.7 and 2.9). Continuous flexure of both bedding and metamorphic foliation has taken place within the layered succession in conformity with the broad plutonic outlines, although, in detail, local truncation of layered structures by the intrusive granite can be observed. Close relations between the structures of granites and greenstones have also been observed in the western Pilbara, and Ryan (1965) has commented o n the conformable relations between the stratification of the Roebourne Croup and the banding and foliation in the adjacent gneisses. In the western Pilbara, the overall tectonic trend of the granite-greenstone system is easterly to northeasterly, as distinct from the mainly northerly trend in the eastern part of the craton. Throughout the Pilbara, discordant faulting is of relatively minor importance, and major faults mostly parallel the principal tectonic trends or coincide with granite-greenstone contacts, e.g. Sholl fault-Sherlock Bay fault, and faults along the western margin of the Shaw Batholith and the southern margin of the Mount Edgar Batholith. These faults are associated with wide zones of extreme attenuation and mylonitization, which followed late magmatic stages and post-consolidation vertical movements during late Archaean and earliest Proterozoic times. These movements resulted in a preferential juxtaposition of basal outliers of the Fortescue Group (Mount Roe Basalt and Hardey Sandstone) with greenstone synclinona and scarcity of such outliers above granitic domes in the eastern Pilbara (Kriewaldt, 1964; Ryan, 1965; Hickman, 1975a). These relations suggest differential isostatic adjustments in the granite-greenstone crust, involving uplift of the low-density granitic sectors and subsidence of the high-density greenstone belts. It will be suggested below that such differential movements played a key role in the evolution of the greenstone belts themselves. Evidently, granite-greenstone contact zones - serving as long-acting discontinuities between rheologically contrasted rigid (granite) and ductile (greenstone) domains - served as loci of repeated dislocation after the principal granite-forming events about 3 Ga ago. The preservation in the Pilbara Block of roof zones of the batholiths in some areas, and of deeply eroded levels of the plutons in other areas renders this terrane significant t o the elucidation of the three-dimensional anatomy of the Archaean intrusions. Examples of high-level zones are: (1)the North Pole dome, where little-deformed volcanic rocks of the Warrawoona Group (Talga Talga and Salgash Subgroups) dip at low angles around a core of intrusive adamellite; and (2) McPhee dome, where volcanics of the Talga Talga Subgroup and Duffer Formation form a domal structure which, from gravity evidence, is probably underlain by granite (Hickman, 1975a). Intermediate to deep levels of the batholiths display subsidiary gneiss domes sepa-
84 rated from each other by sheared synclinal roots of deeply eroded greenstone belts, suggesting an originally polydomal configuration of the batholiths. Projection downwards of greenstone belt structure, and projection upwards of the intervening batholiths, suggests a depth of 30-40 km and a height of 15 km, a.s.1. respectively (Hickman, 1975a). If these extrapolations apply, and since little deformation of the Archaean crust has taken place since the Lower Proterozoic (as shown by the intact structure of dykes and supracrustal cover deposits), the original thickness of the Archaean crust, if equated with the batholiths, may have exceeded 50 km. Of course, this consideration is based on an assumed Archaean age of infracrustal root zones of the batholiths, which, though likely, cannot be confirmed by direct observation. Hickman (1975a) carried out a structural analysis of parts of the granitegreenstone system in the eastern Pilbara, proposing the following tectonic classification: (a) D1 - all structures considered to have pre-dated the main deformation (D2) associated with emplacement of the granitic batholiths; (b) D2 - structures related t o granite emplacement and concomitant subsidence of the greenstone belts, including tight isoclinal folds and steep cleavage sets in supracrustal sequences and concentric foliation in the granites; (c) D3 - strain-slip cleavage striking NE, and a conjugate set of NE- and NW-trending faults displaying dislocations suggestive of N--S compression; (d) D4 - E-W trending subhorizontal crenulations of S2 foliations; (e) D5 - NW- to NE-trending folds effecting the Proterozoic cover, WNWand ENE-striking faults. Both the folding and the faulting imply a maximum E-W compression. Progressive transitions take place from heavily sheared, amphibolite facies granite-intruded (or faulted) margins of the greenstone belts into littledeformed lower greenschist or prehnite-pumpellyite facies synclinal core zones. McCall (1971) described little-metamorphosed to unmetamorphosed ultramafic sills from the Soanesville syncline, as contrasted to garnet-bearing contact aureoles juxtaposed with the granites. The sympathetic relations between metamorphic grade and degree of strain are linked t o thermal and mechanical effects of granite emplacement, respectively, and to subsequent boundary dislocations. Thus, metamorphic patterns in the Pilbara are related in time and place to the granitic batholiths, rather than a regional metamorphic zonation. At the same time, exceptions to the sympathetic metamorphic grade-strain relations are known; for example, little-deformed amphibolite facies rocks may occur along some granitegreenstone contacts. It will be suggested below that, conceivably the Salgash Subgroup, and certainly the Gorge Creek Group, have postdated the intrusion of the major batholiths about 3 Ga ago. If so, it follows that these sequences could not have been thermally affected by the intrusions, and their metamorphosed
85 state must be ascribed to later thermal rise, possibly associated with the c. 2.7-2.6 Ga-old adamellites. This concept is tentatively supported by the more common retention of primary pyroxene in the Salgash Subgroup than in the Talga Talga Subgroup. However, insufficient isotopic and mineralogical data are presently available t o confirm the possibility of such a metamorphic discontinuity. Granitoids Little work has been carried out on granitic terranes in the Pilbara, attention having been focussed on the auriferous and better-exposed greenstone belts. However, excellent exposure of parts of the batholiths render the Pilbara a particularly suitable terrain for detailed study of their structure, petrology and geochemistry. The state of knowledge of the Pilbara granites in the 1940’s was summarized by Forman (1937), who stated: “Granite and gneiss, of which no detailed study has been made, are intrusive into both the Warrawoona and the Mosquito Creek Series”. Maitland (1908), Brandt (1964), Noldart and Wyatt (1962), Hickman (1975a), and M.J. Fitton and G. Sylvester (in Fitton et al., 1975) consider the batholiths as consisting of remobilized older pre-greenstone basement. Miller (1975) regards the basement of volcanic arcs of the western Pilbara as sea-floor basalt. Kriewaldt (1964) compared the oval plutons to Eskola’s mantle gneiss domes. Ryan (1965) favoured an origin of some batholiths by the remobilization of sediments, a view shared by Prider (1965). Clearly, contrasting views exist on the nature and origin of these little-studied rocks. On their map Noldart and Wyatt (1962) divided the granitic rocks into two categories: “granite and granite gneiss” and “granitized Archaean rocks”. The latter consist of wide zones which flank the main batholiths, and their highly sheared state, concordance with the greenstone structure, and abundant greenstone xenoliths were noted. A distinction was made between concordant granites intruded into the Warrawoona Succession and small discordant bodies which pierce the Mosquito Creek Succession (see also Ryan, 1965). De Laeter and Blockley (1972) reported geochronological and petrographic differences between the major Mount Edgar Granodiorite batholith and the smaller Moolyella Adamellite intruded into it. The Mount Edgar complex is typical of the large, dome-like, concordantly foliated intrusions, represented also by the Corrunna Downs, Shaw, Carlindi, Turner, Yule, Portree, Satirist, Caines Well, and George batholiths (Fig. 2.7), whose ages fall within the 3.1-2.9 Ga range (Table 2.V). The Moolyella Adamellite is typical of younger, massive, in places mushroom-shaped, often tin-bearing adamellites whose distribution in the Pilbara was delineated by Blockley (1970), and whose ages fall within the range of 2.7-2.6 Ga (Table 2.V). Other examples of this class of intrusions are the Mondana, Spear Hill, Coondina, Cooglegong and Numbana adamellites. The Mondana Adamellite is
86 intruded along the boundary of the Corunna Downs batholith. A similarly located adamellite occurs along the northern margin of the Yule Batholith south of the Pilbara Mining Centre, and like the Moolyella, Spear Hill, Coondina and Cooglegong intrusions is a source of pegmatitic and alluvial tin deposits. The older granodiorites and the younger adamellites form contrasting terrains: the former underlie smooth t o low-undulating plains, and crop out as isolated boulders, low benches and in creek beds. The drainage pattern is fine-dcndritic, displaying little or no control by joints. The younger adanicllites form rugged, joint-controlled, elevated tors and ridges. De Laeter and Blockley (1972) have described the transition from the rims to t h e centre of the Mount Edgar batholith: the margins are dominated I,y migmatite and lit-par-lit injection gneiss intruded into amphibolites of the Warrawoona Group in zones up t o i km wide. Inwards the migmatite grades i n t o well-banded gneissic granite, which passes in turn into foliated granite. The contacts of the younger adamellite are sharp, and cut across the banding of the older granite. Xenoliths and roof pendants of the latter may occur in the adamellite. The latter is accompanied by swarms of albite pegmatites, the primary source of alluvial cassiterite deposits. Plagioclase is more abundant than microcline in the foliated granite, whose other components are biotite, quartz, and minor t o accessory hornblende, clinozoisite, sphene, chlorite, epidote, zircon, apatite and iron oxide. Amphibole may be rimmed by biotite. Deformation and mylonitization of quartz is common. The Moolyella Adamellite consists of quartz, microcline, sodic plagioclase, biotite, muscovite, minor hornblende and chlorite, and accessory epidote, fluorite, and rarer topaz, apatite, zircon and spessartine garnet. Zoned plagioclase, interleaved biotite-chlorite crystals, and biotite-rimmed amphiboles are common. Little deformation is evident. De Laeter and Blockley (1972) noted departures from the regular granite zonation pattern, e.g., an occurrence of migmatite a t Moolyella, 10 km from the margin of the Mount Edgar batholith. Hickman and Lipple (1975) and Hickman (1975a, b) have shown that these occurrences are related t o a subdivision of the batholith into several subsidiary domes (Fig. 2.7). Much like marginal zones of the batholiths, the synclinal zones between the subsidiary granite domes are heavily foliated (the foliation converging toward the synclinal axis), containing abundant greenstone xenoliths which may form discontinuous trails, and are in places accompanied by shears. These zones merge in places with the greenstone belts, and Hickman (1975a) has demonstrated that they represent synclinal keels of the latter. In our impression, these xenoliths are derived largely from the Talga Talga Subgroup, i.e. the oldest exposed unit of the greenstone belts. Hickman (1975a) also notes occurrences of xenoliths away from interdomal synclines, which he alternatively attributes to an earlier generation of greenstones or t o older granitegreenstone contacts obliterated by younger intrusions. Toward the centres of the domes the foliation weakens and only a weak biotite alignment remains;
87 in places primary flow structure is observed. On the basis of the spatial relations between the younger adamellites, subsidiary gneiss domes and granite-greenstone boundaries, Hickman (1975a) arrived at a most significant observation, namely, a preferential juxtaposition between the younger intrusions and older granite-greenstone boundaries or their deeply eroded, xenolith-rich roots. Such relations apply t o the Moolyella, Mondana and Cooglegong adamellites and their equivalents in the western Pilbara, but are not without exception, e.g. emplacement of the Spear Hill Adamellite into the core of a granodiorite dome within the Shaw batholith. On this basis Hickman (1975a) distinguished three structural levels in y o u n p r granites: (1)deep levels, associated with intrabatholithic synclinal zones; (2) intermediate levels, along granite-greenstone contacts; and (3) high levels, emplaced within the greenstones. Hickmari and Lipple (1975) and Hickman (1975a), as an integral part of a 1 : 250,000 mapping project, airphoto interpretation, and petrographic study, divided the granites of the Marble Bar and Nullagine sheet areas into 12 types: ( 1)Early migmatitic complex: dominated by granodioritic to tonalitic rocks, but also including adamellite. Nebulitic, stromatic, veined, gneissic and agmatitic textures. Gneiss and greenstone xenoliths. Contacts with the greenstones are either intrusive or faulted. ( 2 ) Granodiorite: including both biotite- and hornblende-bearing types. Metamorphic texture. In part occurring as xenoliths in the early migmatites. Many greenstone xenoliths. Characteristic aplite veins. ( 3 ) Alkali-feldspar granite: medium- to coarse-grained, locally porphyritic, intrusive into greenstones. (4) Granodiorite: medium-grained, equigranular, poorly foliated, some mafic schlieren. (5) Biotite tonalite to biotite-hornblende granodiorite: abundant contamination by greenstone xenoliths. (6) Biotite adamellite and biotite granodiorite. (7) Remobilized migmatitic rocks: equigranular, fine- to coarse-grained biotite adamellite concordant with and enclosed by migmatites. Abundant mafic schlieren and greenstone xenoliths. Preserves foliation direction of enclosing migmatites. (8) Granophyre and alkali-feldspar granite: equigranular, medium- to coarse-grained, poorly foliated, strongly jointed. (9) Porphyritic biotite adamellite and granodiorite: well-foliated to locally massive, intrudes greenstones and contains xenoliths. (10) Equigranular biotite adamellite: massive to poorly foliated, fineto medium-grained; intrudes migmatites; strongly jointed. (11) Medium- t o coarse-grained, equigranular biotite adamellite and biotite granite: massive to poorly foliated, locally porphyritic; some xenoliths and schlieren; igneous texture well preserved; strongly jointed; sharp contact
88
against older rocks; source of cassiterite-rich pegmatites. (12) Medium, equigranular albite adamellite: massive to poorly foliated; allotriomorphic-granular; fine-grained margins; sharp intrusive contacts with the early migmatites. Units 1-10 are regarded as pre- to syn-tectonic, and, whereas the early migmatites and granodiorites yield ages in the 3.1-2.9 Ga range, Hickman and Lipple (1975) consider the porphyritic biotite adamellite and granodiorite of unit 9 to have been emplaced as a series of plutons between 3.1 and 2.7 Ga ago. In contrast, units 11 and 1 2 are clearly post-tectonic, and are represented by the Moolyella and Cooglegong adamellites emplaced about 2.7-2.6 Ga ago (Table 2.V). Little information is available t o date on western Pilbara granites, although attention has been given to the problem of granite-greenstone relations in this area (Ryan, 1965; Fitton et al., 1975). The latter authors delineate concentric marginal zones of gneiss and contaminated (greenstone xenolith-bearing) granite around the Portree and Caines Well batholiths, and parts of the Yule and George granites. Internal gneiss zones were charted within the Caines Well and Harding batholiths. The map indicates the occurrence of tonalite and hornblende granite around the Satirist pluton and the western part of the Portree pluton. However, systematic documentation of these intrusions must await further field work. Major and some trace element data for Pilbara granites have been reported by Hickman and Lipple (1975) and Oversby (1976), and some representative analyses are tabulated in Table 2.VI. Plots of K20/Na20 against Rb-Sr and Pb-Pb ages reveal a positive trend, a feature shared by granites from the Eastern Goldfields (Fig. 2.6). Some of the noteworthy geochemical characteristics are: (a) Migmatites from the Shaw and Mount Edgar batholiths (mostly biotite granodiorites) - low K20/Na20;low Rb/Sr; high Li. (b) Mulgandinnah adamellite (syntectonic) - K20/Na20> 1;Rb/Sr = 1.6. (c) Moolyella and Cooglegong adamellites (post-tectonic) - K20/Na20> 1;very high Rb/Sr (>lo); high Sn and Li. The significance of geochemical and isotopic parameters of the granites will be discussed below.
Geoch ro nology Prior t o the advent of isotopic age determinations, concepts on the age of Pilbara rocks were based on lithological and structural similarities t o classic divisions of the Precambrian in Canada and Scandinavia. Thus David (1932) assigned the Warrawoona Series to the Archaeozoic, the Mosquito Creek Series t o the Older Proterozoic, and the Nullagine Series to the Younger Proterozoic. The earliest attempt at isotopically dating Pilbara rocks is due t o
89 Holmes (1927, p. 38) who used the U-Th-Pb method in dating hydrothorite from pegmatite at Wodgina, obtaining an age of 1260 Ma. Subsequent age determinations were reviewed by Compston and Arriens (1968), and later contributions by Hickman (1975a). Principal milestones in this work include: (1)A single granite Rb-Sr age of 3040 Ma (maximum age) by Leggo et al. (1965). (2) A Wodgina pegmatite age of 2900 Ma deduced from Rb-Sr isotopic measurements on muscovite, microcline and lepidolite. K-Ar ages for these minerals are about 2100-2300 Ma old (work by Jeffrey, Greenhalgh and Jeffrey, Riley - reported in Leggo et al., 1965). (3) A whole-rock Rb-Sr isochron age for twelve granites collected from the eastern and western Pilbara (including granite in the Dampier area, and the Satirist, Peaway and Corunna Downs batholiths) yielded a value of 3050+-180 Ma (Compston and Arriens, 1968). (4)A Rb-Sr isochron age for six samples of the Copper Hills Porphyry (Duffer Formation, east of the Coruna Downs batholith) yielded a value of 2880 f 66 Ma with a high Ri value of 0.7303 (de Laeter and Trendall, 1971). (5) Whole-rock Rb-Sr work on the Mount Edgar batholith yielded a poorly defined age of 3125 f 366 Ma for foliated granodiorite and 2670 f 95 Ma for a post-tectonic adamellite intruded into the foliated granodiorite at Moolyella (De Laeter and Blockley, 1972). Similar age relations were documented in the Shaw batholith (De Laeter et al., 1975). Oversby (1976) determined Pb and some Rb-Sr isotope ages of gneiss and granodiorite from the continuous granitic terrane SSE of Port Hedland. In four localities early thermal events about 2950 Ma old are represented, and in at least two of these localities secondary thermal events were superposed about 2600-2770 Ma ago. The granites of three localities reflect repeated isotopic disturbances 2900-2250 or 2900-2000 Ma ago - some of these possibly representing effects of Lower Proterozoic thermal events. Pidgeon (1978) has determined a precise U-Pb discordia age of 3453 f 1 6 Ma on zircons from Duffer Formation dacite collected at Sharks Gully, about 25 km SSW of Marble Bar. Significantly younger Rb-Sr isochron ages of about 2200 Ma were obtained for this rock. Published age data and R ivalues are given in Table 2.V. With the exception of the Copper Hills Porphyry, all suites 3100-2800 Ma-old have low to very low R i values, indicating a relatively short crustal history for their precursor materials (Fig. 2.5). The high R i value of the Copper Hills Porphyry (0.730 f 0.0119) was explained by De Laeter and Trendall (1971) in terms of metamorphic age resetting and a high Rb/Sr ratio of the original volcanic rocks. A crustal history not longer than 60 Ma prior t o metamorphism was suggested. In contrast, the high R i values of the Moolyella and Cooglegong adamellites (Table 2.V) are consistent with their derivation by anatexis of the older granodiorites (De Laeter and Blockley, 1972; De Laeter et al.,
90 1975). In view of Pidgeon’s (1978) results, the question of the igneous ages of Pilbara granites must await U-Pb zircon studies. Relatively little isotopic work has been conducted in the western Pilbara; in particular, the primary age of the Whim Creek Group, Negri Volcanics and Millindinna Complex are unknown, nor are there many data on the granites. The Mount Brown Rhyolite (a member of the Mons Cupri Volcanics) previously suggested as c. 3000 Ma old (Compston and Arriens, 1968) has been reinterpreted as younger by about 500-700 Ma (Fitton et al., 1975).
Mineralization The Pilbara granite--greenstone system is a host for important gold, copper, iron, tin and barite deposits, and for some chrysotile, antimony-gold, molybdenum-copper, copper-lead-zinc and tantalite-columbite deposits. Minor occurrences of wolframite, beryl, lepidolite, spodumene, and fluorite have been mined. Silver, copper, lead, zinc and cobalt are minor byproducts of gold processing. Ingram (1972) classified economic mineral occurrences in the eastern Pilbara in terms of their host units and mode of occurrence. The Warrawoona Succession contains three principal associations of precious and base metals: (1)Strata-bound gold-pyrite: associated with carbonatized tuffs (Warrawoona Mining Centre), carbonatized basalts (Bamboo Creek, North Shaw), carbonatized basic intrusions (Talga Talga, Sharks Well), and altered volcanogenic ultrabasic rocks (Bamboo Creek, Pilgangoora, Halley’s Comet). Auriferous zones are often closely associated with chert intercalations. Ore localization has been largely controlled by strike fault intersections and drag folds. In Ingram’s view, the gold is of volcanic fumarolic origin, and has been secondarily concentrated during deformation. (2) Strata-bound base metals: copper sulphides and Pb-Zn-Ag deposits are principally associated with the upper parts of felsic volcanic piles, particularly the Split Rock Formation (Duffer Formation), often beneath chert beds, and are localized by shears, faults and fractures. Examples are the Copper Hills, Copper Gorge (Reedy Creek Caldera), and Big Stubby (Pb-ZnAg). Stratigraphically higher occurrences are Kelly (copper) and Whim Creek (copper), located within acid volcanic rocks of the Salgash Subgroup and in the Mons Cupri Volcanics, respectively (Table 2.VII). ( 3 ) Nickel sulphides: nickel arsenides and sulphides occur in sheared carbonatized ultrabasic rocks, for example at Bamboo Creek, Cookes Creek, Salgash, Fielding Gully, Sherlock River (west Pilbara). No economic nickel deposits are known to date. Other gold mining centres in the Pilbara located within units of the Warrawoona Succession include North Pole (where Cu, Ba, Bi, Ag also occur), Sharks Gully, Yandicoogina, Twenty Ounce Gully, Lalla Rookh, Tambourah, Western Shaw, North Shaw and McPhee’s Reward workings, most
91 of which are abandoned at present. In the western Pilbara gold has been recovered from the Roeborne Group at Pilbara Mining Centre (including Teichmans and Hong Kong mines), between Pilbara and Station Peak, near Roeborne and at Nickol River. Another important mineral recovered from the Warrawoona Succession is asbestos, mined mainly at the Lionel Mining Centre, north of Nullagine, where it occurs as veins in serpentinized peridotite, and is also known at Nunyerry, Sherlock and Mount Hall in the western Pilbara. The uppermost sedimentary units of the Archaean succession contain in places important gold ores, for example, the Eastern Creek, Billjim, Blue Spec (Au, Sb), and Middle Creek mines within the Mosquito Creek Succession. Antimony-gold ores also occur in slates near Mallina (western Pilbara), at stratigraphic levels probably equivalent with the Mosquito Creek Succession. Copper-iron sulphides occur at Egina and near Croydon (Evelyn Mine) at similar stratigraphic positions, and could represent sedimentary counterparts of the Whim Creek deposits. The banded iron-formation of the Paddy Market Formation (Gorge Creek Group) is a host of large, supergene enriched iron deposits currently mined at Mount Goldsworthy and Shay Gap. The younger, c. 2.7-2.6 Ga-old, adamellites which intrude the older granodiorites and migmatites include pegmatites rich in cassiterite, tantalitecolumbite, and in places beryl, spodumene, lepidolite and gadolinite. Principal mining areas are Moolyella, Cooglegong, Hartigan Creek, Coomba Creek. Hillside, Mondana, Wodgina and Yule River. Other ore-bearing intrusions are mafic to ultramafic sills of the Millindina Complex, which contain titaniferous and vanadiferous magnetite at Balla-Balla and near Roebourne (Fitton et al., 1975).
Evolution o f the Pilbara craton The relative stratigraphic integrity of greenstone belts, the lateral continuity of units around batholiths, their structurally intact, little-faulted nature, and the excellent exposure of supracrustal and often plutonic units render the Pilbara granite-greenstone terrane one of the most complete Archaean crustal segments preserved anywhere. Recent advances in the regional mapping of the greenstones and granites in the eastern and western parts of the Pilbara, coupled with progress in geochronological work, now provide numerous constraints on models of evolution. Current geochemical investigations of the Warrawoona Group (Bureau of Mineral Resources and Geological Survey of Western Australia) promise t o provide further insight into the primary and superposed characteristics of the volcanic rocks. Theories relating t o the origin of granite-greenstone systems hinge to a major extent on interpretations of the nature and origin of their incorporated batholiths. Whilst the volcanicsedimentary nature of Pilbara green-
92 stone successions has been long recognized, the granitoids are variably interpreted as remobilized basement domes antedating the greenstones ( Maitland, 1908; Noldart and Wyatt, 1962; Kriewaldt, 1964), composite intrusions which include both reactivated basement rocks and younger plutons (Hickman, 1975a; Fitton et al., 1975), granitized sediments (Ryan, 1965; Prider, 1965), or as composite intrusions including older mantle-derived phases and younger ensialic anatectic phases (De Laeter and Blockley, 1972; De Laeter et al., 1975; Glikson and Lambert, 1976). The identification of the oldest recognizable components of the Pilbara granite-greenstone system is a necessary step in the elucidation of respective sequences of events within the greenstone belts and batholiths and their relative time correlations. However, no simple answer can be given t o this problem, since both the greenstone belts and the granitoid batholiths constitute composite, long-acting domains whose evolution has t o a large degree overlapped. The greenstone belts contain several important discontinuities, and represent “negative” crustal zones whose subsidence relative to the batholiths probably commenced before 3 Ga ago, and whose lowermost volcanic units are known to be still older (Pidgeon, 1978). These belts still constituted negative crustal domains in the earliest Proterozoic, as evidenced by the preferentially juxtaposed outliers of Fortescue Group volcanics. Likewise, the granitic batholiths evolved during 3.0-2.6 Ga ago, could conceivably contain enclaves of yet older granites (Hickman, 1975a), constituting long-acting “positive” crustal domains. Because neither detailed subdivisions of the greenstones and the granitoids were known prior t o 1970, and earlier concepts on the evolution of the Pilbara craton were based on a few data, the following discussion is restricted to the interpretation of recent studies. Hickman (1975a) has documented the intrusive or faulted relations between the granites and the greenstones; however, he suggested that the intrusive relations may pertain only to younger annular granitoids intruded along unconformable surfaces between a hypothetical older basement and the greenstones. He further recognizes early D,-structures in the granitoids which predate D,-structures related t o their emplacement. However, no relicts of the proposed unconformity were observed, nor have isotopic or geochemical differences been detected between the 3 Ga-old granodiorite/ migmatite and hypothetical pre-greenstone basement. Because the batholiths consist of subsidiary domes separated by trains of greenstone xenoliths this interpretation requires that any relicts of a pre-greenstone granitoid basement may only exist at the cores of these domes. M.J. Fitton and G. Sylvester (in Fitton e t al., 1975) have emphasized the generally concordant relations between granitic batholiths and greenstone belt structures in the western Pilbara, which they interpret in terms of diapiric reactivation of an older basement. The gneissic structure of the plutons was cited in support of this interpretation. In contrast, R.C. Horwitz
93 (in Fitton et al., 1975) stressed the intrusive contacts of the granites and the occurrence of greenstone xenoliths within them, suggesting an older age of the greenstones. It must be borne in mind that the emplacement of composite magmatic complexes - involving synmagmatic, late magmatic and postconsolidation deformations - results in polyphase deformation, which, although related to the igneous event itself, yields an impression of old age. Thus the ascent and emplacement of progressive magmatic increments results in superposed foliation and fold sets that yield no evidence regarding the overall time relation between the intrusions and the greenstone sequences. As pointed out by De Laeter and Blockley (1972) and De Laeter et al. (1975) the low Ri values of the Mount Edgar and Shaw granodiorites indicate derivation from materials with mantle-type isotopic characteristics, thus limiting the crustal history of precursors of the granites. No c. 3.0 Garocks with high Ri values - i.e. derived from a pre-greenstones basement have been encountered to date within the batholiths. Moreover, it must be borne in mind that reactivation of granitic crust by anatexis must result in relatively differentiated melts of high K/Na and Rb/Sr, unlike the granodiorites which form the bulk of the 3.0 Ga-old batholiths. In considering the origin of the Pilbara granite-greenstone system, a possibility is the correlation between the older granodiorites and the dacitic agglomerates of the Duffer Formation as coeval plutonic and extrusive manifestations of one and the same igneous episode. Some evidence for such relations is provided by: (1)the similarity between Rb-Sr ages of the Copper Hills Porphyry and other outcrops of Duffer Formation; (2) the intrusive relations between the granodiorites and the Talga Talga Subgroup as contrasted to the apparent absence of authenticated intrusive relations with the Salgash Subgroup; (3) the general chemical similarity between the granodiorites and the dacites (analysis in Hickman and Lipple, 1975). The zircon age of 3.45 Ga on Duffer Formation dacite (Pidgeon, 1978) casts doubt on this correlation, unless granitoid ages are significantly older than reported, or if the evolution of the batholiths has spanned the period 3.453.0 Ga ago *. Further work is needed to establish this possibility. If correct, the above correlation indicates the following sequence of events: (a) Pre-3.45 Ga: Formation of a regional volcanic sequence consisting of mafic, altramafic, andesitic and dacitic lavas and intercalated chemical sediments (chert, banded iron-formation, carbonate). N o direct evidence is at hand on the nature of the basement of this succession, which in the Pilbara is at least 5 km thick (Talga Talga Subgroup, Friendly Creek Formation), but the ensimatic characteristics of the sequence d o not favour the existence of underlying sial. Had sialic basement existed, its subsidence and partial * N o t e added in proof: Pidgeon (pers. commun., 1 9 7 9 ) reports U-Pb zircon ages of 3417 i 4 0 Ma for migmatites of the Shaw Batholith and 3280 f 20 Ma for foliated granitics of the Mt. Edgar Batholith.
94 melting beneath the thick greenstone column should have resulted in the formation of large volumes of differentiated acid magmas. The latter are rare in the Warrawoona Group, and are mainly known at its top. (b) 3.45-3.0 Ga ago: Deformation, downbuckling, subsidence and partial melting of root zones of the early volcanic crust resulted in progressive formation of large volumes of dacitic magma, erupted as extrusive agglomerates (Duffer Formation) and emplaced as diapiric plutons. This was followed by the region-wide deposition of chert, formed from the fine felsic tuff and by colloidal precipitation from silica-saturated sea water (Marble Bar Chert). (c) 3.0-2.7 Ga ago: Progressive structural definition of interbatholithic troughs and concomitant mafic and ultramafic volcanic activity concentrated in, but not confined to, these belts (Salgash Subgroup). Accompanying intermittent dacitic volcanic activity. Termination of volcanism marked by tkposition of region-wide banded iron-formations (lower part of Gorge (-’reek Group). Further differential vertical movements between the “positive” batholiths and the “negative” greenstone belts resulted in erosion of cmierging granites, development of a regional unconformity ( Fitton et al., 1975), and supply of the granitic detritus t o upper units of the Gorge Creek Group (Lalla Rookh Sandstone, Constantine Sandstone). This was accompanied by deposition of shales and turbidites within synclinal troughs (Mosquito Creek Formation) and by felsic volcanic activity in the west (Mons Cupri Volcanics). The commencement of this sedimentary cycle was accompanied in places by mafic volcanism ( Honeyeater Basalt, Warrambie Basalt). (d) 2.7-2.6 Ga ago: Rising geothermal gradient resulted in ensialic palingenesis of the older granodiorites giving rise to adamellites and accompanying low-grade metamorphism in the greenstone belts. In this model the metamorphism of the Talga Talga Subgroup is mainly related t o c. 3.0 Ga old thermal effects by the intrusive granodiorites, whereas metamorphism of the sequence which postdates the Marble Bar Chert is related t o the 2.7-2.6 Ga old geothermal rise. If this concept is correct, polyphase metamorphism should be restricted t o the lower stratigraphic units. This expectation furnishes a possible isotopic test of the validity of the above sequence of events. Current U-Pb zircon and Sm-Nd isotope studies and geochemical investigations (BMR and GSWA) should provide further constraints on models for the origin of the Pilbar granitegreenstone system. PILBARA-YILGARN
RELATIONSHIPS
Prior to the application of isotopic age determination techniques t o Pilbara rocks, the basic analogy between the lithologies and stratigraphy of greenstone belts in the two cratons was regarded as evidence for their contemporaneity - an approach exemplified by the extension of stratigraphy from the Kalgoorlie area to the Pilhara terrane (Maitland, 1904, 1908;
95 Woodward, 1911; David, 1932). However, following the measurement of ages of about 3.0 Ga in the Pilbara, as contrasted to ages of about 2.7-2.6 Ga in the Eastern Goldfields (Leggo et al., 1965; Compston and Arriens, 1968), the Pilbara craton was generally regarded as older. Subsequent Rb-Sr work by de Laeter and Blockley (1972), De Laeter et al. (1975) and Oversby (1976) showed that both c. 3.0 Ga-old and c. 2.7-2.6 Ga-old granites occur in the Pilbara. The relations between the two terranes are now a central question in the study of the Precambrian of Western Australia. McCall (1971) pointed out that age differences between the Pilbara and the Yilgarn relate to granite intrusion and metamorphism, but do not necessarily reflect different ages of the greenstones themselves. Relations between the Kalgoorlie and Pilbara granite-greenstone terranes could be expressions of the exposure of different crustal levels of the same shield (Glikson and Lambert, 1976). However, such a coeval relationship is now thought t o be unlikely because the Pilbara craton contains stratigraphically and structurally high-level supracrustal rocks affected by low-grade metamorphism in both synclinal and domal positions, a situation suggestive of shallow crustal levels. The pronounced differences in structural style between the two terranes, strongly linear in the Yilgarn and oval to arcuate in the Pilbara, remain unexplained. The distance between the exposed edges of the two terranes is about 300 km, and it is not clear whether their structural difference arises by a gradual transition in the basement underlying the Hamersley and Bangemall basins, or alternatively whether a sharp break exists - signifying an existence of distinct crustal plates in the Archaean. The problems of the relations between the two cratons were discussed by Smith and Horwitz (1975) and Horwitz and Smith (1978) with reference to the palaeogeographical reconstruction of the Hamersley and Nabberu Basins and t o gravity data (BMR, 1976). These authors favour an original continuity of the Yilgarn and Pilbara granite-greenstone terranes, but interpret the positive gravity anomaly which underlies the western part of the Ophthalmian zone (Wyloo Trough) in terms of an underlying mafic slab, possibly formed upon incipient and arrested separation of the cratons. A continuity of epicontinental stratigraphy between the Hamersley and Nabberu Basins wocld preclude crustal separation after the onset of development of the Mount Bruce Supergroup. On the other hand, plate separation prior t o this stage remains a possibility, and Horwitz and Smith (1978) refer t o the extensive mafic volcanic activity represented by the Fortescue Group and related dykes as a possible manifestation of such events. The problem of the relations between the Pilbara and Yilgarn cratons was considered by Gee (1975), who favoured an earlier development of the former, and by Rutland (1976) who likewise suggested that the Warrawoona Succession antedates the Kalgoorlie greenstones. The latter author suggested a possible correlation between banded iron-formation of the Gorge Creek Group and basal banded iron-formation of the Kalgoorlie and Southern
96 Cross greenstone sequences, namely, a contemporaneous development of volcanic sequences in the Yilgarn and sedimentary sequences in the Pilbara. It has been suggested that the Warrawoona Group includes pre-3.0 Ga-old phases, granitoid-related acid volcanic phases, and post-granitoid phases. As in lower greenstone cycles of the Yilgarn Block, only minimum isotopic ages are known for early greenstone units in the Pilbara (e.g., Talga Talga Subgroup *, Friendly Creek Formation), which could therefore be contemporaneous. However, on present evidence, the ages of the main granitoids in the Pilbara and the Eastern Goldfields Province (Tables 2.V and 2.VII), differ significantly, indicating earlier cratonization of the northern Archaean province. Conceivably, differences between the magnitude and orientation of regional stress fields during principal granitoid intrusion phases are the underlying cause for the different structures of the granite-greenstone terranes in the Pilbara and Yilgarn blocks. It follows that the Proterozoic Ophthalmian-Gascoyne mobile belt coincides with, and may be genetically related to, a s i a l s i m a boundary zone which existed during the 3.0-2.7 Ga period between a cratonized region in the north and a simatic region in the south. REFERENCES Anhaeusser, C.R., Mason, R., Viljoen, M.J. and Viljoen, R.P., 1969. A reappraisal of some aspects of Precambrian shield geology. Geol. SOC.Am. Bull., 80: 2175-2200. Archibald, N.J. and Bettenay, L.F., 1977. Indirect evidence for tectonic reactivation of a pre-greenstone sialic basement in Western Australia. Earth Planet. Sci. Lett., 3 3 : 370378. Arriens, P.A., 1971. The Archaean geochronology of Australia. Spec. Publ. Geol. SOC. Aust., 3 : 11-23. Barnes, R.G., Lewis, J.D. and Gee, R.D., 1974. Archaean ultramafic lavas from Mount Clifford. West. Aust. Geol. Surv. Annu. Rep. 1973: 59-70. Baxter, J.L., 1975. Achaean banded iron formation. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5: 202-204. Bennett, R., Page, R.W. and Bladon, G.M., 1975. Catalogue of isotopic age determinations on Australian rocks, 1966-1970. Aust. Bur. Miner. Resour. Rep. 162. Binns, R.A. and Marston, R.J., 1976. Archaean geology of the Yilgarn Block, Western Australia. 25th Int. Geol. Congr., Sydney, 1976, Guideb. Excursion 40A. Binns, R.A., Gunthorpe, R.J. and Groves, D.I., 1976. Metamorphic patterns and development of greenstone belts in the eastern Yilgarn Block, Western Australia. In: B.F. Windley (Editor), The Early History of the Earth. John Wiley, London, pp. 303-313. Blockley, J.G., 1970. Preliminary report on tin granites in the Pilbara Goldfield. West. Aust. Geol. Suw. Annu. Rep., 1969: 34-37. BMR (Bureau of Mineral Resources, Australia), 1976. Gravity Maps of Australia, 1 : 5,000,000. Canberra, A.C.T. Brandt, R.T., 1964. The iron ore deposits of the Mount Goldsworthy area, Port Hedland
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* Note added in proof: Recently Hamilton et al. (in press) reported a Sm-Nd age of 3560 i 32 Ma for the Talga-Talga Subgroup.
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99 Precambrian crust in Western Australia. Tectonophysics, 30 : 55-89. Glikson, A.Y. and Sheraton, J.W., 1972. Early Precambrian trondhjemitic suites in Western Australia and northwestern Scotland and the chemical evolution of shields. Earth Planet. Sci. Lett., 1 7 : 227-242. Gole, M.J., 1976. Metamorphosed banded iron-formations in the Yilgarn Block, Western Australia. 25th Int. Geol. Congr. Sydney, 1976, Abstr., 1: 11-12. Goodwin, A.M. and Shklanka, R., 1967. Archaean volcano-tectonic basins: form and pattern. Can. J. Earth Sci., 4:777-795. Gower, C.F., 1971. Porphyritic dolerites of the Edjudina sheet area. West. Aust. Geol. Surv. Annu. Rep., 1970: 36-37. Gower, C.F., 1973. Explanatory notes on the Laverton 1 : 250,000 geological sheet, Western Australia. Geol. Surv. West. Aust. Rec. 1973/28, 33 pp. (unpubl.). Green, D.H., 1972. Archaean greenstone belts may include terrestrial equivalents of lunar maria? Earth Planet. Sci. Lett., 15: 263-270. Hall, H.I.E. and Bekker, C., 1965. Gold deposits of Norseman. In: J. McAndrew (Editor), Geology of Australian Ore Deposits. Australas. Inst. Min. Metall., Melbourne, pp. 101-107. Hallberg, J.A., 1972a. Geochemistry of Archaean volcanic belts in the Eastern Goldfields region of Western Australia. J. Petrol., 13: 45-56. Hallberg, J.A., 1972b. Spilitic pillow lavas at Mt. Hunt, Western Australia. J.R. SOC. West. Aust., 55: 1-4. Hallberg, J.A., 1976. A petrochemical study of a portion of the western Yilgarn Block. CSIRO Miner. Res. Lab. Div. Mineral. Rep. FP 13, 38 pp. Hallberg, J.A. and Williams, D.A.C., 1972. Archaean mafic and ultramafic rock associations in the Eastern Goldfields region, Western Australia. Earth Planet. Sci. Lett., 15: 191-200. Hallberg, J.A., Johnston, C. and Bye, S.M., 1976a. The Archaean Marda igneous complex, Western Australia. Precambrian Res., 3 : 111-137. Hallberg, J.A., Carter, D.N. and West, K.N., 1976b. Archaean volcanism and sedimentation near Meekatharra, Western Australia. Precambrian Res., 3 : 577-595. Hickman, A.H., 1975a. Precambrian structural geology of part of the Pilbara region. West. Aust. Geol. Surv. Annu. Rep., 1974: 68-73. Hickman, A.H., 1975b. Explanatory notes on the Nullagine 1 : 250,000 geological sheet. Geol. Surv. West. Aust. Rec. 1975/5 (unpubl.). Hickman, A.H., 1977a. Explanatory notes on the Port Hedland 1 : 250,000 sheet. Geol. Surv. West. Aust. Rec. 1977/2 (unpubl.). Hickman, A.H., 1977b. Stratigraphic relations of rocks within the Whim Creek belt. West. Aust. Geol. Surv. Annu. Rep., 1976: 53-56. Hickman, A.H. and Lipple, S.L., 1975. Explanatory notes on the Marble Bar 1 : 250,000 geological sheet. Geol. Surv. West. Aust. Rec. 1974/20 (unpubl.). Hickman, A.H. and Chin, R.J., 1977. Explanatory notes on the Yarrie 1 : 250,000 geological sheet. Geol. Surv. West. Aust. Rec. 1976/16 (unpubl.). Holmes, A., 1927. The Age of the Earth. Ernest Benn Ltd., London. Honman, C.S., 1916. The geology of the country to the south of Kalgoorlie. Geol. Surv. West. Aust. Bull., 66, 76 pp. Horwitz, R.C. and Smith, R.E., 1978. Bridging the Yilgarn and Pilbara Blocks, Western Australia. Precambrian Res., 6 : 293-322. Horwitz, R.C. and Sofoulis, J., 1965. Igneous activity and sedimentation in the Precambrian between Kalgoorlie and Norseman, Western Australia. Australas. Inst. Min. Metall. Proc., 214: 45-59. Hudson, D.R., 1972. Evaluation of genetic models for Australian sulphide nickel deposits. Australas. Inst. Min. Metall., West. Aust. Conf. Pap., pp. 99-109.
100 Ingram, P., 1972. Geology of Portion of the Pilbara Goldfield, Western Australia. Peter Ingram and Associates, Perth (unpubl.). Jaques, A.L., 1976. An Archaean tholeiitic layered sill from Mt. Kilkenny, Western Australia. J. Geol. SOC.Aust., 23 : 157-168. Jutson, J.T., 1934. The physiography of Western Australia. Geol. Surv. West. Aust. Bull., 95. Kriewaldt, M.J.B., 1964. Dampier and Barrow Island, Western Australia. Geol. Surv. West. Aust. 1 : 250,000 Geol. Series Explan. Notes. Sheets SF/50-2 and SF/50-1. Kriewaldt, M.J.B. and Ryan, G.R., 1967. Pyramid, Western Australia. Geol. Surv. West. Australia 1 : 250,000 Geol. Series Explan. Notes. Sheet SF/50-7. Larking, A., 1969. Acid Rocks and their Relationship t o the Gold Mineralization at Norseman, Western Australia. B.Sc. (Hons.) Thesis, University of Adelaide (unpubl.). Leggo, P.J., Compston, W. and Trendall, A.F., 1965. Radiometric ages of some Precambrian rocks from the northwest division of Western Australia. J. Geol. Soc. Aust., 12: 53-65. Lewis, J.D., 1971. Spinifex texture in slag as evidence for its origin in rocks. West. Aust. Geol. Surv. Annu. Rep., 1970: 45-49. Lipple, S.L., 1975. Definitions of new and revised stratigraphic units of the eastern Pilbara region. West. Aust. Geol. Sum. Annu. Rep., 1974: 58-64. Low, G.H., 1965. Port Hedland, Western Australia. Geol. Surv. West. Aust. 1 : 250,000 Geol. Series Explan. Notes. Sheet SF/50-4. Maitland, A.G., 1904. Preliminary report on the geological features and mineral resources of the Pilbara Goldfield. Geol. Surv. West. Aust. Bull., 15. Maitland, A.G., 1908. The geological features and mineral resources of the Pilbara Goldfield. Geol. Surv. West. Aust. Bull., 40. Marston, R.J., 1976. Geochemistry of Archaean felsic t o ultramafic metasediments in relation t o crustal evolution, northeastern Yilgarn Block, Western Australia. 25th Int. Geol. Congr. Sydney, 1976, Abstr., 1: 17-18. Marston, R.J. and Travis, G.A., 1976. Stratigraphic implications of heterogeneous deformation in the Jones Creek conglomerate (Archaean), Kathleen Valley, Western Australia. J. Geol. SOC.Aust., 23: 141-156. Mathur, S.P., 1976. Crustal structure in southwestern Australia from seismic and gravity data. Tectonophysics, 24: 151-182. McCall, G.J.H., 1969. The Archaean succession to the west of Lake Lefroy. J.R. SOC. West. Aust., 52: 119-128. McCall, G.J.H., 1971. Some ultrabasic and basic igneous rock occurrences in the Archaean of Western Australia. Spec. Publ. Geol. SOC.Aust., 3 : 429-442. McCall, G.J.H., 1972. The nickel-sulphide-bearing ultramafic rocks and their environment in the Archaean of Western Australia. 24th Int. Geol. Congr., Sect. 1: 354-362. McCall, G.J.H., 1973. Geochemical characteristics of some Archaean greenstone suites of the Yilgarn structural province, Australia. Chem. Geol., 11: 243-269. McCall, G.J.H. and Leishman, J., 1971. Clues t o the origin of Archaean eugeosynclinal peridotites and the nature of serpentinisation. Spec. Publ. Geol. SOC.Aust., 3: 281299. McCall, G.J.H., Braybrook, J.C. and Middleton, D.D., 1970. The Merougil Creek sub-area. J.R. SOC.West. Aust., 53: 9-32. McKinstry, H.E., 1945. Precambrian problems in Western Australia. Am. J. Sci., 234A: 448-466. McMath, J.C., 1953. The geology of the country about Coolgardie, Coolgardie Goldfield, Western Australia. Geol. Surv. West. Aust. Bull., 107. Miller, L.J., 1975. Archaean geosyncline of the Pilbara. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5: 55-63.
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103 Williams, I.R., 1969. Structural layering in the Archaean of the Kurnalpi 1 : 250,000 sheet area, Kalgoorlie region. West. Aust. Geol. Surv. Annu. Rep., 1968: 40-41. Williams, I.R., 1970. Kurnalpi, Western Australia. Geol. Surv. West. Aust. 1 : 250,000 Geol. Series Explan. Notes. Sheet SH/51-10. Williams, I.R., 1974. Structural subdivision of the Eastern Goldfields Province, Yilgarn Block. West. Aust. Geol. Sum. Annu. Rep., 1973: 53-59. Williams, I.R., 1975. Eastern Goldfields Province. In: Geology of Western Australia. Geol. Surv. West. Aust. Mem., 2: 33-54. Williams, I.R., 1976. Regional interpretation map of the Archaean geology southeast part of the Yilgarn Block. Geol. Sum. West. Aust. 1 : 1,000,000 map. Williams, I.R., Gower, C.F. and Thom, R., 1976. Edjudina, Western Australia. Geol. Sum. West. Aust. 1 : 250,000 Geol. Series Explan. Notes. Sheet SH/51-6. Wilson, A.F., 1958. Advances in the knowledge of the structure and petrology of the Precambrian rocks of southwestern Australia. J.R. SOC.West. Aust., 41: 57-83. Winchester, J.A. and Floyd, P.A., 1976. Geochemical magma type discrimination: application to altered and metamorphosed basic igneous rocks. Earth Planet. Sci. Lett., 28: 459-469. Woodall, It., 1965. Structure of the Kalgoorlie Goldfield. In: J. McAndrew (Editor), Geology of Australian Ore Deposits. Australas Inst. Min. Metall., Melbourne, pp. 71-79. Woodall, R., 1975. Gold in the Precambrian shield of Western Australia. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5: 175-184. Woodward, H.P., 1911. The geology and ore deposits of the west Pilbara Goldfield. Geol. Surv. West. Aust. Bull., 41. Reference a d d e d in P r o o f : Hamilton, P.J., Evenson, N.M., O’Nions, R.K., Glikson, A.Y. and Hickman, A.M., 1980. Sm-Nd dating of the Talga-Talga Subgroup, Warrawoona Group, Pilbara Block, Western Australia, 2nd lnt. Archaean Symp., Perth, 1980. Ahstr. Vol., p. 11.
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Chapter 3
PROTEROZOIC GEOLOGY OF WESTERN AUSTRALIA
*
A.D.T. GOODE
INTRODUCTION
Knowledge of the Proterozoic geology of Western Australia is at present in a state of flux with major new discoveries altering our understanding of all terrains. The most notable of these has been the discovery of the Nabberu Basin in late 1973, but other major advances have also taken place in the Hamersley and Bangemall Basins. More complex terrains such as the Fraser, Albany and Gascoyne Blocks have also been the subject of new work. Consequently, the summary presented here must at best be seen as an arrested view point of a rapidly changing picture. At present the focus of most attention is on the intracratonic basins and fold belts separating the Archaean Yilgarn and Pilbara Blocks. In the present climate it was, therefore, felt best to attempt to present as far as possible an objective review of the available geological data, and then to try and relate this to more interpretive regional pictures of geological evolution. This latter aim will obviously be much more subjective, but it is hoped that it will provide a basis for stimulating discussion by other workers in the field. The Proterozoic rocks of Western Australia considered in this chapter are those from the Western Australian Shield (Gee, 1976); Proterozoic units €rom the northern Australian Kimberley region and the central Australian Arunta-Musgrave area are described in this volume by Plumb et al. (Chapter 4) and Rutland et al. (Chapter 5). The Precambrian Western Australian Shield is surrounded by the undeformed Phanerozoic sediments of the Canning, Officer and Eucla Basins to the east, and the Perth and Carnarvon Basins t o the west. In keeping with the recommendations of James (1978), the Archaean-
* Published by permission of Dr. R.S.H. Fardon, General Superintendent Exploration, The Broken Hill Proprietary C o . Ltd.
106
Fig. 3.1. Tectonic subdivision of the Western Australian Shield.
107 Proterozoic boundary is taken a t 2500 Ma. The Archaean basement to the Western Australian Proterozoic has been described by Hallberg and Glikson (this volume, Chapter 2). The boundaries between Early, Middle and Late Proterozoic are taken at 1750 Ma and 1000 Ma respectively (Glaessner, 1966). Apart from thin basal sediments of the Officer Basin (Lowry et al., 1972), evidence of Late Proterozoic deposition is lacking on the Western Australian Shield. REGIONAL TECTONIC UNITS
The tectonic sub-division of the Precambrian Western Australian Shield has been discussed by several authors, e.g. Daniels and Horwitz (1969), Trendall (1975a), Gee (1976; 1979), Rutland (1976), Horwitz and Smith (1978), Plumb (1979). For the purposes of this review, the Shield is divided into three major units (Figs. 3.1 and 3.2).
I
/
-
Grawty hrgh
-Grovrty low
....... Dipole Ainorndy
- 0
500km
Fig. 3.2. Synoptic map of Bouguer gravity anomalies (after Gee, 1979), showing also the location of the steep gradients associated with the major dipole gravity anomalies (after Wellman, 1978). (With permission of Tectonophysics and The Director, Bureau of Mineral Resources).
108 Archaean granite-greenstone
terrains
The Archaean of the Yilgarn and Pilbara Blocks (Hallberg and Glikson, this volume, Chapter 2) consists essentially of metavolcanic/metasedimentary (greenstone) belts intruded by -2600 Ma granites. Both terrains also contain -3000 Ma granites, of indeterminate relationship t o the greenstones in the Yilgarn but definitely intrusive into the greenstones in the Pilbara. The Pilbara greenstones (Warrawoona Group) have been dated -3400 Ma (Pidgeon, 1978). In the Yilgarn the “older” high-grade granite gneisses are confined to the western margin of the block. The Pilbara also contains “younger” metasedimentary sequences (e.g., Gorge Creek Group, Mosquito Creek Group) which predate the Early Proterozoic sequences. Structures within the Archaean basement are significant in the depositional and deformational history of the subsequent Proterozoic basins. Conversely it is hoped to show that the Proterozoic basinal and fold belt configurations are important in helping t o unravel one of the most fundamental and critical problems in the Western Australian Archaean, viz. the relationships between the apparently younger Yilgarn Block (dominated by 2600 Ma granites) and the apparently older Pilbara Block (dominated by 3000 Ma granites). The results have significance in similar terrains in North America and South Africa, and probably elsewhere.
Intracratonic Proterozoic basins and fold belts The Pilbara and Yilgarn Blocks are separated by several successions indicative of repetitive development of depositional basins followed by deformational episodes (Daniels, 1966a). This intracratonic belt has been termed the Median Belt by Horwitz and Smith (1978), and corresponds to changes in the major trend pattern of gravity anomalies-(Wellman, 1976). The Early Proterozoic portion of this zone has been termed the Capricorn Orogen by Gee (1979). However, it may be preferable to employ the term Capricorn Tectonic Zone for both Early and Middle Proterozoic rocks in this belt, as this would better indicate their common physical relationships. The Early Proterozoic basins comprise the well-known Hamersley Basin flanking the Pilbara Block (Macleod, 1966; Trendall, 1976a) and the recently discovered Nabberu Basin flanking the Yilgarn Block (Hall and Goode, 1978). Both basins are notable for the development of major thicknesses of banded iron formations characteristic of the Early Proterozoic. Both basins become increasingly deformed away from the Archaean basement blocks, resulting in the formation of the Ophthalmian Fold Belt in the north and the Stanley Fold Belt in the south, together with the intrusion of granites. Increasing metamorphic grade is associated with increasing degree of deformation in both cases. A t present, two of the most critical problems which may help resolve the relationships between the Pilbara and Yilgarn cratons
109 and the Capricorn Tectonic Zone are: (a) the stratigraphic relationships between the Hamersley and Nabberu Basins; and (b) the structural relationships between the Ophthalmian and Stanley Fold Belts. The small Middle Proterozoic Bresnahan and Mt. Minnie basins (Daniels, 1975a) postdate the Ophthalmian folding, and may be resultant molasse deposits. The Middle Proterozoic Bangemall Basin occupies the central part f the intracratonic basinal area, covering the critical zone between the arnersley and Nabberu Basins and their respective fold belts (Brake1 and Muhling, 1976; Williams et al., 1976; Goode and Hall, in prep.). It has itself been deformed in the Late(?) Proterozoic by the Edmundian Fold Belt in the west and the Paterson Fold Belt in the northeast. Other minor occurrences of sedimentary rocks, at least some probably related to the major Proterozoic successions, occur along the western and southern margins of the Yilgarn Block.
k
Proterozoic metamorphic complexes
The Yilgarn Block is flanked on its eastern, southern and western margins by Middle Proterozoic metamorphic complexes or “mobile belts”. These complexes include the arcuate Albany-Fraser Province of high-grade metamorphics and granites (Doepel, 1975; Bunting et al., 1976). Other lesser important complexes occur along the western margin of the Yilgarn Block and within the Paterson Fold Belt. The Gascoyne Block has been included in this group, although it may eventually be better included with the Early Proterozoic basins and fold belts. The geochronologic, stratigraphic and structural relationships of these terrains to the other Precambrian rocks of the Western Australian Shield continues to promote vigorous discussion and controversy. INTRACRATONIC BASINS
Early Proterozoic Hamersley Basin
The Hamersley Basin sediments (known as the Mt. Bruce Supergroup) comprise thick clastic, chemical and volcanic sediments, outcropping over a present area of about 100,000 km2 (Fig. 3.3). The main basin can be structurally divided into a central basin to the south of the Archaean Pilbara basement, and a subsidiary extension to the northeast of the Rat Hill Anticline (or basement ridge). Minor sub-basins at the margins of the main basin include the Peawah, Nullagine, Meentheena and Yarrie sub-basins. Outliers of the lower parts of the Mt. Bruce Supergroup also occur within the Pilbara Block (e.g., near Marble Bar).
111 The Mt. Bruce Supergroup comprises the basal Fortescue Group (dominantly basic volcanics with lesser clastics), the middle Hamersley and Turee Creek Groups (banded iron formations, shales, dolomites, dolerites and acid volcanics) and the upper Wyloo Group (clastics, carbonates and volcanics). The Wyloo Group is unconformable on the underlying sequences. In the northeast, the Manganese Group forms the uppermost sequence unconformably above the lower Hamersley Group, and is probably equivalent to the Wyloo Group. The present structure of the basin comprises very shallow dipping, essentially undeformed sequences on the Pilbara Block which become more folded and metamorphosed to the south and east. Strong faulting separates much of the Fortescue-Hamersley sequences from the Wyloo Group along the western parts of the basin. Basement doming was also active in this region prior to and during Wyloo deposition. The basinal axis of deposition appears to have been displaced to the south at this time. Previous work in the Hamersley Basin is summarized in Macleod (1966), Trendall (1975b, c, 1976a), and Trendall and Blockley (1970), and the following discussion draws heavily on this work by the Geological Survey of Western Australia. Recent mapping by various investigators has led t o additional information and some stratigraphic revision, particularly in the northeastern Hamersley Basin, and at the base of the Wyloo Group in the southwestern part of the basin.
Fortescue Group The stratigraphic nomenclature and principal lithologies of the Fortescue Group are summarized in Table 3.1 (after Trendall, 1975b, 1976a). Trendall (1976a) has noted that the Fortescue Group requires reassessment to resolve fully all problems of intrabasinal correlation. The bulk of the Fortescue Group is directly or indirectly volcanogenic, and dominantly mafic. Extensive flows of dark, fine-grained, massive, amygdaloidal or vesicular basalt, locally pillowed, are separated by equally extensive sheets of mafic tuff, much of which has clearly accumulated in shallow water. Long periods of lava extrusion broadly alternated with long periods of volcanoclastic accumulation, t o give a broad stratigraphic alteration of lava and tuff, as distinct from continual thin intercalation of the two lithologies (Trendall, 1976a). Smith (1976,1979) has recognized numerous altered flow units averaging 50 m thickness within the Maddina Volcanics. Felsic and acid lavas also occur, and many tuffs have an admixture of terrigenous clastic (often granitic) debris. Shallow-water and stromatolitic carbonate intercalations are locally well developed, particularly in the northeast. The Jeerinah Formation, and more particularly the Roy Hill Shale Member, may be regarded as an exception to the bulk of the Fortescue Group in that its mafic volcanic affinity is not demonstrable, it is widespread, and its
112 TABLE 3.1 Stratigraphy of the Fortescue Group (After Trendall, 1976, and others) Unit
Formation
Maximum thickness * (m)
Principal lithologies
Upper sediments (fine clastics)
R o y Hill Shale Member Jeerinah Formation Nallanaring Volcanic (Lewin Member Shale) Warrie Member Woodiana Sandstone Member
u p t o 1000
shale (black pyritic), siltstone, chert, subordinate sandstone, some volcanic intercalations
Maddina Basalt (Bunjinah pillow 600-800 Lava; Pearana Basalt; “Upper” Little de Grey Lava)
Upper lavas
I
m
.-0
massive to amygdaloidal basic and intermediate lava with subordinate tuffaceous siltstone and cher; locally pillowed
A
2 Pillingini Tuff (Kuruna Siltstone,
200-650
stratified tuffaceous sandstone and siltstone with cross bedding and ripple marks, minor basalt, stromatolitic limestone
200-800
massive t o amygdaloidal basic and intermediate lavas, locally pillowed
Lower sediments (mediumgrained clastics)
Hardey Sandstone (Cliff Springs Formation with Lyre Creek Agglomerate Memher; Beatons Creek Conglomerate; Isabella Formation?)
0-1300
tuffaceous and terrigenous sandstones and siltstones, local conglomerate and agglomerate
Lower lavas
Mt. Roe Basalt
0-300
amygdaloidal, vesicular, and massive basalt, locally columnar; basal sandstone
Middle sediments (pyroclastics)
4
Nymerina Basalt, Tumbiana Pisolite, Pyradie Pyroclastics)
cs 4
Middle lavas
7
1
Kylena Basalt (Boongal Pillow Lava; “Lower” Little d e Grey Lava)
*Thickest sequences usually found in Mt. Bruce 1 : 250,000 Sheet (De la Hunty, 1965); Mt. Roe Basalt confined t o the northern margin of the basin (Fig. 3 . 3 ) .
lithology is closely similar t o the shales of the Hamersley Group. Like many of these it is, where fresh, black due to the presence of free carbon (up to 17%), commonly pyritic and cherty. Although thin lavas are intercalated
113 locally even at this level the volcanicity which contributed directly to the first phase of basin development either waned or manifested itself as thick sills of both basic and acid material injected at various levels within the sequence. These include the Gidley Granophyre, the Cooya Pooya Dolerite, the Spinaway Porphyry and the Bamboo Creek Porphyry. Some Fortescue units are quite distinctive and unusual. For example the Tumbiana Pisolite is a thick, accretionary lapilli tuff probably similar t o the Zwartkoppie spheroidal rocks of the Archaean Onvenvacht Group in South Africa (Trendall, 1965; Lowe and Knauth, 1978).Hardey Sandstone/Beaton’s Creek Conglomerate equivalents in the Nullagine Sub-basin contain small sub-economic concentrations of uraninite in carbonaceous pellets within fluvial pyritic conglomerates and sandstones (Richards, 1972). Detrital gold also occurs in the basal fluvial Beaton’s Creek Conglomerate in the same subbasin. These deposits have many similarities to the classical Late Archaean Au/U-bearing pyritic quartz-pebble conglomerates of the Witwatersrand (South Africa) and Huronian (Canada) Basins (e.g., Minter, 1978). The lateral continuity of the lowest stratigraphic units in the Fortescue Group is generally restricted, and their thickness varies sharply along strike. Along the northern margin of the basin, the basal Mt. Roe Basalt, for example, is confined to palaeovalleys in the topography (now preserved as subbasins), see Fig. 3.3. Higher in the sequence individual units tend t o be more easily traceable over large strike distances, with less variation in thickness, suggesting progressive “drowning” of the palaeotopography by the flood basalts. The palaeovalleys can be recognized by their thicker, more complete stratigraphic sequences and by the fluvial nature of their clastic sediments (e.g., Richards, 1972; Hickman and De Laeter, 1977). In the Yarrie palaeovalley (or sub-basin) for example, flat-lying basal Fortescue Group sediments are preserved within relatively steep-sided hills of Archaean greenstones. Hickman and Lipple (1975) have noted that the palaeolows tend to be developed preferentially on the Archaean greenstone belts with the granite plutons forming relative palaeohighs. Hickman and Lipple also believe that the greenstone belts were progressively down-warped during Proterozoic sedimentation; similar tectonosedimentary effects have also been noted in the Early Proterozoic Transvaal (Anhaeusser and Button, 1976) and Nabberu (Hall and Goode, 1978) Basins.
Hamersley Group In contrast to the Fortescue Group the stratigraphic sub-divisions and nomenclature of the Hamersley Group, initially established by Macleod et al. (1963), are clearly defined and demonstrably valid over its entire outcrop area; only in a few marginal localities, and in the northeastern part of the basin, is there any uncertainty of internal correlation. The subdivisions of Macleod et al. (1963), as modified by Trendall and Blockley (1970), appear in Fig. 3.4.
114
Fig. 3.4. Stratigraphic columns of the Hamersley Group (left) and the Dales Gorge Member of the Brockman Iron Formation (right), showing numbering and thicknesses in metres of the 33 macrobands. Lithology is represented as follows: BIF - solid black; shale and siltstone including tuffaceous material - white; dolerite - diagonal hatching; acid volcanic rocks - v; dolomite - brick pattern. The vertical divisions in the Mount McRae Shale and Dales Gorge Member in the left-hand column indicate diagrammatically the approximate proportion of different components too closely mixed for clear representation at the scale used (after Trendall, 1976a). (With permission of the Director, Geological Survey of Western Australia).
115 TABLE 3.11 Proportions of lithological types in the Hamersley Group (After Trendall, 1976a) Rock Type
Thickness (m)
Volume percentages total
Iron formation Acid lava Dolerite (sills) Shale Dolomite Tuff
1024 659 386 303 125 49
Total
2546
without:
40.2
4::;
11.9 4.9 1.9
} 41.1
dolerite
dolerite, lava, and tuff
47.4 30.5
70.5
14.0 5.8 2.3
20.9 8.6
In Table 3.11 the proportion of the different lithological components of the Hamersley Group are set out. This emphasizes that it is the abundance of BIF which led t o the initial definition of the group, although the upper part of the underlying Jeerinah Formation has features which indicate its affinity with the Hamersley Group. Disregarding the igneous components of the Hamersley Group it is clear from Fig. 3.4 that there is a broad alternation between shale, or shale and dolomite, and BIF, usually with subordinate shale (Table 3.111). Each of the iron formation units in the sequence is subtly but lithologically distinct from the others (see Trendall, 1976a, table 3). However, because it is the unit which has been most thoroughly studied, only the Dales Gorge Member of the Brockman Iron Formation will be described in detail. The type section of the Dales Gorge Member (Trendall and Blockley, TABLE 3.111 Lithological variations in the upper Fortescue t o Turee Creek Groups (excluding igneous components) (After Trendall, 1976a, 1979) Lithology
Formation
Approximate thickness ( m )
Siltstone, sandstone BIF, shale Dolomite, shale BIF, shale Shale, dolomite
Turee Creek Brockman, Weeli Wolli, Boolgeeda Wittenoom (Carawine), Mt. Sylvia, Mt. McRae Marra Mamba Jeerinah
3 00-4 50 1070 260 180 150-300
116
Fig. 3.5. The eastern side of Wittenoom Gorge showing macrobanding in the Dales Gorge Member of the Brockman Iron Formation (after Trendall and Blockley, 1970). (With permission of the Director, Geological Survey of Western Australia).
1968) has been divided into 33 major BIF or shale units (macrobands), see Figs. 3.4 and 3.5. The following descriptions of the Dales Gorge Member are modified from Trendall’s (1976a) recent review. Details may be found in Trendall and Blockley’s (1970) classic paper (see also Trendall, 1972). The S (shale) macrobands consist mainly of dark green t o black, iron-rich, stilpnomelane-bearing shale, often finely laminated, and of interbanded chert and green siderite, which may be very finely laminated, more or less structureless, or thinly bedded with a ghost clastic structure defined by slight color variations within the fine-grained siderite. Limestone and breccia bands also occur within the S macrobands. Thin bands of stilpnomelane with the shale have textural variations which define the shapes of volcanic shards (La Berge, 1966). The BIF macrobands of the Dales Gorge Member conform in general lithology with “typical” Precambrian banded iron-formation of many continents, but resemble most closely those of the Transvaal Basin of South Africa (Trendall, 1968). Alternate thin bands (called mesobands) of chert and fine-grained iron-rich material were designated chert-matrix by Trendall and Blockley (1970). Chert mesobands are mostly 5-15 mm thick, and in common with other types have by definition a minimum thickness of 1mm
117 the frequency of thicker cherts falls off rapidly up t o the measured maximum of 87 mm. Mesobands of chert-matrix have similar thicknesses, and these two main mesoband types, chert and chert-matrix, make up about 60 and 2096, respectively, of the total BIF volume. The remainder are magnetite, carbonate, stilpnomelane, riebeckite, and minor miscellaneous types. Within most chert mesobands, a small-scale regular lamination defined by layers of some iron-bearing mineral within the general fine mosaic of quartz is present, and is known as microbanding (Fig. 3.6). Although microbands were initially defined (Trendall, 1965), as consisting of simple iron-rich/ iron-poor couplets, subsequent detailed examination shows that a good deal of fine structure may be distinguished within this. Up t o six internal subdivisions of a single microband couplet may be present, some of which may themselves possess evident sub-structure. The defining mineral of the main iron-rich component of each microband is normally hematite, siderite, or ankerite, but may exceptionally be magnetite, stilpnomelane or riebeckite.
Fig. 3.6. A. Microbanding within chert mesobands, separated by chert-matrix, in the lowermost BIF unit in the Dales Gorge Member. B. Cores showing the same upper part of the chert-magnetite group of a Calamina cyclothem, Dales Gorge Member. The drill holes are 6.6 km apart. C. Cores of the same interval in the Dales Gorge Member from drill holes 8 2 km apart. (All photographs from Trendall and Blockley, 1970; reproduced by permission of the Director, Geological Survey of Western Australia)
118 In many BIF macrobands of the Dales Gorge Member there exists a cyclic sequence of mesoband types, over a stratigraphic thickness of 10-15 cm, which Trendall and Blockley (1970) have called the Calamina cyclothem (Fig. 3.6). In this sequence a single thick chert mesoband with fine microbanding, or a group of such cherts separated by thin magnetite mesobands, alternates with a mixed mesoband group of chert-matrix, magnetite, and comparatively thin and coarsely microbanded cherts; these two alternating components of the Calamina cyclothem are called the chert-magnetite group and the mixed group, respectively. Riebeckite occurs within restricted stratigraphic sections of the Dales Gorge Member in some areas, mainly in the form of thick massive mesobands which are stratigraphically equivalent to, and result from the replacement of, the chert-magnetite group of particular Calamina cyclothems. A remarkable feature of the Dales Gorge Member is the lateral continuity of the detailed internal stratigraphy. The 33 macrobands are easily identifiable in the field throughout the outcrop area. Wherever sufficient work has been carried out in well-exposed sections chert mesobands have also proved to be continuous, and microbands within some chert mesobands have been correlated over distances of up t o 300 km. By contrast with the subtle, yet distinctive, differences between the five main iron formation units of the Hamersley Group (Trendall, 1976a, table 3), their chemical composition is relatively uniform, and it is noteworthy that BIFs of quite different lithologies tend to have a uniform total iron content of close to 30%; however, the ratio of ferric to ferrous iron may vary widely, even in unweathered material (see Trendall and Pepper, 1977). Relatively little detailed work has been carried out on stratigraphic units of the Hamersley Group other than the iron formations. The units which contain shale as the main, or an important, constituent include the Mount McRae Shale, the Wittenoom Dolomite, the Mount Sylvia Formation, and the Whaleback and Yandicoogina Shale Members of the Brockman Iron Formation. Although poorly exposed, much of the shale appears to share common characteristics; it is very fine-grained, dark grey or green to black when fresh, and white to buff when weathered, often finely laminated, and usually with a proportion of intermixed chert and carbonate, and sometimes a little iron formation. The upper part of the Mount McRae Shale contains bands of shard-bearing volcanic ash like those in the S macrobands of the Dales Gorge Member, but there is no other direct evidence for volcanic contribution. It appears that there are at least two distinct types of shale: a very fine-grained, soft, poorly laminated black shale rich in free carbon and pyrite, and a dark green, finely laminated shale often with some microbanded chert and with evidence of volcanic activity. The lower part of the Wittenoom Dolomite is composed of massive, medium to thin-bedded dolomite with rare beds of black chert. I t contains sedimentary structures such as cross-beds and slumps, and such features as stylo-
119 lites and chert nodules. Stromatolites appear t o be absent in the formation. The igneous rocks which form over 40% of the Hamersley Group comprise thick dolerite sills largely within the Weeli Wolli Formation, and acid volcanics of the overlying Woongarra Volcanics. Over most of the region over half the thickness of the Weeli Wolli Formation consists of dolerite in several separate sills. The dolerite is uniformly massive, homogeneous, medium- t o coarse-grained, and invariably shows strong deuteric alteration (De Laeter et al., 1974). The acid Woongarra Volcanics include discrete bands of both nonporphyritic and porphyritic rhyolite and dacite as well as thinner horizons of stratified tuff and agglomerate, at least one major intercalation of BIF and a dolerite sill. Although there is apparently a high order of lateral stratigraphic continuity and concordance with the overlying rocks, recent unpublished reexamination has raised some doubt whether the volcanic rocks may be in part intrusive rather than lavas (Trendall, 1976a). Only the lower part of the Hamersley Group is present in the northeastern part of the basin, and is unconformably overlain by the Manganese Group. The basal Marra Mamba Iron Formation thins and becomes more siliceous towards the north and east, being absent or represented by a thin chert in these areas. The Carawine Dolomite (equivalent t o the Wittenoom Dolomite) is a thick bedded dolomite with abundant chert layers (it is rarely ankeritic or calcitic). Shallow-water features are absent except for rare occurrences of cross-bedding, symmetrical ripples, intraclastic conglomerates and stromatolites * indicating local shoaling. The top of the Carawine Dolomite is only observed in the extreme east of the outcrop area where it is interbedded with shales, and appears conformable to and transitional with an overlying sequence of shales and siltstones with rare, thin cherts, “dirty” iron formations and sandstones. The cherts and “dirty” iron formations are pelletal, the first such recognized in the Hamersley Basin. They indicate a shallow-water origin (cf. Hall and Goode, 1978). In view of its stratigraphic position above the Carawine Dolomite, this unit (informally known as the Arrowhead Shale) may be equivalent to the Mt. Sylvia/Mt. McRae Formations in the western Hamersley Basin.
Turee Creek Group The Turee Creek Group (Formation) has only recently been separated from the Wyloo Group (Trendall, 1979). It is conformable with the Hamersley Group, whereas the Wyloo Group is unconformable on all the older sequences. The transition from the Hamersley Group is expressed as a gradual increase in fine clastic debris which dilutes and eventually obliterates
* Stromatolites have been found within t h e Carawine Dolomite in recent mapping. However, stromatolites previously recorded in t h e formation b y D e La H u n t y (1964) have since been shown t o belong to younger units.
120 the characteristic lithology of the iron formation. The unit consists of up to 450 m of thinly bedded shale and litharenite, with local intercalations of quartzite, conglomerate, breccia, dolomite and basalt. Trendall (1976b) has noted faceted and striated sandstone and acid volcanic clasts (Woongarra Volcanics?) of possible glacial origin in a mixtite within the Turee Creek Formation. Wyloo Group The present outcrop of the Wyloo Group is confined to the southern margin of the Hamersley Basin. Metamorphosed equivalents are present in the Gascoyiie Province, further south. N o type sections have yet been estab lished for the Wyloo Group and the accepted nomenclature is the provisional stratigraphy of Macleod et al. (1963), De la Hunty (1965), Daniels (1968, 1970) and Trendall (1979), see Table 3.IV. Recent independent mapping by exploration geologists and by the Western Australian Geological Survey (Trendall, 1976a, 1979) has shown that the Wyloo Group lies with strong unconformity on the underlying rocks whereas it was previously regarded as conformable. For example the entire Hamersley Group has been stripped off around the parts of the Wyloo Dome, and only the lower part of the Fortescue Group is preserved on the basement granites (e.g. Horwitz, 1978). In the northwestern Wyloo Dome, relatively high units of the Wyloo Group (e.g., Duck Creek Dolomite) lie on terrestrial chert breccias above the Fortescue basalts (A.D.T. Goode, S.V. Bell, H.S. Madan, unpublished). The chert breccias are similar to the Pinjian Chert Breccia (basal Manganese
TABLE 3.IV Stratigraphy of the Wyloo Group (After Daniels, 1968, 1970; Trendall, 1979) Formation
Capricorn Formation Ashburton Formation
Duck Creek Dolomite Mt. McGrath Formation - Karlathundra Conglomerate Member - Coolbye Shale Member - Cheela Springs Basalt Member - Nummana Member Beasley River Quartzite
Approximate thickness ( m ) 1oo+ 7200?
300-1000
400 30 1500 30 100-200
Principal lithologies
sandstone, greywacke, dolomite, shale shale, greywacke, minor sandstone, conglomerate, BIF, dolomite, basalt, acid volcanics stromatolitic dolomite conglomerate, sandstone shale amygdaloidal basalt siltstone, quartzite sandstone, minor conglomerate
121 Group) in the northeastern Hamersley Basin. The Pinjian Chert Breccia and minor clastics of the Manganese Group unconformably lie on the Wittenon Dolomite in the Fortescue Valley (M. Kneeshaw, pers. commun., 1978) indicating similar sub-aerial exposure and stripping of upper Hamersley Group sediments along the northern margin of the main basin. The basal Beasley River Quartzite thins from east to west, from about 180 m in the Turee Creek Sheet are2 t o about 90 m in the Mount Bruce Sheet area. This trend continues farther west and northwest, and the formation is not recognizable in the far northwest. It is typically a clean, coarse-grained quartz sandstone. Where stratification is visible cross-bedding and ripple marks have been noted, together with thin bands of breccia, conglomerate, and shale. The clastic component of the lower part of the formation is largely derived by erosion of the underlying Hamersley Group, e.g., BIF, chert, and rarely (but very importantly) high-grade iron ore. Iron-oxide heavy mineral banding occurs locally within rippled and crossbedded siltstones and arenites. The four members of the overlying Mount McGrath Formation can only be distinguished locally; elsewhere the formation rapidly lenses along strike. In the west, lavas are apparently absent, while to the south and southeast, the Cheela Springs Basalt Member is the dominant component. This has a variable thickness up to 1500 m and is a homogeneous dark, fine-grained basalt which is frequently epidotised, and vesicular or amygdaloidal. In these areas the associated sediments locally reach a thickness of 900 m, and include laterally variable shale, sandstone, quartzite, greywacke and con glomerate. The Duck Creek Dolomite, which succeeds the Mount McGrath Formation, is the most consistent unit of the Vyloo Group. It has been mapped over a lateral distance of 320 km. A measured section (Daniels, 1970) includes 990 m of variable dolomite, subdivisible by variations in colour, bedding thickness, or the abundance and nature of stromatolites. The range of stromatolites from the Duck Creek Dolomite is reviewed by Walter (19?2), while Knoll and Barghoorn (1975) describe microfossils from the formation. The overlying Ashburton Formation occupies the major part of the outcrop area of the Wyloo Group. It consists mainly of a monotonous sequence of brown-weathering shale and greywacke with local developments of sandstone, conglomerate or more rarely, banded iron-formation, dolomite, quartzite and basalt. Acidic t o intermediate pyroclastic volcanics are locally important, particularly north of the Wyloo Dome (e.e., Doust, 1975). The virtual absence of any distinctive internal markers, and the usual presence of a well-developed axial plane cleavage obscuring the weakly defined bed ding, makes it virtually impossible t o assess the thickness of this formation. Daniels (1968) estimated 7200 m, mainly of shale, to be present in part of the Turee Creek Sheet.
122
Manganese Group Considerable confusion in stratigraphic terminology exists in the northeastern Hamersley Basin. This is particularly true for that group of rocks known as the Manganese Group, which includes those formations which are unconformably younger than the Hamersley Group sediments and older than the basal units of the Bangemall Group. De la Hunty (1964) in mapping the Balfour Downs Sheet originally referred to part of this sequence as the “Manganese Group” with the following subdivisions (in descending order): (a) Noreena Shale: chocolate coloured blocky shale--siltstone with pellets of braunite. ( b ) Balfour Shale: green shale, micaceous t o sandy, calcareous and cherty. (c) Bee Hill Sandstone: banded cross-bedded reddish sandstone with pebble bands and white spots. (d) Coondoon Conglomerate: conglomerate containing angular boulders and sandy, ferruginous and manganiferous shales. De la Hunty (1963) mistakenly considered the Manganese Group t o be equivalent t o the Waltha Woora Beds of Noldart and Wyatt (1962). The Manganese Group name was also applied by De la Hunty (1969) t o formations to the south, which can be also traced northward t o the younp;er Bangemall Group. Thom et al. (1973) used the name Waltha Woora Formation, which they included in the lower part of the Bangemall Group, in mapping the Ripon Hills and Woodie Woodie areas described by Nolrlart and Wyatt. Recent mapping by the Broken Hill Proprietary Co. Ltd. has shown, however, that the Ripon Hills stratigraphy is classical Manganese Group which can be shown elsewhere t o unconformably underlie the Bangemall Group. Furthermore, rocks similar t o the Waltha Woora Formation can be shown to be the basal unit t o the Bangemall Group in this region. The previous subdivision has tended to lead t o confusion in the past (partly due t o proliferation of names), and does not have a very good stratigraphic basis in that the units seem t o have been deposited in numerous small separated basins. Due t o consequent facies variations from basin to basin, local irregularities in stratigraphic position frequently occur (e.g., in the Ripon Hills, “Balfour Shale” underlies “Bee Hill Sandstone”). Furthermore, the separation of Noreena Shale from Balfour Shale has largely been made on the presence of braunite pellets in the former. These can be shown t o be concretionary, which places serious doubts on the validity of the stratigraphic subdivision of the two units (similar concretionary pellets have for instance also been found in the older Hamersley Group). The recent mapping has shown that the Manganese Group can be divided into three distinct lithofacies:
(1)Pinjian Chert Breccia. This unit is normally separated from the Manganese Group but is best included as its basal unit. Although little geological work
123 has been done on the Pinjian Chert Breccia itself, it is apparent that several different types exist. A true breccia composed of angular chert fragments is a common variety, but “laminated” cherts are also present. Other breccias, apparently transported over greater distances, have a high proportion of sand-size chert grains. No exotic detrital grains have been observed in the unit. The Pinjian Chert Breccia is thickest and most widely developed on the Carawine Dolomite. Its contact with the underlying Carawine Dolomite is very irregular, with “islands” (hums) of dolomite occurring within the Pinjian Chert Breccia throughout the area. It is highly likely that the chert was derived from the underlying dolomite, both directly, by eluvial deposition after solution of the carbonates, and indirectly by solution and reprecipitation of silica as chert, and possibly later brecciation due t o slumping instigated by further solution of underlying dolomite. The palaeotopography is typical of a terrestrial karst terrain.
( 2 ) Clastics, uolcanics. These include the formations initially defined by De la Hunty (1964). These sediments generally lie in well-defined (solution?) depressions in the Pinjian Chert Breccia. Folding in these depressions (e.g., poljes) is generally box-like, with steep sides and flat or undulating floors. The absence of steep dips in the underlying Carawine Dolomite suggests a localized, non-tectonic karst origin for the steep dips on the sides of the “basins” . The sedimentary successions in each “basin” are variable, but general lithologies are similar. These include litharenites with minor chert pebble conglomerates, hematitic sandstones, glauconite sandstones, vitric tuffs, hematitic welded tuffs and rare oolitic hematite rocks. Shales, siltstones and minor tuffs are more common t o the east. (3)Carbonates. These occur only in the eastern part of this region and appear to be stratigraphically equivalent t o (or partly above) the clastic facies. The carbonates (informally known as the Enacheddong Formation) are dominantly shaly dolomites with minor, basal stromatolitic, conglomeratic and sandy dolomites, chert, and hematitic and glauconitic sandstones. Edgewise conglomerates near the base are indicative of shoreline deposits. Manganese Group sediments extend t o the south of the R,at Hill Rise where they unconformably overlie both the Hamersley and Fortescue Groups and the Archaean granite basement of the Sylvania Dome. blanganese-bearing shales and glauconitic sandstones typical of the Group occur t o the south of the dome, but have previously been regarded as part of the Bangemall Group (e.g., Daniels and Macleod, 1965). The exact stratigraphic status of units in this region is under review, but a current interpretation is given in Figs. 3.1 and 3.3. Chert breccias and clastic sediments typical of the Manganese Group are also known in the alluvium-covered Fortescue Valley, and are probably more extensive than shown in Fig. 3.3.
124
Fig, 3.7. Structural map of the Hamersley Basin showing the fold trends of the Ophthalmia Fold Belt in the south (modified from Gee, 1979). Dotted line represents the boundary between the prehnite-pumpellyite (north) and greenschist (south) metamorphic facies zones (after Smith, 1976). Faults shown as heavy lines.
The Manganese Group of the eastern Hamersley Basin is considered the equivalent of the Wyloo Group in the southwest. Both have similar relation ships to the Hamersley Group, and both were deposited after the initiation of folding in the Ophthalmian orogeny, and the beginning of uplift of Archaean granite domes along the southern margin of the early Hamersley Basin. In both cases stripping of earlier sediments is more pronounced on the southern side of the domes. Subaerial exposure of the earlier sequence is indicated in both instances by the presence of terrestrial chert breccias. The relationships between the Manganese/Wyloo and the Hamersley Groups are probably similar to those between the Baraga and Menominee Groups in the Animikie Basin (Cannon and Gair, 1970), and the Pretoria/Postmasburg and the Chuniespoort/Griquatown+ampbell Groups in the Transvaal Basin (Button, 1976a).
Structure and metamorphism The Archaean Pilbara Block represents structurally a very gentle regional
125 arch about an approximate east-west axis. The Fortescue Group dips off it along its southern, western and eastern edge. The gentle southwesterly dip on the southern margin is maintained up t o the axis of a major regional synclinorium - the Hamersley Range Synclinorium (Fig. 3.7) - south of which dips remain low but generally northeast. Within the area covered by this synclinorium, local alterations of dip due t o internally nested folds of continuously smaller scale, are common. Dips are almost entirely less than 20". and mainly less than lo", but abrupt steepening is locally present (Trendall, 1976a). South of the Hamersley Range Synclinorium the whole of the Mount Bruce Supergroup in its central outcrop area is moderately deformed into broad open folds whose limbs have average dips of 30"-50", and their scale is such as to re-expose locally the Archaean floor of the basin; at least some of these basement domes were active during sedimentation and represent a "primary geanticline" (Horwitz and Smith, 1978). The main folds of these central areas maintain a consistent east-west regional axial trend. Along the whole of the southern outcrop boundary the folding becomes still more intense, while maintaining the same regional east- -west trend. At Newman the southern limbs of synclines are commonly overturned, so as t o dip steeply south (Fig. 3.8). The south limb of the Hardey Syncline, in the same narrow zone of intense deformation, has apparently been faulted out. Individual fold structures in the bulk of the Wyloo Group on the southern S
N
High grade hematite ore (av>64% Fe)
Low
- medium
grade ore (54-62%Fe)
-.-
---
L
Major synclinal axis Fault
Jaspilite
I0
Shale Amphibolite & Shale
Scale ~n Metres
0
50
Fig. 3.8. Cross-section of the Mt. Whaleback iron-ore deposit (after Kneeshaw, 1976).
126 margin of the basin are difficult t o recognize due t o the lack of marker horizons, but axial plane cleavage is well developed. This orogenic belt is known as the Ophthalmian Fold Belt. Its development was accompanied by metamorphism increasing up t o amphibolite facies grade towards the southwest (e.g., Smith, 1976: Becker and Clayton, 1976; Horwitz and Smith, 1978), although the possible contribution of contact metamorphism by intrusive granites in these latter areas is not known. A t both the eastern and western ends of the Hamersley Basin the fold axes curve northwards, away from their general east-west direction of the main central area. “The full structural significance of these terminations is not completely understood” (Trendall, 1975b). In the northeastern Hamersley Basin gently folded Fortescue Group basalts are faulted against the Gregory Range granitic complex, at least part of which is Archaean in age (1)c Laeter et al., 1977). The relationship between these structures and the TJatc2 Proterozoic? Paterson Fold Belt is unknown, although there is some indication that the Gregory Range deformation is older. A t the western and southern margins of the Hamersley Basin, a zone of complex faulting separates folded Wyloo Group rocks from folded Hamersley Group rocks t o the north (Fig. 3.3 and 3.7), and it is uncertain t o what extent this zone may have been defined by movements at the close of Hamersley Group deposition. The western region is also associated with a sharp swing in the metamorphic zonation boundaries (Fig. 3.7). A zone of nortlieast-trending faults at the southeastern end of the basin near the bylvania Dome (Fig. 3.3) appears t o control the distribution and sedimentation of younger Middle Proterozoic rocks. This zone corresponds t o the extension of the important Flint Hill Lineament in the Bangemall Basin (Brake1and Muhling, 1976).
Geochronology (1)Basement: the Hamersley sequences lie with angular unconformity on the older (3000 Ma) granites of the Pilbara Block, and are also presumed younger than the c. 2600 Ma tin-bearing granites (Cooke’s Creek Granite 2560 ? 37 Ma, De Laeter et al., 1977; Moolyella Adamellite - 2670 k 95 Ma, De Laeter and Blockley, 1972; Cooglegon Adamellite - 2606 5 128 Ma, De Laeter et al., 1975). Detrital cassiterite, almost certainly from these granites, occurs in the Hardey Sandstone (Hickman and De Laeter, 1977). (2) Fortescue Group: acid volcanics from the Jeerinah Formation have been dated at 2190 ?r 100 Ma (Compston and Arriens, 1968), while the intrusive Gidley Granophyre (De Laeter and Trendall, 1971) and Spinaway Porphyry (Trendall, 1975d) have been dated at 2196 k 26? Ma and 2124 195 Ma respectively. The Black Range and associated dykes in the basement, suggested as possible feeders for the lower Fortescue basalts, have been dated at 2329 2 89 Ma (Lewis et al., 1975). The Hardey Sandstone unconformablv overlies the Cajuput Dyke of this suite (Hickman and De Laeter, 1977).
*
127
( 3 ) Hamersley Group: the Woongarra Volcanics have been dated a t 2000 k 100 Ma (Compston and Arriens, 1968). The age of intrusion of the Weeli Wolli dolerites is suggested as c. 2100 Ma by De Laeter e t al. (1974). (4) Wyloo Group: acid igneous rocks have been dated a t 2020 k 165 Ma by Compston and Arriens (1968). (5) Manganese Group: a K/Ar age of 1050 Ma on qlauconite adjacent t o the younger Davis Dolerite is almost certainly a reflection of the dolerite’s age rather than that of the sediments (BHP, unpublished data). (6) Intrusive granites: the Boolaloo Granodiorite has been dated a t 1720 Ma by Leggo et al. (1965). This age corresponds to various possible updating events in the Hamersley sequence (see the review by Trendall (1976a, fig. 12).
History of sedimentation and tectonism ( 1 ) Initial sedimentation and basin development. Sedimentation in the Hamersley Basin was initiated approximately 2200-2400 Ma ago by the advent o f major flood basalts accompanied by minor fluvial sedimentation. The earliest sediments and volcanics accumulated in palaeovalley systems, preferentially located over Archaean qreenstone belts and separated by broad granite plateaus. However, the sources of the lavas may have been concurrent fissures (now represented as major N--S dykes) located preferentially in the Archaean granite plutons (e.g., Lewis et al., 1975; Horwitz and Smith, 1978). Subsequent “drowning” of the topography by the repeated eruption of basalt flows and pyroclastics was accompanied by regional depression of the central part of the basin (Tom Price area) and associated marine incursion. The central part of the basin contains the thickest development of Fortescue Group strata, and possibly a higher proportion of pillowed (subaqueous) t o massive (subaerial) basalts (Trendall, 1976a). Early marine conditions were shallow water as indicated by abundant sedimentary structures and stromatolites. ( 2 )Deposition o f the Hamersley and Turee Creek Groups. The late Fortescue transgression continued into Hamersley Group time. In t h e major part of the basin sedimentation, while still retaining shelf characteristics, was probably relatively deep water. The almost complete lack of benthonic stromatolites and primary shallow water sedimentary structures, together with the great lateral continuity of microbands in the BIF, indicate that sedimentation took place in quiet water a t depths greater than 100-200 m (i.e. below wave base and below the photic zone). The complete lack of coarse terrigenous sedimentation in the main basin indicates either great distance from a shoreline, low relief of surrounding hinterland, and/or a general aridity of climate. The occurrence of minor shallow-water cherts and iron formations, and minor amounts of coarser terrigenous clastics in the northeastern part of the
128
Fig. 3.9. Isopach pattern of the Dales Gorge Member (modified from Trendall, 1976a). The isopach interval is approximately 6.1 metres. The outcrop area of the member, and locations of the measured sections on which the isopachs are based, are given by Trendall and Blockley (1970). (With permission of the Director, Geological Survey of Western Australia).
basin suggests that this area was shallower and nearer the shoreline of the main basin. The Marra Mamba Iron Formation thins and becomes more siliceous t o the north and east in this area, which by comparison with iron formations in the Nabberu Basin (Hall and Goode, 1978), would be consistent with nearer shore deposition. The shallow-water iron formations of the Nabberu Basin itself may also represent southern shallow shelf equivalents t o the Hamersley Group, but more data on regional correlation are required before equivalence can be established (Hall and Goode, 1978). The area of deepest water sedimentation in the Hamersley Basin corresponds t o the area of greatest thickness of sediments. Isopachs are available for only the Dales Gorge Member of the Brockman Iron Formation (Fig. 3.9). These show maximum thickness along a west-northwest-trending trough in the central part of the Mount Bruce Sheet area, with a steady outward decrease in thickness in all directions, giving a generally elliptical pattern with some irregularities *. Although the regional variation in thickness of the other formations, with the exception of the Woongarra Volcanics, appears to follow a similar pattern, few detailed measurements are available t o support this (Trendall, 1976a). Consequently the characteristics of basin development initiated in Fortescue Group time continued during deposition of the Hamersley Group. The thinning of the Dales Gorge Member over the Wyloo Dome, however, also indicates t h e early structural prominence of this feature (Trendall and Blockley, 1970, p. 69). Similarly thinning of Hamersley sediments near the Sylvania Dome (Kneeshaw, 1975) suggests the early
* N o allowance for structural shortening due to folding was made by Trendall and Blockley (1970), consequently the directional asymmetry of the isopachs may be in part or in total a geometrical illusion.
129 importance of this structure. However, other structural highs such as the Rocklea Dome and Milli Milli Anticline appear t o be entirely of later origin. Ignoring the volcanic and intrusive components of the stratigraphy, the sedimentary sequence of the upper Fortescue, Hamersley and Turee Creek Groups shows a relatively simple and symmetrical variation (Table 3.111). This variation can be interpreted as indicating a continuation of the Fortescue transgression followed by regression leading to the development of the coarser Turee Creek clastics, and eventually culminating in uplift and exposure of basement domes prior t o Wyloo Group time. By this interpretation the deposition of iron formations at the time of maximum transgression is consistent with Hall and Goode’s (1978, p. 178) model for the deposition of the Nabberu iron formations (see later). The occurrence of carbonates (Wittenoom and Carawine Dolomites) may indicate a minor regression within the overall transgression by this model. It is felt that the evidence is more consistent with transgressive-regressive relationships being associated with an interplay between vertical tectonic effects ( downwarping, uplift) and sedimentation rather than solely eustatic changes in sea level. However, the occurrence of possible glacial units in the Turee Creek Group (i.e. associated with the regression) must be considered. Trendall (1976b) has noted possible intercontinental correlatives of these mixtites, thus suggesting the potential for interpreting a world-wide glacial episode. However, the association of the Wyloo mixtites with areas of active basement uplift leaves open the possible association with either local montane glaciation, or even tectonically derived mudflows.
( 3 ) Deposition of the Wyloo and Manganese Groups. Major iron formation deposition in the Hamersley Basin essentially ended with the regression at the end of Turee Creek Group time. The widespread development of chert breccias in the basin indicates extensive sub-aerial exposure, particularly in the shallower parts of the basin at the beginning of both Wyloo and Manganese Group sedimentation. It is believed that the general uplift of the Pilbara Craton in the north and the basement domes in the south was responsible for both the regression at the end of Hamersley-Turee Creek Group time, and the stripping and sub-aerial exposure of the northern and southern Hamersley Group sediments. The lowermost clastic sediments in both the P7yl00 and Manganese Groups may have been terrestrial in origin, with a transgressive marine influence becoming noticeable later. Shallow marine conditions existed during carbonate deposition. It is clear that the areas of greatest Wyloo Group deposition lay well south and west of the region of known maximum Hamersley Group sedimentation. Therefore, it appears that the transition between these groups represents not only an important change of depositional regime, but also a marked change in the configuration of the basin (Trendall, 1976a). Trendall
130 and Blockley (1970) have interpreted the Cheela Springs Basalt Member as an indication of volcanicity restricted to, and tectonically associated with, a hinge line forming the northern edge of this trough, running approximately along the present junction of the Hamersley and Vlyloo Groups. Volcanics are also prevalent, although not previously widely recognized, within the upper Wyloo Group. Gee (1979) has excluded the depositional area of the Wyloo Group (“Ashburton Trough”) from the remainder of the Hamersley Basin, believing it to be a significantly younger feature. However, on analysis of the presently available data it seems t o represent an integral part of the crustal development of this region from Fortescue Group time through t o the final Ophthalmian orogenic phase. The recognition of a zone of active basement highs (“Paraburdoo Hinge Zone”) raises an interesting question. While Trendall and Blockley (1970) regard the major depositional area of FortescueHamersley Group sedimentation t o be confined t o the north of this zone, Horwitz and Smith (1978) regard the hinge as a geanticline with even thicker sedimentation t o the south. The latter view is probably more consistent with the subsequent history of Wyloo Group sedimentation and Ophthalmian folding but requires more work t o substantiate as the proposed older rocks are now largely concealed by the younger Wyloo and Bangemall sediments. The hinge zone broadly corresponds t o a major dipole gravity anomaly between the Fortescue Regional Gravity Complex t o the north and the Ashburton Regional Gravity Ridge t o the south (Wellman, 1978). The configuration of the northeastern Hamersley Basin at this time is not clear. Greater stripping of pre-Manganese Group sediments occurred towards the Pilbara craton, while more marine Manganese Group sediments appear t o be confined t o the easternmost exposed areas. There is no evidence at this time that the Gregory Range granitic basement complex was an active structural high like the Paraburdoo Hinge Zone during sedimentation. ( 4 ) Orogenic finale. The sedimentary history of the Hamersley Basin came t o a close approximately 1700-1800 Ma ago with the development of the Ophthalmian Fold Belt. The asymmetric structure of the fold belt indicates that deformation resulted from northward-directed orogenic forces, centred to the south of the present limits of the basin. The orogeny was accompanied by regional metamorphism up t o at least amphibolite facies grade, and by the intrusion of major granitic plutons in the southwest. It is possible that the Middle Proterozoic Mt. Minnie and Bresnahan Groups represent a molasse facies developed as a result of the orogeny (see later).
(Ire deposits The Hamersley Basin is most notable for its outstanding deposits of highgrade iron ore, which provide the basis of Australia’s huge iron-ore industry
131 (largest exporter and second largest producer in the world after the U.S.S.R., Klinger, 1978). However, smaller but significant deposits of manganese and crocidolite are also known, as well as minor occurrences of gold, base metals, uranium and fluorite. General summaries of mineralization in the Hamersley Basin are given by Macleod (1966), Trendall (1975b) and Blockley (1975a) while Button (1976a) has reviewed the similarities and differences between the Hamersley and Transvaal Basins. ( 1 ) Iron ore. Present official estimates of reserves of high-grade (>55% Fe) iron ore are 18,000 million tonnes of readily usable low-phosphorus ore (<0.08% P) and 17,000 million tonnes of high-phosphorus ore (identified paramarginal resources, see Mead (1975), Pratt (1976). Substantial submarginal resources of hematite mineralization grading less than 55% Fe occur in association with high-grade deposits or as separate deposits, but little information on them is available. Immense submarginal resources of banded iron formation also exist. In the Brockman Iron Formation alone of the Hamersley Basin sediments Lord and Trendall (1976) estimated that 6.4 X 10’‘ tonnes of material grading greater than 30% contained iron is available for extraction by open-cut methods. To estimate the total quantities available in the Hamersley Basin it would also be necessary to add estimates for the Marra Mamba and Boolgeeda Iron Formations and the additional quantities available if underground mining was considered. 1977 production from the four operating Hamersley mines at Newman, Tom Price, Paraburdoo and Pannawonica was 76 million tonnes (82% of Australia’s total production). The largest mines and deposits occur as hematite-goethite enrichments within the Hamersley Group iron formations, although a large proportion of the resources of the region occur in Tertiary limonite deposits (Robe Pisolite) ultimately derived from the iron formations (see Macleod, 1966; Harms and Morgan, 1964). General descriptions of the ore deposits may be found in Macleod (1966) and Trendall ( 1 9 7 5 ~ ) . Despite their economic importance, the Hamersley hematite-goethite deposits are poorly understood, particularly with regard to the mechanism and timing of enrichment. Trendall ( 1 9 7 5 ~ )has estimated that hematite bodies constitute less than 2% of the entire present outcrop area of BIF. Larger ore deposits include Mt. Whaleback at Newman (1400 million tonnes, 64% Fe - Kneeshaw, 1976) and Mt. Tom Price (610 million tonnes, 63% Fe - Gilhome, 1975). Other major deposits include Paraburdoo and Koodaiderie (Baldwin, 1975), Area “C” (Neale, 1975), Jimblebar or McCamey’s Monster (Ward et al., 1975), Brockman deposits (Evans and Clint, 1975), Rhodes Ridge, Marillana and Marandoo. The ore deposits consist essentially of hematite with minor quartz and magnetite (Ayres, 1971). Goethite tends t o be more common near the surface, suggesting it is a Tertiary weathering feature of the ore. It is also more common in the newly discovered deposits in the Marra Mam-
132 ba Iron Formation (e.g., Kneeshaw, 1976). The distribution of phosphorus in the ores is discussed by Graham (1973) and Morris (1973). There appear t o be two major controls on the occurrence of ore bodies: (a) Stratigraphic - enrichment is greatest in the Dales Gorge Member and to a lesser extent the Joffre Member, although at Mt. Whaleback high-grade ore also occurs in most other members of the Brockman Iron Formation and also in the overlying Mt. McRae Shale (Kneeshaw, 1976). The Marra Mamba Iron Formation has recently been found t o also contain significant deposits such as Area “C” (Neale, 1975), Orebody 26 at Newman (Kneeshaw, 1976) and Marandoo. The rapid lateral transition from host BIF t o ore is normally accompanied by a reduction in stratigraphic thickness of about 50%. The ore forming process seems to have involved leaching of most “impurities” (including rare earths and phosphorus) from the iron formations leaving the residual iron oxide textures (magnetite is converted to hematite), although it appears iron has been added in some situations. (b) Structure - the larger ore bodies of higher grade are mostly associated with faulted synclines (Fig. 3.8), e.g. Mt. Whaleback (Kneeshaw, 1975), Tom Price (Gilhome, 1975). The occurrence of ore bodies has previously also been related to the Tertiary “Hamersley Surface”, but drilling has indicated hematite enrichment to at least 400 m casting some doubt on this proposition. The temperature of the ore-forming fluids at Mt. Whaleback has been estimated by Linley (1975) t o be less than 2OO0C,on the basis of isotopic fractionations of quartz-hematite pairs. A trend surface analysis of Fe, P, S O z , Al,03 and L.O.I. values undertaken by Linley (1975) suggested that the direction of flow of the ore-forming fluids was downward and northerly within the BIF units. This flow was apparently blocked by the relatively impermeable shales and amphibolites of the Jeerinah Formation and consequently channelled to the west. The flow of the fluids was thus concentrated, resulting in a hematite orebody of unique purity and size (Kneeshaw, 1976). The formation of the hematite ore may date from the time of major faulting. There is palaeomagnetic evidence to suggest the ore-forming processes stopped some 1500-1600 Ma ago (Porath, 1967; Linley, 1975). The surface goethitic ores were formed in relatively recent times (Tertiary?) by supergene enrichment by meteoric waters. The occurrence of rare enriched hematite pebbles in the basal Wyloo Group would tend t o suggest that enrichment began prior t o c. 1800-2000 Ma. However, as no ore bodies show post-enrichment deformation (Trendall, 1976a) most enrichment presumably took place after the development of the Ophthalmian Fold Belt. Similar ages of enrichment have been suggested in the Middleback region of South Australia (Furber, 1976; Chamalaun and Dempsey, 1978). Hematite pebbles derived from Early Proterozoic Nabberu iron formations in Western Australia occur within Bangemall Group rocks (Bunting et al., 1978), also indicating enrichment prior to 1100 Ma ago.
133 ( 2 ) Manganese. The Pilbara manganese province lies in the northeastern part of the Hamersley Basin (De la Hunty, 1963; Blockley, 1975b). Production to the end of 1972 was about 1.5 million tonnes averaging 49.1% Mn. Substantial reserves of ferromanganese ore have also been outlined. Deposits are of several types and are similar t o those in the Postmasburg region of South Africa (Anhauesser and Button, 1976; Button, 1976a). (a) Replacement of Pinjian Chert Breccia: these deposits are common and small but relatively important, e.g. 30,000 tons of ore have been mined from Skull Springs. They are characterized by their poddy, irregular nature, and form by replacement of the Pinjian Chert Breccia which hosts them. (b) Replacement of basal shales of the Manganese Group: this is also a common type of deposit (e.g., Ripon Hills, Balfour Downs), and is probably the most important, e.g. Ripon Hills contains 20 million tonnes of proved ore at 19%Mn, 26% Fe with 50 million tonnes of indicated and inferred ore. Braunite, bixbyite, concretionary pyrolusite, and cryptomelane have been identified in the ores. (c) Replacement of Carawine Dolomite: manganese oxides also replace Carawine Dolomite or infill caves and joints in the dolomite in many areas. These deposits are generally small. Larger deposits, sometimes very important, occur in the Carawine Dolomite adjacent to large faults, e.g., Woodie Woodie (reserves in excess of 100,000 tons), Mt. Sydney. Braunite, hausmannite, psilomelane, pyrolusite and cryptomelane have been recorded from these deposits (De la Hunty, 1963). As in South Africa, the development of these deposits is probably related to the Proterozoic karst topography in early Manganese Group time. Manganese enrichments are also found associated with the Marra Mamba Iron Formation in vqrious parts of the Hamersley Basin. ( 3 ) Crocidolite. Cross-fibre riebeckite (crocidolite) and the more abundant but noncommercial mass-fibre (felted) riebeckite are confined mainly t o 30 to 60 m thick stratigraphic intervals (riebeckite zones) in the Marra Mamba Iron Formation, and Dales Gorge Member of the Brockman Iron Formation (Trendall and Blockley, 1970; Blockley, 1976). Mineable seams of crocidolite appear in more restricted horizons (5 cm t o 4 m thick), which are generally within the riebeckite zones although one, the Yandicoogina Horizon, is isolated from other occurrences of riebeckite. The seams comprise one or more bands of crocidolite lying parallel to the bedding planes of the host banded iron formation, with the fibres perpendicular or at a high angle to the bedding. The Colonial and Wittenoom mines yielded about 150,000 tonnes of fibre of 5-8 cm average length prior t o their closure. The deposits are similar t o those in South Africa (Trendall and Blockley, 1970; Button, 1976a), although they are shorter in fibre length.
( 4 ) Others. Only minor quantities of lead, zinc, silver, gold and copper are known fiom quartz veins, commonly within the Fortescue and Wyloo
134 Groups (Blockley, 1971, 1975a). Stratiform copper mineralization in graphitic, pyritic shales has been recorded in minor quantities in the Jeerinah Formation (Blockley, 1975a; see also Smith, 1976). Gold and uranium occur in pyritic conglomerates near the base of the Fortescue Group near Nullagine and Marble Bar (Richards, 1972; Blockley, 1975a). The deposits resemble the Au/U bearing pyritic quartz-pebble conglomerates of the Witwatersrand (South Africa) and Huronian (Canada) Basins. Uranium also occurs as small, rich shoots on fault zones in metamorphosed Wyloo Group sediments in the southwest of the basin (Blockley, 1975a). Significant fluorite vein deposits within Fortescue Group basalts occur east of Marble Bar in the Meentheena area (Hickman, 1974).
Early Proterozoic Nabberu Basin The Nabberu Basin lies along the northern margin of the Archaean Yilgarn Block (Hall and Goode, 1975, 1978; Horwitz, 1975a, b), extending at least 700 km in an east-west direction and covering an area of approximately 60,000 km2 (Figs. 3.1 and 3.10). The basin was recognized in 1973 following publication of regional aeromagnetic maps of the eastern part of the region by the Bureau of Mineral Resources (Hall and Goode, 1975). Conspicuous east-west trending magnetic anomalies were interpreted as indicating the presence of banded iron formations, and subsequent field investigations revealed the presence of a large sedimentary basin containing iron formations of the Lake Superior type, the first of their kind discovered in Australia. Further field work has extended the known basin t o the east and west. West of the Robinson Ranges much of the Early Proterozoic cover has been eroded. The basin comprises an eastern Earaheedy Sub-basin, a central Glengarry Sub-basin, and a western Padbury Sub-basin (Hall et al., 1977). Probable tectonic control of sedimentation has led to the development of distinctive facies assemblages of differing thicknesses in each sub-basin. Deformation of the Nabberu Basin in the late Early Proterozoic(?) resulted in the formation of the northern Stanley Fold Belt, in which the sediments are strongly deformed and metamorphosed. To the south the Kingston Platform is virtually undeformed (Fig. 3.11). The boundary between these two tectonic subdivisions extends approximately east-west to west-southwest across the central part of the Earaheedy Sub-basin. The Padbury Sub-basin is entirely within the Stanley Fold Belt.
Earaheedy Sub-basin The Earaheedy Group, comprising nearly 6000 m of shallow water sediments was deposited within the Earaheedy Sub-basin (Fig. 3.12). In the west
w 0)
Y
2 ‘D
0
0
0
0 Ln
0
BANGEMALL
I
+++
+ I + -
Fig. 3.11. Structural map of the Nabberu Basin. The dotted line represents the boundary between “unmetamorphosed” (south) and lower greenschist metamorphic facies (north) sediments, and also the southernmost extent of penetrative slaty cleavage (after Hall and Goode, 1978). This line also delineates the division between the Stanley Fold Belt and the Kingston Platform.
I
I
137
Fig. 3.12. Landsat photograph of the eastern Nabberu Basin (Earaheedy Sub-basin). Note the prominent development of banded iron formations (Frere Formation) outlining the Earaheedy Synclinorium and the Lake Teague Dome (lower left of centre).
and north of the sub-basin slightly deeper water-facies rocks appear, some of the units recognized t o the southeast die out, and thicknesses of the lower units in particular show marked changes. The division of the Earaheedy Group into a lower Tooloo Sub-group and an upper Miningarra Sub-group is based on its deposition in two distinct cycles of sedimentation (Table 3.V). The stratigraphic units were defined by Hall et al. (1977) and type sections were described for each formation.
(1) Tooloo Sub-group. The basal Yelma Formation in the southeast is composed of medium- t o very coarse-grained clean (supermature) quartz sandstones, sometimes arkosic, with minor bands of quartz pebble or quartz cobble conglomerates. Siltstones become increasingly abundant higher in the formation and to the north and west. Rounded glauconite pellets are present in the Yelma Formation in the southeast (Bunting and Chin, 1975; Bunting et al., 1975), but have not as yet been observed further west. Probable equivalents of the Yelma Formation in the deformed northern portion of the Earaheedy Sub-basin include a lower medium-grained, impure quartzite unit, with scattered rounded quartz pebbles and thin grit bands, now psammitic quartz-mica schists, and an upper phyllite and slate unit con-
138 TABLE 3.V Stratigraphy of the Earaheedy Group (After Hall and Goode, 1978) Formation
Li thologies
Comments
quartz sandstone, minor shale. limestone
cross-bedding, slumps, mud pellets, minor load casts, glauconite
300
limestone, minor sandstone, shale
stromatolites, oolites, intraformational conglomerates
2000
fine micaceous sandstone, minor siltstone, limestone
festoon cross-bedding, ripple marks, slump rolls, scour channels, mud pellets, rare mud cracks
Maximum thickness; (m)
Miningarra Su.b-Group: Mulgarra Sandstone 100
Kulele Creek Limestone Wongawol Formation
Princess Ranges Quartzite
200
quartz sandstone, siltstone, minor grit
cross-bedding, glauconite, mud pellets
Wandiwarra Forma tion
350
fine micaceous sandstone, shale
cross-bedding, minor glauconite, ripple marks, mud pellets
Tooloo Sub-Group:
Windidda Formation
1200
limestone (ankeritic, chamositic a t base), shale, minor sandstone, chert
stromatolites, intraformational conglomerates
Frere Formation
1300
granular iron formation, chert, siltstone, minor carbonates, laminated iron formation. Dolomite at base
local stromatolites, oolites, gypsum pseudomorphs, cross-bedding
sandstone, siltstone, minor conglomerate, iron formation, carbonate
cross-bedding, ripple marks, glauconite
Yelma Formation
130+
taining thin bands of schistose quartzite and conglomerate. These rocks are equivalent to the Malmac Formation of Horwitz (1976) and the Troy Creek Beds of Bunting et al. (1978), but their stratigraphy and thickness have not been determined with certainty. The Frere Formation consists of interbedded ferruginous chemical sediments (iron formations) and fine-grained clastics and minor carbonates. It lies between the Yelma Formation and the Windidda Formation east of Lake
139 Teague, but to the west it is conformably overlain by the Wandiwarra Formation, lowest unit of the Miningarra Sub-group. The formation thickness of probably less than 1300 m appears to remain fairly constant along the southern margin of the Earaheedy Sub-basin, but within this interval an iron formation/shale ratio of about 60 : 40 in the west changes to a ratio of much less than 1 0 : 90 in the southeast as individual bands of iron formation thin from a maximum of 200 m thick in the west to generally less than a metre in the southeast. The iron formations also tend to become more siliceous eastwards, and the top of the Frere Formation is marked in some areas by a distinctive c h e r t s h a l e unit (Hall et al., 1977). The iron formations have been described in detail by Goode and Hall (1976) and Hall and Goode (1978). The most common iron formation has a distinctive intraclastic (pelletal) appearance, comprising rounded t o rarely angular intraclasts of ferruginous chert (hematite-quartz magnetite) and/ or chert up to 2 mm in diameter, in a chert or chalcedonic matrix (Fig. 3.13a, b). Detrital phases such as muscovite and tourmaline are occasionally present. The pellets are generally sub-spherical in shape, although in some units they may be irregularly tabular and curved. Shrinkage or syneresis cracks filled with clear, nondusty quartz are common in individual intraclasts. The intraclastic iron formations are usually internally unbanded and lenticular. Bands of microbanded chert are commonly disrupted or fragmented, especially in the upper part of the Frere Formation, and large, distinctive tabular chert clasts up t o 4 cm long are incorporated in the pelletal units in which shallow-water sedimentary structures are well developed. These features are interpreted as indicating later or contemporaneous shallow-water reworking of iron-rich chemical precipitates, and are very similar to textures and structures in the granular iron formations of the Lake Superior region and Labrador Trough of North America (Gross, 1972; Mengel, 1973; Dimroth and Chauvel, 1973). Although such features may not necessarily form in high-energy environments (see Adeleye, 1975), it is considered that the bulk of evidence indicates that most of the Frere iron formations were reworked during storm, wave and tidal action of varying intensity. Thin units of oolitic iron formations occur within the pelletal ferruginous cherts, especially in the southeast of the Earaheedy Sub-basin. The ooliths are usually elliptical in shape, and rarely contain multiple nuclei of ferruginous chert or chert (Fig. 3.13~).In the same area abundant algal filaments and microfossils occur in a thin oncolitic and stromatolitic iron formation interbedded with pelletal and oolitic ferruginous cherts (Walter et al., 1976). The micro-fossils show strong similarities t o benthonic micro-organisms from the Early Proterozoic Gunflint and Biwabik Iron Formations of the Lake Superior region, although their preservation is not as good (cf. Barghoom and Tyler, 1965). Large syneresis cracks in the oncolite cores are filled with small rounded intraclasts indicating gel dewatering and crystallization at an early stage prior to final deposition.
*
140
Fig. 3.13. Textures in the Frere iron formations, Nabberu Basin (after Hall and Goode, 1978). a. Rounded to subangular intraclasts (pellets) of fenuginous chert matrix. b. Angular intraclasts of ferruginous chert in chert matrix. c. Ferruginous chert ooliths, including rare multiple ooliths, in chert matrix. d. Laminated ferruginous chert, Illagie Iron-Formation Member.
141 In the western part of the Earaheedy Sub-basin a 200 m thick banded iron formation, the nlagie Iron-Formation Member, is developed at the top of the Frere Formation. This is a microbanded hematite-rich ironstone with minor chert (Fig. 3.13d), and is believed t o be a deeper water, non-reworked equivalent of the pelletal ferruginous chert (i.e. an Algoman or Hamersley-style BIF). This member thins t o the east, and is absent in the eastern Frere Range. The clastic units separating individual bands of iron formation in the Earaheedy Sub-basin are generally siltstones, shales and sandy shales up to 100111thick (e.g., Lockeridge Shale, Wy-Uda Shale, and Murgurra Shale Members). The siltstones are finely laminated and often have small-scale cross-bedding, and frequently include bands of chert less than 3 m thick. Climbing ripples are observed in some siltstone units. In some units, unusual thin, angular “broken glass” slivers of quartz are common. The origin of these grains is unknown, but Hall and Goode (1978) have suggested either a tuffaceous or reworked Archaean silcrete derivation. Carbonate units occur both near the base and top of the iron formations. Dolomites appear t o be more common than limestones, and contain stromatolites of various types which have been described by Hall and Goode (1978) and Preiss (1976). Recently, gypsum pseudomorphs indicating an evaporitic environment (sabkha) have been discovered in the basal dolomites (W.D.M. Hall and P.G. Haslett, pers. commun., 1979). Silicified vadose pisoliths in the upper part of the formation also indicate local sub-aerial exposure of the unit in the east (Hall and Goode, 1978). The Windidda Formation, consisting of carbonates and fine-grained clastic sediments, is confined to the southeast of the basin, although t o the west and north zones of no outcrop may correspond t o its stratigraphic position. In the type area the basal few metres of the Windidda Formation, transitional to the Frere Formation, consist of brown ankeritic carbonates with interbedded, pelletal ferruginous cherts, pelletal chlorites, black sulphidic cherts and stromatolitic horizons. Sedimentary chlorite with a chamositic composition occurs as very fine-grained material on bedding planes and algal laminations, and as reworked pellets within carbonates and cherts. Rare detrital chlorite is also present, and some laminated chamosite has been found finely intergrown with chert. Chamosite and also glauconite, which is found elsewhere in the Nabberu Basin, are not commonly observed in Precambrian rocks (James, 1966). Apatite, ankerite, chalcopyrite and sphalerite also occur in some of the thin, dark chert bands. The basal units in the type area are overlain by locally conglomeratic, laminated limestones containing abundant stromatolitic beds, and interbedded with calcareous shales which become more common higher in the sequence. Beds of sandstone and intraformational limestone conglomerate also become more common higher in the sequence. The top of the Windidda Formation is marked by a prominent band of intraformational limestone
142 conglomerate containing large (5 cm) flat clasts of laminated limestone in a quartz-glauconite sand matrix. Irregularly layered calcite of variable grain size and with both normal and reverse size grading occurs in the conglomerate. Microtextural studies show that this calcite has a polygonal fitted texture with some polygons preserving a fibrous radiating morphology with rare concentric banding. Large, wavy concentric pisoliths are also found in the overlying Wandiwarra Formation, and are very similar t o those described by Dunham (1969) from pisolitic carbonates of vadose origin in the Capitan Reef. This implies sub-aerial exposure of the uppermost P'ind-idda carbonate sediments in a dry climate. The unit is intruded by Middle Proterozoic dolerite sills in the eastern part of the sub-basin (Preiss et al., 1975).
(2)Miningarra Sub-group. The base of the Wandiwarra Formation is marked by a thin, but prominent, medium-grained, ferruginous, mature quartzose glauconitic sandstone, which is overlain by a similar ferruginous glauconitic sandstone containing flat shale pebbles and distinctive calcite pisoliths of possible vadose origin. These pisoliths exhibit a radial structure and an unusual wavy concentric banding unlike that observed in normal sedimentary pisoliths. In the type area the Wandiwarra Formation appears t o be conformably overlain by the Princess Ranges Quartzite representing the incoming of extensive mature quartz sand deposition (beach or barrier sand environment?), but appears t o be absent in the eastern part of the sub-basin. The glauconite in the Wandiwarra Formation is very aluminous, and has a skolitic composition. The aluminous nature of the glauconite may indicate that it was reworked from a non-marine environment (Hall and Goode, 1978), and that it formed during the period of sub-aerial exposure following deposition of the Tooloo Sub-group. The Wongawol Formation consists of a rather monotonous sequence of finely laminated, very fine-grained micaceous sandstone and minor siltstone units about 2000 m thick. Strongly developed sedimentary features include small-scale festoon cross-bedding, ripple marks, small scour channels and rounded mud pellets up to 2 cm across, and are indicative of shallow-water sedimentation. Rare mud cracks indicate local sub-aerial exposure of the formation. A few thin bands of shale and small pebbles also occur in the sandstone. Small slump rolls are particularly abundant in the upper part of the formation. In the upper part of the formation (Sholl Creek Member), the sandstones are interbedded with bands up t o two metres thick of micaceous shale, fine calcareous sandstone, oolitic limestone, oolitic limestone conglomerates, and occasional one metre thick stromatolitic limestones. The stromatolites are domal, similar t o those in the overlying Kulele Creek Limestone. The carbonates usually contain minor quantities of coarse terrigenous silt, grading up rarely to a medium sand fraction. The limestones and associated silt are
143 usually mature and well-winnowed sediments with little, if any, micritic matrix, and are typically shallow-water sediments. The ubiquitous presence of ooliths indicates strong tidalcurrent action. Carbonate intraclasts in the conglomerates are frequently micritic, and indicate quieter depositional environments in the vicinity of the observed conglomerates. The Kulele Creek Limestone comprises repeated (cyclic) sequences several metres thick. At the base of the formation the cycles consist of purple shales and carbonate conglomerates which give way to massive t o crossbedded, oolitic limestones, which are in turn overlain by stromatolitic limestones with interdomal conglomeratic and pisolitic carbonate. At the top of the formation shales are absent, although minor carbonate conglomerates are present. This upper part of the formation is dominated by stromatolitic limestones with a pisolitic limestone matrix. The cycles have been interpreted by Hall and Goode (1978) as indicating a general deepening of water within and between cycles, i.e. a rhythmical deepening and progressively transgressive sequence from supratidal mud-flats to shallow sub-tidal reefal conditions. The stromatolites are large domal structures, up t o 1-2 m across and a metre high, indicative of shallow sub-tidal conditions (Serebryakov and Semikhatov, 1974). The trapping of sand-sized material as well as the association of pisoliths, ooliths and coarse carbonate clasts in interdomal hollows probably indicates strong current activity in the vicinity of the stromat olite colonies. The Mulgarra Sandstone is the youngest known unit of the Earaheedy Group, and is exposed only in the extreme east of the Earaheedy Sub-basin. It possibly represents the beginning of a third transgressive phase of sedimentation, or alternatively represents a regression at the end of the second sedimentary cycle.
Glengarry Sub-basin The Glengarry Sub-basin in the central part of the Nabberu Basin contains a much thicker and generally finer-grained sequence than the Earaheedy Sub-basin. The Glengarry Group contains a basal sequence of mature, crossbedded quartz sandstone and minor conglomerate to shale-siltstone with sand lenses (Finlayson Sandstone of Bunting et al., 1977). The sandstones are thickest along the southern margin of the Sub-basin, and probably represent strandline deposits. This unit is conformably overlain by siltstone and shale with minor sandstone, conglomerate, chert and laminated, stromatolitic dolomites (Maraloou Formation of Bunting et al., 1977) in the south, and interbedded slate and psammitic schist (turbidites) in the north. The turbidites are volcanoclastic in the east and quartzose in the west (W.n.M. Hall, pers. commun., 1979). A feature of the Glengarry Sub-basin is the occurrence of large piles of metabasaltic rocks or thick dolerite sills interbedded with the lower sedimentary units. A volcanic origin for the basalts is supported bv the occur-
144 rence of abundant volcanoclastic greywackes (Thaduna Beds) and basic tuffs. The greywackes are locally cross-bedded, poorly sorted and immature. They are composed of angular igneous rock and mineral fragments, suggesting a local source, and well-rounded quartz grains, indicating an admixed, second more mature provenance (Hall and Goode, 1978). The fluviatile Kaluweerie Conglomerate (Allchurch and Bunting, 1976) and other outliers up to 120 km south of the Glengarry Sub-basin may be equivalent to the basal sandstones of the Earaheedy and Glengarry Subbasins. Abundant detrital? pyrite occurs in heavy mineral bands in feldspathic litharenites in one of these outliers at Mt. Yagahong (W.D.N. Hall and A.D.T. Goode, unpublished). Pad b ury Su b-basin The Padbury Sub-basin is broadly separated from the Glengarry Subbasin by the Archaean Goodin Dome for convenience but the sequences are transitional. Bunting et al. (1977) regard the Goodin Dome as being active during sedimentation although this has yet to be confirmed. The Padbury Group has been described by F4acleod (1970), Barnett (1975) and Mall and Goode (1978). The Peak Hill Beds are composed of fine- to coarse-grained clastic sediments, now quartz-mica and chlorite schists and quartzites with occasional conglomeratic beds and rare fuchsite schists. Thin beds of iron formations are also reported by Macleod (1970) who suggests a minimum thickness of 1700 m for the Peak Hill Beds. The Peak Hill Beds also contain local accumulations of thick basaltic metavolcanics similar to those in the Glengarry Sub-basin. The Peak Hill Beds are overlain by the Horseshoe Range Beds, Labouchere Formation and Robinson Range Formation (Macleod, 1970; Barnett, 1975). These units total about 3000 m thick, and consist dominantly of mediumto fine-grained clastics including flysch-like units, with rare pelletal iron formation, rare thin dolomite, and banded iron formation at the top (Robinson Range Formation) and bottom (Horseshoe Range Beds) of the succession. Metamorphosed, often talcose, dolomites are found above thin iron formations in the western part of the Padbury Sub-basin. The occurrence of large blocks of ferruginous chert (“banded iron formation”) within the Peak Hill Beds has been suggested by Horwitz and Smith (1978) t o represent olistostromes and interpreted as indicating slumping of the Frere Formation or its equivalents into the Peak Hill Beds. However, such features may simply represent chert boudinage structures within the slate sequence (W.D.M. Hall, pers. commun., 1978). The Millidie Creek Formation includes slate, chert, quartz schist and conglomerate, and is the youngest formation preserved in the J3obinson RangePeak Hill area (Barnett, 1975).
145
Relationships between the sub-basins This topic has been the subject of the most controversy in the discussion of the Nabberu Basin (see Bunting et al., 1977). All workers seem to agree on the general equivalence of the basal Padbury Group (Peak Hill Beds) with the Glengarry Group. However, the relationships between the eastern Earaheedy Group and the western sediments comes most discussion. Horwitz (1975a) and Horwitz and Smith (1978) believe the Glengarry Group to unconformably overlie the Earaheedy Group, although Horwitz and Smith (1978) record “olistrostromes” of Frere iron formations within the Glengarry Sub-basin, suggesting contemporaneity (however, see discussion above). Bunting et al. (1977), while conceding the equivalence of the Padbury and Earaheedy iron formations, regard the Earaheedy Group as unconformably overlying the Glengarry Group (i.e. the basal Padbury Group). However, it should be noted that no unconformity has been found between sediments of the Earaheedy and Glengarry Groups, and that almost identical basal sediments of each group rest unconformably on Archaean basement. Hall and Goode (1975, 1978), on the other hand, maintain that all the available lithological, structural and facies data are entirely consistent with the broad equivalence of all the groups, and that the variations are those that would be expected with simple transgressive-regressive relationships and with structural control on sedimentation (Fig. 3.14). That structural control was important in the development of the Glengarry Sub-basin in particular is shown by Hall and Goode (1978) and by Bunting et al. (1977). Hall and Goode suggest that the iron formations of the Padbury and Earaheedy Sub-basins are equivalent, and that the thin clastics of the Yelma Formation in the southeast are facies equivalents of the much thicker, but generally finer-grained, clastics of the northern Earaheedy Sub-basin, Glengarry Sub-basin and basal Padbury Sub-basin. The basal Padbury clastics also thin markedly to the west. The lack of iron formations in the Glengarry Subbasin is largely due to structural causes (i.e. the easterly plunge of the Earaheedy Synclinorium) as facies variations within the Raraheedy Sub-basin indicate that thicker, deeper-water iron formations formerly existed towards the west (now preserved further west as the Padbury iron formations). Structure and metamorphism The Earaheedy Sub-basin is deformed into a large synclinorium which opens out to the southeast, and progressively younger stratigraphic units are preserved as the present centre of the structure, immediately north of Lake Carnegie, is approached (Fig. 3.11). In this region the entire stratigraphic sequence, with the exception of the two uppermost units, can be studied in both a deformed and undeformed state, and the progressive development of structural features provides an excellent model to which deformation in the remainder of the basin can be related (Hall and Goode, 1978). In the south on the Kingston Platform, sub-horizontal t o gently dipping
146 sediments rest unconformably on Archaean granitic rocks and narrow, north northwest-trending belts of metasediments and metavolcanics. The gentle northerly dips persist almost to the centre of the synclinorium. In the centre of the sub-basin these are followed by the appearance of saucer-shaped folds with gentle dips of up to 20". Northwards the folds become ovalshaped, and limb dips increase quite rapidly up to 50" as the folds become more elongate and canoe-shaped. Cleavage, dipping 40-60" northwards, becomes noticeable in the finer-grained rocks, and small north-dipping thrusts appear. Further north the cleavage becomes increasingly more penetrative and the folds tighter and slightly overturned southwards (see Hall and Goode, 1978, fig. 4). In the north of the fold belt the resistant iron formation units form elongate, tight to isoclinal, doubly plunging (i.e. closed) structures which are commonly asymmetric and overturned southwards, and often dislocated by northdipping thrusts. Locally within this part of the fold belt the cleavage and the individual folds to which it is related have been refolded into structures which may be slightly overturned southwards. Archaean basement rocks reappear in the north of the fold belt in the central Stanley Sheet (e.g., Malmac Dome of Horwitz, 1975b), and have a moderately developed planar fabric imparting a slightly gneissic appearance. To the west, basement rocks become increasingly involved in the deformation, and increasingly gneissic. Sheared granite appears in anticlinal crests (e.g., Goodin and Marymia Domes of Bunting et al., 1977), and eventually occupies extensive anticlinorial areas on either side of the major Mt. Seabrook synclinorium of schistose Early Proterozoic rocks. This deep, tight synclinorium closes t o the west, and only isolated patches of strongly deformed, intensely metamorphosed Early Proterozoic rocks are preserved within amphibolite and granulite facies gneisses further west. In the western parts of the fold belt, the schistosity of the Early Proterozoic rocks is strongly refolded along generally east-west trends. A second axial plane fabric locally transposes the first, and numerous steeply plunging second folds occur, as well as large-scale fold interference patterns. In these areas the basal unconformity becomes increasingly blurred, and is marked by the contact between granitic gneisses and coarse-grained, locally conglomeratic psammatic schist derived from arkose and impure quartzite. This orogenic belt is known as the Stanley Fold Belt (Hall and Goode, 1978); its formation is related to two distinct pulses of deformation in the east and three in the west. The intensity of each deformation phase increases westwards. One of the most notable features of the Stanley Fold Belt is the abrupt change in trend in the northwest of the Earaheedy Sub-basin. PThile changes in strike of the major structures at the western end of the fold belt can be related to later deformations, no such cause can be attributed to this change as no significant second phase structures are evident. Hall and Goode
Fig. 3.14.Palinspastic diagram showing the interpreted longitudinal section of the Nabberu Supergroup (after Hall and Goode, 1978).
148 (1978) have suggested that important basement structures are the controlling influence on this change in trend. Regional changes in metamorphic grade accompany the variations in structural deformation in the Nabberu Basin and its Archaean basement (Hall and Goode, 1978; see also Horwitz and Smith, 1978). The assemblages within various rock suites indicate a general increase in metamorphic grade from essentially unmetamorphosed sediments on the Kingston Shelf (even lower grade than in the northern Hamersley Basin) through lower and upper greenschist facies to possibly a maximum grade of granulite and amphibolite facies in the strongly deformed western parts of the Stanley Fold Belt (Fig. 3.11). However, it is uncertain whether the highest grade assemblages are related to the development of the Stanley Fold Belt, or represent an earlier event within the basement only. The spectacular Lake Teague dome and associated rim syncline (Figs. 3.11 and 3.12) on the southern margin of the Earaheedy Sub-basin are probably related to a piercement structure. Butler (1974) suggested that the structure is an astrobleme, but unusual alkaline rocks within the core suggest an intrusional origin (A.D.T. Goode and C.D. Arndt, unpublished). Bunting et al. (1977) suggested the structure originated by cold emplacement of a syenite plug at high strain rates, while Horwitz (197513) proposed a fold interference mechanism. Geochronology The sediments from the eastern Nabberu Basin have previously been considered to be equivalent to the lower division of the Bangemall Group (Daniels, 1966b) and thus of Middle Proterozoic age, while those from the central, more deformed and metamorphosed part of the basin were mapped as Archaean by Macleod (1970). However, Macleod noted the possibility of the sediments being Early Proterozoic. Hall and Goode (1978) believe the Early Proterozoic age t o be more likely on the following grounds: (1)The Nabberu sediments unconformably overlie the Archaean basement of the Yilgarn Block. This relationship can be directly observed in the field, and is also well illustrated by the magnetic discontinuity between Mabberu and Yilgarn iron formations. Granites elsewhere in the Yilgarn basement have been dated at 2550-2750 Ma (Compston and Arriens, 1968). (2) The westernmost part of the Nabberu Basin is adjacent t o the Gascoyne Block in which anatectic granites intruding Early Proterozoic sediments have been dated at about 1700 Ma (Compston and Arriens, 1968). Williams et al. (1978) have also dated various granites immediately northwest of the western Padbury Sub-basin as between 1550 and 1700 Ma old. These granites are probably related t o the development of the Stanley Fold Belt. (3) Samples of glauconite from basal Nabberu sediments in the southeastern Earaheedy Sub-basin gave minimum K/Ar ages of around 1700 Ma and minimum Rb/Sr ages between 1590 and 1710 Ma (Preiss et al., 1975), c.f. granite ages above.
149 (4) Abundant, thick iron formations of the Lake Superior type are present in the Nabberu Basin and contain microfossils identical to those from the Early Proterozoic Gunflint and Biwabik Iron Formations of the Lake Superior district (Walter et al., 1976). Such iron formations are considered typical of the Early Proterozoic (Cloud, 1973). (5) In the northern Earaheedy Sub-basin folded Nabberu sediments are unconformably overlain by sandstones of the Bangemall Group which have a minimum age of 1100 Ma (Compston and Arriens, 1968). (6) The Nabberu succession is intruded by unmetamorphosed dolerite sills and plugs in several areas; one of these in the eastern part of the basin has been dated at 1050 Ma by Preiss et al. (1975). History of sedimentation and tectonism ( 1 ) Initial sedimentation and basin development. Pre-existing Archaean structures appear t o have had an important control on sedimentation in the Nabberu Basin. The Glengarry Sub-basin, which contains a much greater thickness of basal sediments, is bounded in the east by the extension of a number of major north-northwest-trending faults parallel t o and within the Wiluna greenstone belt (e.g., Celia Lineament). These structures can be shown to have been active during sedimentation as the basal Earaheedy Group sediments (Yelma Formation) thicken quite rapidly immediately t o the west of each fault (Hall and Goode, 1978). Volcanics which occur in the basal sediments in the Glengarry and Padbury Sub-basins may be associated with this faulting and tectonic downwarping during the early development of the basin. Volcanics appear t o be absent on the stable Kingston Platform. Sedimentation in the Nabberu Basin began during a marine transgression over Archaean basement rocks. The presence of glauconite in the basal clastics, the upward lithological variation in the Too100 Sub-group of the Earaheedy Sub-basin from sand to silt t o chemical sediments, and the occurrence of probable deeper-water iron formations at the top of the Frere Formation indicate progressively more offshore and deeper-water conditions associated with transgression. East of the Glengarry Sub-basin the outcrop trace of the basal unconformity is essentially linear for some 300 km despite the very low dip of the sediments. This suggests that the dominantly granitic basement was essentially a peneplain at the time of the Early Proterozoic transgression, as also suggested by Allchurch and Bunting (1976), although the basal Yelma Formation locally drapes over low hills (Hall and Goode, 1978). The peneplain concept is supported by the general lack of thick, coarse-grained clastic sediments in the basal Earaheedy and Glengarry Groups. Higher in the Earaheedy Group abundance of silt is further evidence of the lack of relief in the Archaean hinterland. Hall and Goode (1978) have also shown that the minera-
150 logical composition of clastics in the Earaheedy Group is consistent with a deeply weathered basement. Surprisingly, plagioclase appears t o have been more stable than K feldspar in the inferred Late Archaean weathering profile. ( 2 ) The first sedimentary (Tooloo) cycle, The earliest Tooloo Sub-group sediments in the Earaheedy Sub-basin were deposited during a regional transgression. However, features at the top of the Windidda Formation indicate sub-aerial exposure, suggesting that deposition of the Tooloo Sub-group ended in a regressive phase of sedimentation (algal marsh conditions?). This is partly supported by the presence of reworked aluminous glauconite in the overlying Miningarra Sub-group, possibly indicating non-marine deposition of glauconite following the regression but prior t o the Miningarra transgression. The occurrence of the Frere Formation above transgressive carbonates and below regressive carbonates suggests that the banded iron formations were deposited during the time of maximum transgression, i.e. they represent a facies of chemical deposition further off-shore than carbonate precipitation and occupy a specific stratigraphic niche in the succession. The presence or absence of reworked textures in the chemical sediments depends on basinal water depth, as well as distance off-shore (i.e. it will depend t o some extent on basin floor slopes). The compression of the transgressive facies, and attenuation of the regressive facies (in terms of relative thicknesses) may result from either a relatively slower regression, or, alternatively, slower sedimentation rates during the transgression. The bulk of the sediments of the Tooloo Sub-group were deposited in shallow-water conditions, and the presence of algal stromatolites throughout the sequence indicates water depths of less than 100 m (i.e. within the photic zone). In fact, depths were probably much less than this along the southern margin of the Kingston Platform. Sabkha conditions existed immediately before the iron formations were deposited in some areas. Decreasing water depths and increasing proximity t o the basin edge south and southeast are suggested by a large number of observations in the Earaheedy Subbasin : (a) The Yelma Formation thins eastwards along the southern margin of the Earaheedy Sub-basin and becomes thicker and more argillaceous on the northern margin. (b) Banded iron formations of the Frere Formation thin and become more siliceous eastwards, where shales dominate over chemical sediments. Deeperwater (unreworked) iron formations are only found at the western end of the Earaheedy Sub-basin. (c) Thick shallow-water carbonates of the Windidda Formation are confined to the southeast of the Earaheedy Sub-basin. Similar conditions existed in the western Padbury Sub-basin where thin
151 basal clastics and carbonates (equivalent t o the Yelma and Windidda Formations?) are interbedded with thin pelletal iron formations similar t o those in the Frere Formation. The majority of the banded iron formations in the eastern Padbury Subbasin are of the Hamersley (deeper-water) type. Together with the abundant turbidite sequences, this supports the contention that the eastern Padbury and Glengarry Sub-basins represent the area of maximum downwarping in the Nabberu Basin. The apparent continuity of sedimentation in the western Earaheedy Sub-basin compared to the evidence for sub-aerial exposure in the eastern areas is further support for this concept. ( 3 ) The second sedimentary (Miningarra) cycle. Present outcrop of the Miningarra Sub-group is confined t o the Earaheedy Sub-basin ,although Hall and Goode (1978) have suggested the Millidie Creek Formation of the Padbury Group may be equivalent to the basal Wandiwarra Formation. As in the Tooloo Sub-group all units exhibit abundant shallow-water sedimentary structures, and rare features indicating local sub-aerial exposure. Increasing distance off-shore t o the northwest (as for the Tooloo sedimentary cycle) is supported by: (a) The Wandiwarra Formation appears to be absent in the extreme southeast of the Sub-basin, and becomes finer-grained t o the northwest. It rests disconformably on the Tooloo Sub-group in the southeast, but appears t o be conformable in the northwest. (b) The Princess Ranges Quartzite thins towards the northwest. (c) The Wongawol Formation becomes finer grained t o the northwest. The presence of glauconite in the basal parts of the Miningarra Sub-group, and the general transition from coarse t o fine clastics t o carbonates indicate that the group was deposited in a second major transgression. Detailed analysis of cyclic units in the basal Kulele Creek Limestone near the top of the Miningarra Sub-group shows that the transgression was still continuing at that time (Hall and Goode, 1978), although the Mulgarra Sandstone may represent a regressive phase of sedimentation. ( 4 ) Development of the Stanley Fold Belt. The abundance of detrital plagioclase in the Miningarra Sub-group (Hall and Goode, 1978) indicates fairly rapid erosion of the basement which may in turn be related t o tectonic reactivation of the basement, in contrast to the generally stable relations existing during deposition of the Tooloo Sub-group. Thus it appears that the Stanley orogeny may be traced t o a relatively early stage of basin development. Limited geochronological data suggest that the orogenic-metamorphic peak occurred at about 1550-1700 Ma, with the development of the Stanley Fold Belt and the intrusion of granites. The asymmetric structure of the fold belt indicates deformation resulted from southward-directed forces, centred to the north of the present limits of outcrop.
152 Ore deposits Mineral production from the Nabberu Basin is limited and essentially confined t o the Padbury Sub-basin (Macleod, 1970). Despite early hopes for iron ore discoveries in the Frere Formation, it appears that high-grade enrichment is small and localized, as might be expected for Lake Superior-style iron formations. The best areas of enrichment occur in the Hamersley-style iron formations in the Robinson Range Formation, but no commercial ore bodies have been discovered. The Horseshoe manganese deposits near Peak Hill produced about 400,000 tons of metallurgical grade ore (35-40% Mn, 15--20% Fe). The deposits are associated with Tertiary drainage channels and scree slopes from the iron formations of the Horseshoe Range Beds (Blockley, 1 9 7 5 ~ ) . Minor gold and copper deposits associated with quartz veins and fault zones have been mined in the past (Low, 1963; Macleod, 1970). White talc schist hom the Mt. Seabrook deposit in the western Padbury Sub-basin constitutes the only production from the region at present. Reserves are indicated at approximately 7 million tonnes (Lipple, 1975). The deposit occurs in sandy dolomite at the nose of a major synclinorium. Relationships between the Hamersley and Nabberu Basins
The discovery of the Nabberu Basin in 1973 was very significant as two neighbouring and possibly connected Early Proterozoic basins are now known in Western Australia. Thick banded iron formations are common in both basins but exhibit markedly different features. The question of the relationship between the two basins and their fold belts is probably the most presently argued feature of the Western Australian Precambrian, and one which is of great importance in unravelling the origin of banded iron formations in general. Bunting at al. (1977) believe that the Nabberu Basin and particularly the Earaheedy Group could be considerably younger than the Hamersley Basin (G1700 Ma?). Gee (1979) agrees with the “younger” interpretation for the Earaheedy Group, but considers in addition that the Glengarry-Padbury Groups may be equivalent t o the Wyloo Group. Honvitz and Smith (1978) also consider the equivalence of the Glengarry and Wyloo sequences, but in contrast regard the Earaheedy Group as older and equivalent to the Hamersiey Group. Hall and Goode (1978) tentatively suggested that the Nabberu sequences may represent the shallow-water or marginal equivalents of the Hamersley sediments in a proto-Hamersley Basin, but noted that much more work was needed to be done within each basin before firm conclusions could be reached. This situation remains but some of the remarkable similarities in the sedimentary and tectonic evolution of each basin will be briefly reviewed.
153 Sedimentation (1) Both basins show similar controls with early sedimentation related t o the location of Archaean greenstone belts. (2) Basal sequences in both basins consist of clastics and basaltic volcanics. The Glengarry Sub-basin has less basalts than the Fortescue Group, and the Earaheedy Sub-basin has no volcanics and only thin clastics. However, these features could be expected in marginal platform equivalents t o a subsiding basin. Similar features are observed between the “miogeosynclinal” and “eugeosynclinal” facies of the Early Proterozoic Labrador Trough in Canada (e.g., Frarey, 1976). (3) Lewis et al. (1975) have suggested that the Black Range dolerite suite in the Pilbara Block may be feeders for the Fortescue basalts. E-W dykes (Widgiemooltha Suite) of the same age are prevalent in the Yilgarn Block (see later), and may similarly be related t o feeders for the Padbury basalts. (4) The differences in the types of iron formation in both basins has been ascribed largely t o differing basinal water depths (Goode and Hall, 1976), and is therefore largely superficial. Such variations are the same as for the equivalent Biwabik (Minnesota) and Negaunee (Michigan) iron formations in the Animikie Basin (Sims, 1976). (5) Facies and thickness variations in the Nabberu sequences indicate increasing water depths and distance from shore t o the north and also into the Glengarry Sub-basin. Similarly, recent work has suggested that the early Hamersley Basin was not closed and that basinal depths and distance from shore probably increased southwards. The potentially interesting confluence of these trends is presently covered by younger sediments. (6) In both basins, the units of the second sedimentary cycle (WylooManganese Groups in the Hamersley Basin; Miningarra Sub-group in the Nabberu Basin) lie with unconformity/disconfomity on the older iron formation bearing sequences. Tectonic reactivation of basement seems t o be involved in both cases. (7) Presently available absolute geochronology, while not definitive, certainly is consistent with both basins being active in the time range 2400 t o 1700 Ma, and does not favour Nabberu being younger than Hamersley. Fold belts The age relationship between the Stanley Fold Belt and the Ophthalmian Fold Belt which deform the Early Proterozoic sediments of the Nabberu and Hamersley Basins respectively is uncertain. While deformation in the Stanley Fold Belt increases northwards that in the Ophthalmian Fold Belt increases southwards, so that the Yilgarn and Pilbara Shields are flanked by two Early Proterozoic fold belts with opposite symmetry. If the fold belts are the same age, as limited geochronology on associated granites and the lack of obvious overprinting deformational/metamorphic effects on either fold belt seems to suggest, the problem of tectonic mecha-
154 nism is raised. There could have been an active central zone (now hidden by the younger Bangemall sediments) which resulted in thrusting and folding outwards t o the north and south. Whether this is a result purely of vertical tectonics or involved some form of Precambrian plate-tectonic model is unknown. However, it should be noted that this duality of the fold belts is unusual. In other Early Proterozoic basins of this type (e.g., Labrador, Animikie, Transvaal), the sedimentological and structural picture is a single, asymmetric one. Consequently, tectonically it may be easier t o envisage the two Western Australian fold belts as having developed separately as individual asymmetric structures. Further geochronologic data are needed t o establish which alternative is correct. The development of the Early Proterozoic basins themselves, although probably related t o the development of the fold belts, remains a separate question. Many alternative relationships can be argued (e.g., see Bunting et al., 1977; Hall and Goode, 1978; Horwitz and Smith, 1978; Gee, 1979). However, if all evidence is carefully evaluated and if due regard to variations in tectono-sedimentary environments is given, it is felt that the Hamersley and Nabberu sequences could be equivalent, and that the Nabberu iron formations could represent the shallow-water equivalents of those in the Hamersley Basin. Furthermore, if this correlation is not correct, more problems in*the geologic evolution of the region seem to be raised than solved.
Depositional model for the iron formations The origin of banded iron formations has long remained a complex question, and it is beyond the scope of this chapter t o review fully the data or t o attempt t o explain their origin in detail. However, the increasing amount of data available from the Early Proterozoic basins all over the world is placing restraints on depositional models for BIF, and is also allowing fresh insights into their origin. In Western Australia, the occurrence of two major well-preserved sequences of iron formations of different character in “separate” basins has facilitated these studies (Trendall and Blockley, 1970; Hall and Goode, 1978). The granular Nabberu iron formations are of shallow-water origin (Superior type), and comparable t o those in the Labrador Trough and the northwestern Animikie Basin. The laminated Hamersley iron formations are interpreted as being of deeper-water shelf origin (Algoman-type), and are similar t o those in the Transvaal, Brazilian (Minas Gerais, Serra dos Carajas) and southeastern Animikie Basins. The textural type of BIF is dependent on the particular depositional environment. The transition from Ca-Mg carbonates t o oxide facies iron formations via a complex, usually thin, iron carbonate, silicate and rarely sulphide assemblage is becoming increasingly well known in classic Early Proterozoic sequences, e.g., Nabberu (Hall and Goode, 1978), Hamersley (Davy, 1975),
155
and light penetration)
low relief hinterland
sandstones
transition zone iincluding iron silicates. iron Carbonates)
Fig. 3.15. Schematic facies model showing relations between near-shore clastics, carbonates and off-shore iron formations during sedimentation of the Tooloo Sub-group (after Hall and Goode, 1978). The intersection of the wave base-photic zone limit with the basin floor will determine presence or absence of sedimentary reworking and photosynthetic benthonic organisms in the sediments.
Transvaal (Button, 1976b). In some regions the iron carbonatesilicate zone is much thicker, e.g., Animikie Basin (Sims, 1976); Middleback Range, South Australia (S.V. Bell and N. Lemon, pers. commun., 1978). In the Animikie Basin and Labrador Trough the Ca-Mg carbonate zone is missing. The essential point of these observations is that iron formations occupy a specific niche in the stratigraphic record, and are gradational with other sediments in the basin. The origin of iron formations therefore does not appear t o have an extra-basinal influence. An idealized sedimentary facies model (Hall and Goode, 1978) for deposition in the Tooloo sedimentary cycle of the eastern Nabberu Basin is shown in Fig. 3.15. This is based on the observed lateral and vertical facies relationships in the basin. In reality, of course, deposition is much more complicated than this, being especially dependent on local environmental/palaeogeographic factors and on the interplay between clastic supply, volcanism and chemical precipitation. This diagram proposes that the deposition of BIF occurs offshore from areas of carbonate deposition, and therefore BIF appears in any given stratigraphic column at a time of maximum transgression (note that in most Early Proterozoic sequences, BIF occurs near the top of any transgressive sequence prior to a subsequent erosional or regressive event). This relationship of course is difficult t o reconcile with evaporitic theories of origin for BIF. The close and transitional association with Ca-Mg carbonates suggests a common or related origin for these chemical sediments. Since it is generally agreed that Phanerozoic, and presumably Proterozoic, carbonates were precipitated as a result of organic activity (e.g., Bathurst, 1976), it would seem reasonable t o propose that iron formations were also precipitated as a result of organic activity (see also Becker and Clayton, 1972; Trendall, 1976a). Microfossils were certainly prolific at certain times in the Early Proterozoic
156 (note Hamersley black shales with up t o 17% free carbon). Dominantly benthonic? forms are known in several shallow-water iron formations (Walter et al., 1976), but planktonic forms although suspected (e.g., La Berge, 1973) have yet t o be confirmed. The thickness, extent and broadly synchronous nature of the Early Proterozoic iron formations appear t o be due t o a unique combination of at least three major partially related factors: (1)The development of widespread, large stable shelves suitable for major chemical deposition (low clastic influx due t o commonly associated peneplained basements and arid climates; limited or confined volcanicity). (2) Major algal growth (due to evolution, atmospheric or climatic conditions, availability of clean-water shelf environments?). ( 3 ) Atmospheric evolution which led t o the development of huge, soluble, ferrous iron reservoirs in the ocean basins over Archaean time during reducing conditions, and was perhaps indirectly related t o the organic precipitation of this iron.
Middle Proterozoic M t . Minnie and Bresnahan Basins Though the Bresnahan Group occupies a considerably larger area and is much thicker than the Mount Minnie Group, both have several features in common and can be broadly correlated, though their exact ages are unknown (Daniels, 1975a). Both lie unconformably on the deformed Wyloo Group of the Hamersley Basin, and are themselves unconformably overlain by the Bangemall Group (Fig. 3.1). They are of shallow-water, probably marine origin and appear t o have developed rapidly in small tectonically controlled basins, largely on the southern side of the pre-existing trough responsible for the accumulation of the Wyloo Group. The sequences may represent molasse deposits related t o the final orogenic phase of the Ophthalmian Fold Belt (c.f., Miall, 1978).
Bresnahan Basin The Bresnahan Group is a thick (4900 to possibly 11,900 m), slightly glauconitic, arenaceous sequence (Kunderong Sandstone) with a locally welldeveloped basal conglomerate (Cherrybooka Conglomerate), mLiy D f the clasts of which appear t o have been derived from the Early Proterozoic rocks to the north and the Archaean Sylvania Dome t o the east. Sedimentation appears t o have proceeded in a small basin, the limits of which were probably controlled by major faults (Daniels and Macleod, 1965; Daniels, 1968). Moderate folding along westerly t o northwesterly axes has affected the Bresnahan Group prior t o deposition of the Bangemall Group (Halligan and Daniels, 1964).
167
M t . Minnie Basin As in the Bresnahan Group, marked thickness and facies variations occur over short distances in the Mt. Minnie Group and the nearby similar Tanpool Beds (Williams, 1968; Daniels, 1970). The basal Brogadee Sandstone and equivalent Katanga Conglomerate are overlain by the Wabco Shale and/or the Warramboo Sandstone. Shallow-water sedimentary features are common, and glauconite has been reported from the Tanpool Beds. The deposits probably developed in several small, perhaps unconnected, tectonically active basins. Trainor In1ier The Trainor Inlier occurs as a palaeohigh within flat-lying Bangemall Group sediments in the eastern Bangemall Basin (Williams et al., 1976; A.D.T. Goode and W.D.M. Hall, unpublished). The sequence is dominated by cross-bedded quartz-rich clastic rocks (Cornelia Sandstone of Brake1 and Leech, 1979), and is at least 2500 m thick. The fine- t o coarse-grained sandstone is mature t o supermature, and contains mud pellets, bands of heavy minerals, grit, quartz pebble conglomerate, and fine micaceous sandstone. The inlier contains a number of broad, WNW-trending folds with gentle plunges. Dips are up t o 80", but are more commonly 10"--30". Some strike faulting is evident, and a WNW-trending cleavage dipping 70" northwards has developed in slate along the northern side of the area. Folding becomes more intense northwards. Although the sequence is of unknown age, it could be correlated with the Mt. Minnie and Bresnahan Groups on the basis of lithology and its location along strike. Alternatively it may be related t o the thick Scorpion Group of Bunting et al. (1978), although the latter contains thick carbonate units as well as clastics. Its correlation with the Nabberu Supergroup is not supported by its decreasing degree of deformation southwards and by its vastly greater thickness of quartz arenites than any unit in the Nabberu Basin. Scorpion Group The recently recognized Scorpion Group unconfonnably? underlies the Middle Proterozoic Bangemall Group, and is separated from the older? Early Proterozoic Nabberu Supergroup sediments t o the south by the major Salvation Fault (Bunting et al., 1978; Commander et al., 1979). The sequence is at least 10,000 m thick (minimum), and is composed of shallow-water carbonates and clastics. Deposition of conglomerates in the upper part of the succession appears related t o activity on the Salvation Fault (Bunting et al., 1978), and suggests a possible correlation with the similar Sydney Heads Pass Conglomerate (Commander et al., 1979) which unconfonnably overlies folded Nabberu iron formations south of the fault. Stromatolites, including Conophyton, occur in the lower part of the sequence (Bunting et al., 1978; A.D.T. Goode and W.D.M. Hall, unpublished).
158 Middle Proterozoic Bangemall Basin While the western part of the Bangemall Basin is well known, the eastern extent of the basin has only recently been delineated in any detail. This has led to major revisions of the earlier knowledge (c.f. Daniels, 1975b; Gee, 1975a). The southeastern part of the region has been redefined as the Early Proterozoic Nabberu Basin (Hall and Goode, 1978) while the Manganese Group in the north has been included in the Hamersley Basin succession (Horwitz, 1975b; Button, 1976a; see also this chapter). The northeastern Paterson Province (Blockley and De la Hunty, 1975) has been the recipient of new work following the discovery of the Telfer gold deposit; the younger sediments in this area have been variably included with the Bangemall Group (Goode and Hall, in prep.) or separated as an older succession, the Yeneena Group (Williams et al., 1976). The eastern part of the basin within the Gibson desert has been only recently mapped during independent surveys by the Geological Survey of Western Australia (e.g., Williams et al., 1976) and BHP (Goode and Hall, in prep.). For ease of discussion the western and eastern parts of the basin will be described separately. A regional facies change in the Bangemall Group occurs in the central part of the basin near the 119"E meridian and is the result of an easterly lensing out of dolomite, chert, and sandstone units (Brake1 and Muhling, 1976). It partly corresponds to a major northeasterly trending zone of faulting and basement arches. Correlation of much of the eastern succession with the formations t o the west and north is uncertain.
Geochronology The Bangemall Basin contains Middle Proterozoic sediments, varying from flat-lying and unmetamorphosed in the east, t o folded and metamorphosed to lower greenschist facies in the west and north(?). The sediments lie with angular unconformity on the older Proterozoic Hamersley , Nabberu, Bresnahan and Mt. Minnie successions, and cover an area of approximately 150,000 km2. Isotopic age data are limited, but R,b/Sr determinations on acid lavas near the base, and black shales towards the top of the western Bangemall Group give ages of 1098 k 42 Ma and 1080 f 80 Ma, respectively (Compston and Arriens, 1968; Gee et al., 1976). Intrusive dolerites in the southeastern part of the basin (e.g., Parker Range) are very similar to a nearby sill dated at 1050 Ma (Preiss et al., 1975). Western Bangemall Basin ( 1 ) Stratigraphy, Stratigraphic sub-division was first undertaken in the western region of the Bangemall Basin by Halligan and Daniels (1964) and modified by Daniels (196613), who recognized the regional facies change between
159 the northern and western areas of the basin. Daniels (1969) drew attention to the effect on sedimentation of northeasterly trending basement structures in the Bangemall Group, and the increasing intensity of folding from north to south in the basin. Recent mapping in the centre of the basin by Brakel and Muhling (1976) has outlined another regional facies change and delineated distinct structural provinces. Dolerite sills are common throughout the Bangemall Basin, and were probably intruded before the main period of folding of the Bangemall Group. Dykes of dolerite occur either as feeders t o the sills, or as later NE-trending intrusions that transect the sills and the folds. Three major facies provinces, each with distinctive lithostratigraphic assemblages, are recognized in the western part of the basin (Fig. 3.16). The western facies contains the thickest and most varied succession. The transition between the western and northern facies is the result of both rapid lensing out of some units from the west and changes in lithology along strike
Fig. 3.16. a. Generalized geological map of the western Bangemall Basin (after Brakel and Muhling, 1976).
160
Fig. 3.16. b. Facies provinces in the western Bangemall Basin (after Brakel and Muhling, 1976). (With permission of the Director, Geological Survey of Western Australia).
of others. The western and northern facies share the same upper three formations. The change from the western t o the eastern facies is due to the easterly lensing out of dolomite, chert, and sandstone units and the persistence of shale and siltstone beds. The eastern and northern facies provinces are juxtaposed by faults. The following summary of the various facies provinces is based on Brakel and Muhling (1976). (a) Western facies. The stratigraphy and depositional environments of the western facies are outlined in Table 3.VI. The equivalent Mt. Augustus Sandstone, Tringadee Formation and Yilgatherra Member (Irregully Formation) are the basal units in the basin, but are only locally developed with maximum thicknesses of 600 m, 1000 m and 80 m, respectively. They consist of coarse, often pebbly sandstone and lenses of pebble, cobble, and boulder conglomerate. The clasts are mainly vein quartz, but granitic and metamorphic clasts are also widespread. Siltstone and fine-grained sandstone lenses are rare. Elsewhere the Irregully Formation is the lowest formation of the Bangemall Group. In the northwest Daniels
161 (1970) has sub-divided this unit into nine members. In the west Daniels (1965) records irregular sheets of sedimentary breccia thought t o be the result of local intraformational erosion. The Kiangi Creek Formation overlies and interfingers with the Irregully Formation such that the two are laterally equivalent in part, and in some areas the Kiangi Creek Formation overlies the basal unconformity. Similarly the Jillawarra Formation interfingers with the arenites of the Kiangi Creek Formation, but is absent in most of the northwestern part of the basin. The Discovery Chert, conformably overlying the Jillawarra Formation, is a distinctive, remarkably persistent chert unit which forms the best marker horizon in the Bangemall Group. On the northern margin of the basin it overlies the Irregully Formation and lenses out eastwards. It consists of black, massive chert with diffuse light-coloured laminations that are either planar, wavy, or contorted. A streaky texture is characteristic. The bedding is generally 10 t o 30 cm apart and is wavy, containing irregular depressions and swells. Crystal casts after pyrite are common, while possible gypsum casts are also present. The upper and lower contacts are usually transitional over a thickness of about 3 m ythe beds becoming thinner and less siliceous away from the chert. The Discovery Chert is the first unit in the basin which does not show major facies variations. At the time of its deposition, the sedimentation pattern became simpler and tectonic conditions were relatively quiet. However, slump structures have been found in places and may indicate some seismic activity. TABLE 3.VI Stratigraphy of the western facies of the Western Bangemall Basin (after Daniels, 1975b; Brake1 and Muhling, 1976, and others) Formation
Maximum thickness (m)
Lithology
Kurabuka Formation
1500
shale, mudstone, thin carbonate
300
well-sorted quartz sandstone (local siltstone clasts)
herringbone cross-bedding, ripple marks, flute, groove, and bounce casts common
near shore, high energy (e.g. barrier bar)
micaceous shale, minor quartz sandstone, chert, greywacke
groove, bounce and flute casts, cross-bedding, ripple marks, slumps
open marine shelf, usually below wave base
Mt. Vernon Sandstone
Fords Creek Shale
1300-1900
Comments
Depositional environments
*
coastal lagoons
162 TABLE 3.VI (Continued) Formation
Maximum thickness (m)
Coodardoo Formation
0-350
Curran Forma tion Ullawarra Formation
Devil Creek Formation Discovery Chert
Lithology a n d thickness
Comments
Depositional environments
greywacke, siltstone near t o p
sole marks a t base; better sorted to east
turbidite flows, marine reworking
*
marine shelf, pro-delta
75
shale, mudstone, chert
1800
shale, mudstone, quartzose siltstone, fine sandstone, minor dolomite, chert
abundant load casts
marine platform local shoals
dolomite (locally pyritic), silty shale
stromatolites, oolites, cross-bedding
lagoonal, local shoals
thin-bedded black, pyritic chert
possible gypsum casts
stagnant (anaerobic), a t times hypersaline, usually below wave base
pyrite, gypsum? halite? casts locally abundant
shelf muds below wave base (locally stagnant, hypersaline)
0-800
350
Jillawarra Formation
1400
silty shale, minor chert, dolomite, sandstone
Kiangi Creek Formation
1800
quartz sandstone, increasing shale to east, intershale, minor dolomite a t t o p fingers with Irregully Formation
near-shore marine (shoals, barrier islands)
Irregully Formation
2000
dolomite (locally stromatolitic, pyritic), shale, minor chert, sandstone, conglomerate, breccia
minor dessication cracks
shallow lagoonal, tidal flat
locally developed coarse pebbly sandstone, pebbly to boulder conglomerate
subaerial braided streams, alluvial fans
Mt. Augustus
600 Sandstone, Tringadee Forma- 1650 tion, Yilgatherra 300 Member
* From Brake1 and Muhling (1976), Davy e t al. (1978).
163 The Curran Formation, higher in the sequence, is a distinctive unit of shale, mudstone, and chert extending over a distance of 85 km and overlying the Ullawarra Formation. The overlying Coodardoo Formation is a partially reworked turbidite unit with a sheet-like form at least 175 km long. It is better sorted t o the east, but disappears from the sequence by lensing out. The sands were introduced into the basin from the northwest by turbidity flows and in part transported further east by bottom traction currents. The western facies is dominantly marine with those units from the centre of the stratigraphic column (e.g., Discovery Chert) deposited below wave base (Table 3.VI). The basal clastic wedges were deposited in terrestrial braided streams and alluvial fans, while the uppermost units represent littoral marine deposits. Consequently the overall sequence can be broadly interpreted as forming in a simple marine transgression-regression. (b) Northern facies. The basal sequence (Top Camp Dolomite) of the Bangemall Group along the northern margin of the basin consists of over 2000 m of dolomite, shale, siltstone, chert, and sandstone. Dohmite beds, some with stromatolites, occur at several levels. Some of the bedded cherts appear t o be silicified lutites. The sandstones are dominantly fine-grained, well-bedded, and have internal lamination. Bedding surfaces commonly show current lineations, flute and groove moulds, load casts, and ripple marks. Intraformational conglomerate and slump structures have also been noted (Brakel and Muhling, 1976). South of the Sylvania Dome, a thick wedge of arkose and conglomerate with jaspilite boulders, known as the Prairie Downs Beds (Halligan and Daniels, 1964), occurs within the Top Camp Dolomite. A nearby derivation from the north and northeast is postulated (Daniels, 1966a). The wedge thickens t o the northeast and disconformitg has been recorded at its base. It appears this deposit originated as an alluvial fan related. t o contemporaneous tectonic activity on the Prairie Downs Fault. A glauconitic, manganese-bearing, dolomite-poor sequence t o the east, previously regarded as equivalent t o the Top Camp Dolomite (Brakel and Muhling, 1976) may instead be an older sequence equivalent t o the Manganese Group of the Hamersley Basin (Fig. 3.1). The poor quality of outcrop and the general lack of distinctive rock types makes interpretation and correlation difficult, but it is likely that major revisions in stratigraphy may be required in this area t o the south of the Sylvania nome. Along the regional axis of the basin, the northern facies shares the same upper three formations as the western facies (i.e. Fords Creek Shale, Mt. Vernon Sandstone, Kurabuka Formation), so that the Top Camp Dolomite is equivalent t o the Coodardoo Formation and all the units below it.
164 (c) Eastern facies. The Coobarra Formation, a thick accumulation of coarse-grained sandstone and conglomerate, analogous t o the Mount Augustus Sandstone and Tringadee Formation, unconformably overlies a basement inlier of sheared granitic rocks. In its northwest extent, conglomerate makes up about half of this formation, but elsewhere sandstone predominates. The conglomerate contains well-rounded pebbles, cobbles, and boulders of vein quartz, together with clasts of jaspilite-bearing conglomerate that may be derived from the Early Proterozoic Labouchere Formation (Barnett, 1975) on the southern margin of the Bangemall Basin. Other clasts include sandstone of intraformational derivation. The sandstones of the formation are medium to coarse-grained, moderately t o well-sorted, pebbly quartz arenites. Cross-stratification is common. Bedding trends in the formation are continuous with those in the adjoining Kiangi Creek Formation t o the west, and the two units are equivalent in part. The upper beds of the Kiangi Creek Formation overlie the western portion of the Coobarra Formation and consist of fine-grained laminated dolomite, dolomitic sandstone, medium-grained quartz arenite, and minor siltstone and shale. The eastern portion of the Coobarra Formation is overlain by the Backdoor Formation. Several small rhyolite lava domes up t o 30 m thick occur near the base of the Coobarra sequence (Gee et al., 1976). Brake1 and Muhling (1976) interpret the Coobarra Formation as originating as a terrestrial alluvial fan, in which the proportion of gravel and the grain size of the sand decreased away from the source area. Where the fan entered the sea, reworking of the material in beach and near-shore conditions resulted in sorting of the sand t o produce the Kiangi Creek Formation. Piedmont deposition was ended by a marine transgression. The Backdoor Formation is the lowest laterally extensive unit of the eastern facies but lenses out t o the east (Williams et al., 1976). It consists of shale and siltstone (locally pyritic), with lesser chert, claystone, dolomite and sandstone. Some of the chert is similar in appearance t o the Discovery Chert and the chert beds in the Jillawarra Formation. The sequence was deposited in a shallow marine shelf environment. The Calyie Sandstone, a prominent slightly pyritic sandstone unit over 1000 m thick, overlies the Backdoor Formation. Beds are up to 3 m in thickness and are uniform in composition and grain size. Current grooving, ripple marks and cross-stratification are present, and fragments of siltstones are widespread. The rock is usually a medium-grained quartz arenite but interbeds of shale and siltstone also occur. A glauconitic sandstone bed 10 m thick locally forms the base of the unit. The unit forms a regionally uniform sheet about 150 km long in the area dealt with here, but it extends at least another 220 km farther east (Williams et al., 1976). It is consid-ered t o be a marine deposit. The ngarari Beds form the top unit of the eastern facies, and consist chiefly of shale, siltstone, claystone, and fine-grained sandstone. It is thought t o be a marine deposit.
165
Fig. 3.17. Diagrammatic section showing east-west facies changes in the lower Bangemall Group in the central Bangemall Basin (after Brakel and Muhling, 1976). (With permission of the Director, Geological Survey of Western Australia).
(d) Relationship between eastern and western facies. The bottom portion of the Backdoor Formation is equivalent t o the Jillawarra and Ullawarra Formations of the western facies, so that deposition was also contemporaneous with the Discovery Chert and Devil Creek Formation of the western facies (Fig. 3.17). The relative stratigraphic position of the younger rocks of the eastern facies is uncertain. The Calyie Sandstone may be equivalent t o either the Mount Vernon Sandstone or the Coodardoo Formation. At present, the evidence for either view is inconclusive (see discussion by Brakel and Muhling, 1976), although the former is strongly suggested by relationships in the Ned’s Gap Fault area. If this is correct, the Ilgarari Beds would correlate with the Kurabuka Formation. ( 2 ) Structure. Two major fold trends are present in the western Bangemall Basin (Brakel and Muhling, 1976) and together form the Rrlmundian Fold Belt (Daniels, 1966a), in which the intensity of folding increases southwards. The dominant trend is arcuate from southeast t o easterly, and is parallel t o the major lithological and structural trends in the older Hamersley craton. The other trend is a northeast t o easterly arcuate trend. which is roughly parallel to the northern margin of the Yilgarn Block (Fig. 3.18a). Dips in the Bangemall Basin vary from horizontal t o vertical, and even slightly overturned. Fold styles are essentially concentric, with well-developed, broad synclines and poorly developed cuspate anticlines. Slaty cleavage is only sporadically developed and metamorphism is minimal. The age of this folding is unknown, but possibly related pegmatitic muscovites from Yinnietharra in the Gascoyne Block t o the west give ages of approximately 900-1000 Ma (Aldrich et al., 1959). Regional variations in the intensity and orientation of folds outline four major structural provinces (Fig. 3.18b). Brakel and Muhling (1976) postulate that these provinces relate directly t o the nature of the underlying base-
166 ment. The following descriptions and interpretations are based on their analysis. The Bangemall Basin is broadly synclinorial. The southern flank, which contains numerous folds, contrasts with the gently dipping northern flank. The northern province includes this northern limb which is a relatively undeformed apron around the periphery of the Hamersley-Pilbara craton. The western province extends from the synclinorial axis t o the southern margin of the basin. It is characterized by relatively tight folds that have the typical southeasterly trend of Edmundian folds. Slaty cleavage is more common than in the other provinces. There is n o diminution in intensity of folding at the southern margin, and some of the folds are slightly overturned t o the north. Outliers of Bangemall Group t o the southwest are tiqht synclinal keels. Typically the folds are doubly plunping with axial lengths of about 20 t o 75 km. The terminations of these folds usually occur in northeasterly trending zones. Palaeogeomorphic highs of basement rock are exposed through the Bangemall Group and form marked indentations in its southern margm; some of these basement hiphs are the loci of basal conglomeratic beds. All these features together delineate northeasterly trending regional arches.
167
Fig. 3.18. Structural map (a) and structural provinces ( b ) of the western Bangemall Basin (after Brake1 and Muhling, 1976). Structural provinces are: N - northern; W - western; I - intermediate; E - eastern. (With permission of Director, Geological Survey of Western Australia).
In the southeastern part of the basin (eastern province), folding is gentle and is characterized by irregularly shaped domes and basins having a d-ominant east-northeast direction. Major faults are absent and the structural style indicates a stable basement. The fourth (intermediate) province is characterized by elongate folding with a principal direction that is distinctly arcuate from northeast t o eastearly. The boundaries of this intermediate province are defined by northeasterly trending lineaments which persist well t o the northeast beyond the intermediate zone. The eastern boundary against the eastern province is marked by the Tangadee Lineament, which is expressed by two important features: a line of basement inliers that were highs during initial Bangemall Group sedimentation, and which influenced the formation of wedges of basal conglomerate and sandstone, and a line of east-southeasterly or easterly trending faults arranged en-kchelon. The boundary of the transitional zone
168 with the western province is called the Flint Hill Lineament. Apart from being a conspicuous photolineament, it is a line of structural mismatch across which the southeasterly trending folds of the western province are replaced by northeasterly trending folds. Easterly trending faults in the central part of the basin and north of the intermediate province cut the axial trend of the southeasterly trending folds, and some of these faults are known t o be south-side-up thrusts. The age relationship of the northeasterly trending folds and the more typical southeasterly trend is uncertain and no superposition has been recognized. However, the deformation styles, which are notable for the lack of cleavage, make this relationship inconclusive. There is n o marked mismatch of stratigraphy across the Flint Hill Lineament so that it is not interpreted as a sinistral wrench fault that dragged easterly trending folds into a northeasterly direction. Instead the regional fold patterns suggest synchronous formation of folds of all trends by independently behaving basement segments (Brakel and Muhling, 1976). ( 3 ) Relation of structural provinces to basement. A notable feature of fold trends in most of the western Bangemall Basin, and especially the western province, is the parallelism with earlier Ophthalmian fold trends in the Hamersley craton (Daniels, 1966b), and those in the Gascoyne Province. A basement control t o the Edmundian folding is therefore indicated. The transitional boundary between the northern and western structural provinces in the overlying Bangemall Group may reflect the position of the southern margin of the buried Hamersley craton (Brakel and Muhling, 1976), although Honvitz and Smith (1978) place such a boundary further north. The Bangemall Basin broadly corresponds t o the Ashburton Regional Gravity Ridge (Fraser, 1976) rather than t o the Fortescue Regional Gravity Complex which corresponds more t o the Pilbara-Hamersley craton. Similarly, the northeast trend of the eastern province parallels the structural trend in the northern periphery of the Yilgarn Block, and a similar basement control is indicated. However, the irregular style of folding suggests an interaction with the trend of the western province. The nature of the boundary between the Gascoyne Province and the Yilgarn Block is not properly understood at present, but it appears that on approaching the Yilgarn Block, the southeast trends of the Gascoyne Province are intersected by northeast trends that parallel those in the Yilgarn Block. The transitional zone is a migmatite complex about 70 km wide, trending northeast. The boundaries of this zone, when extrapolated into the Bangemall Basin correspond with the boundaries of the intermediate province, namely the Flint Hill and Tangadee Lineaments. The structural provinces in the Bangemall Basin therefore appear t o reflect movement of basement blocks (Brakel and Muhling, 1976).
169 ( 4 ) Relation o f sedimentation to tectonics. There is a correlation between the major sedimentary facies boundaries and structural province boundaries (c.f. Fig. 3.16b an 3.18b). Furthermore, there is a relationship between the complexity of sedimentation and the degree of deformation. Thus the stratigraphicdly simple northern facies is little deformed, whereas the western facies, which shows complex relationships between many stratigraphic unib , is more strongly deformed. Brakel and Muhling (1976) consider that these regional variations in the style of sedimentation were therefore controlled by the tectonic behaviour of the individual basement segments. ( 5 ) History of sedimentation and tectonics. The sedimentary history of the western part of the Bangemall Basin is one of progressive extension of the area of deposition accompanied by a marine transgression from the west to east. Brakel and Muhling (1976) have discussed the development of this region, and consider that early terrestrial deposition was restricted t o alluvial fans advancing into the basin from local topogranhic high points, and that there was a gentle seaward gradient from the shoreline with a very broad. intertidal and shallow lagoonal zone, in which the Irregully Formation and lower Top Camp Dolomite were laid down. The Kiangi Creek Formation originated as sandy shoals and islands near the basin axis. As the area of deposition extended, these zones migrated outwards, especially towards the south, and laminated muds of the Jillawarra Formation settled out of suspension in the deeper water. Lateral movements of province boundaries resulted in complex interfingering of the various sediments. The more pronounced and fluctuating migration of the southern margin can be attributed to a more tectonically active segment of basement rocks underlying this region. If the interpretation of some crystal casts as evaporite minerals is correct, there were periods when the water in the basin was restricted and hypersaline. During such a stagnant period the Discovery Chert was precipitated in shallow water, perhaps on a tidal flat over most of the existing basin area. Carbonates and muds of the Devil Creek Formation were deposited in lagoonal and shoal environments, while muds of the Ullawarra Formation were laid down in offshore areas. The absence of quartz sand in the shoal zone at this time reflects a lack of supply of medium-grained clastics. Later sedimentation in the west took place on an open marine shelf and resulted in laterally extensive lutite formations in which green shales are conspicuous. At least one sand sheet deposited in the higher sequences was introduced into the basin by turbidity flows, and all others show signs of working by bottom traction currents. Subsequently, barrier bar and lagoonal environments appear t o have prevailed, and this shallowing of the basin could have been a prelude t o the end of deposition. Dolerites were intruded following deposition, and at a later stage the Edmundian Fold Belt was formed. In the older crust activity on the boundaries of the basement tectonic units increased. Those segments which had
170 shown greatest instability during sedimentation were again the most active and the folding in the sedimentary cover overlying them was consequently most intense (Brake1 and Muhling, 1976). Eastern Bangemall Basin The eastern Bangemall Basin was first mapped on a reconnaissance basis in 1975 (Williams et al., 1976; Goode and Hall, in prep.). While there is general agreement on the geology and structure of the region (Fig. 3.19), there are major differences in interpretation between these two surveys. ( 1 ) Stratigraphy. The eastern Bangemall Basin is dominated by generally flatlying or gently dipping fine- t o coarse-grained quartz sandstones with minor carbonates. The sandstones typically are well-sorted and well-rounded. Certain sequences such as that exposed at Constance Headland are feldspathic, but in general feldspar is rare or absent. Siltstones and shales are uncommon in the arenaceous sequences, although they are interbedded with pyritic sandstones in the northwestern part of the region (within the “Yeneena Group” of Williams et al., 1976). Sedimentary structures indicative of formation in shallow water are abundantly preserved throughout the eastern Bangemall Basin. Medium scale planar oblique, planar tabular and minor trough cross-bedding are common throughout the basin, as are intraclastic mud pellets. The feldspathic sandstones and siltstones of the Constance Headland succession are characterised by the presence of shallow marine, near-shore sedimentary features such as symmetrical (rarely asymmetrical) ripples, small- and medium-scale tabular to planar oblique cross-bedding, mud pellets and rarely desiccation cracks indicating local sub-aerial exposure. Local scouring is observed in other parts of the basin. Gypsum has been observed in Yeneena Group dolomites in the northern part of the basin. Pebbly sandstones are common in certain parts of the sequence in the northeastern Bangemall Basin. They generally occur relatively low in the sequence, and are characterized by thin bands of scattered, well-rounded and often poorly sorted pebbles and cobbles. The pebbles comprise vein quartz with lesser quartzite and chert. Williams et al. (1976) have divided the extensive flat-lying sandstones in the central eastern Bangemall Basin into two main but equivalent units, the Calyie (southwest) and McFadden (northeast) Sandstones. Both these sandstones exhibit a number of unusual and apparently contradictory features (Goode and Hall, in prep.): (a) The great majority of clastic grains are well-rounded, well-sorted, monocrystalline quartz, with low proportions of polycrystalline (metamorphic) quartz and labile rock/mineral fragments, i.e. the sandstones have an apparent high maturity index. However, rare euhedral clay pseudomorphs (after ferromagnesian crystals?) are found in some sandstones. (b) Matrix clay is common in the sandstones throughout the region predominating over authigenic quartz cement.
171
Fig. 3.19. Simplified geological and structural map of the eastern Bangemall Basin (after Goode and Hall, in prep.; Williams et al., 1 9 7 6 ; Bunting et al., 1978; Commander et a]., 1979).
172 The combination of features suggesting high maturity of the arenite components with features suggesting a lack of winnowing presents something of a problem when occurring over such a large area. Goode and Hall (in prep.) suggest that much of the clay was supplied by the contemporaneous addition of large quantities of volcanic ash (airborne and reworked aqueous?) to already maturesupermature marine sands. This hypothesis is given strong support by the following features. Petrographic examination of a large number of samples from throughout the eastern Bangemall Basin has revealed volcanic quartz crystal tuffs in several areas. The quartz from these tuffs has a distinctive morphology - embaged, corroded and broken crystals diagnostic of igneous quartz. When partly rounded by sedimentary processes these grains preserve a distinctive scalloped margin which cannot be produced by abrasion (Goode and Hall, in prep.). Subsequently these volcanic quartz grains have been recognized as minor t o major components in most of the southeastern Bangemall sandstones. The crystal tuffs, together with sandstones containing volcanic rock fragments and weathered mineral clasts seem t o be confined t o the McFadden Sandstone in a broad northwest-trending belt along the northern margin of the Trainor Inlier (Fig. 3.19). Volcanic detritus has not been recognized in the southern or northern areas. Along part of the extreme southern margin of the basin in the Carnarvon Range area (Fig. 3.19), the basal Calyie Sandstone is composed of spectacular large-scale, trough cross-bedded, quartz sandstones over a thin, basal, flat-bedded sandstone sequence. The lower sandstone sequence contains horizons of scattered discoidal pebbles and cobbles of rounded quartzsericite schist and angular vein quartz derived from the underlying basement. Goode and Hall (in prep.) and Bunting et al. (1978) have tentatively interpreted the largescale cross-bedding as being of marine or possibly coastal aeolian origin. The latter is in accord with its restricted location on the extreme southern margin of the basin, and its apparent gross similarity with known aeolian deposits (see Walker and Middleton, 1977). It is interesting to note the occurrence of comparable large-scale cross-bedding in the Middle Proterozoic Waterberg Supergroup in South Africa (Meinster and Tickell, 1976), which has been similarly interpreted as being of aeolian origin. Further east along the southern margin of the basin, the Calyie Sandstone is underlain by clastics of the Coonabildie and Marlooyanoo Formations (Commander et al., 1979), while t o the west it is partially equivalent t o the Wonyulgunna Sandstone (Williams et al., 1976). Thin sequences of carbonate rocks have been located in several areas of the eastern Bangemall Basin, generally low in the stratigraphic succession. At least two of these sequences are stromatolitic. The Skates Hills Formation (Williams et al., 1976) forms the basal unit overlying the Trainor Inlier and consists of lenticular, interbedded stromatolitic dolomite, conglomerate and. finer clastics. Columnar stromatolites have been identified as Acaciella c.f.
173 Austrulica by Grey (1978). The Waltha Woora Formation probably forms the basal unit of the Bangemall Group on the northwestern margin of the eastern basin (Goode and Hall, in prep.). It also consists of shallow water stromatolitic dolomites and fine sandstones. Stromatolites from this unit have been classified as Buicalia cupricornia by Thom et al. (1973). This form was initially described by Walter (1972) from the Bangemall Group in the western part of the basin. An apparent unconformity within the upper Bangemall sequence occurs in argillaceous quartz sandstones in the Durba Hills region south of Lake Disappointment. Over numerous outcrops at least several hundred metres in extent, flat-bedded, medium- t o coarse-grained, argillaceous quartz sandstones (McFadden Sandstone) maintain regular dips of up t o 30" below subhorizontal sequences of pebbly argillaceous quartz sandstones (Durba Sandstone of Williams et al., 1976). The underlying, even-dipping McFadden sandstones form sets up t o at least 30 t o 40 m thick. In some localities these apparently regular-dipping sequences can be shown to consist of large to very large-scale trough cross beds. Variation of dips can also be observed on a regional scale between outcrops. The base of the sequence has not been observed, but no shallowing of dips is observed in the outcroppingsequences. The sediments are interpreted as representing very large-scale foresets in a prograding sequence. Rather than indicating a traditional unconformable relationship (Williams et al., 1976), Goode and Hall (in prep.) have tentatively interpreted the deposits as being of deltaic origin and similar t o Carboniferous deltaic sequences in England (Collinson, 1968). In this latter sequence the giant cross-beds are thought t o be infilling facies of major erosional distributory channels in the delta slope deposits (McCabe, 1977), while the overlying sediments were deposited in migrating, fluviatile channels. Further t o the north in the eastern Bangemall Basin (around the Rudall Inlier) the sediments appear t o become progressively more deformed and metamorphosed up to lower greenschist facies grade (Goode and Hall, in prep.). Williams et al. (1976), however, regard the deformed shallow-water rocks in the Paterson Fold Belt as forming the Yeneena Group (Table 3.VII), and being older than the flat-lying Bangemall sediments t o the south (Fig. 3.19). Equivalent rocks immediately t o the northwest unconformably overlying the older Hamersley Basin sequences and Gregory Range Complex have been previously regarded as Bangemall Group correlatives, and have been divided into the Googhenama Conglomerate (base), Bocrabee Sandstone and Wandy Wandy Shale (De la Hunty, 1964; Thom et al., 1973). This latter sequence is underlain by dolomites and sandstones very similar to the Waltha Woora Formation about 70 km t o the north (A.D.T. Goode, unpublished). (2)Structure. As noted above, the rocks of the eastern Bangemall Basin become increasingly deformed northwards toward the Paterson Fold Belt.
174 TABLE 3.VII Stratigraphy o f the Yeeneena Group (After Williams et al., 1976; Chin et al., 1979; and others) Formation
Approximate thickness ( m )
Lithology
Kaliranu Formation Wilki Quartzite Telfer Formation Malu Quartzite Isdell Formation
350+
1000 c. 500 0-7 5 0 >lOOO?
Choorun Formation
2000
Broadhurst Formation
500-1000?
silty dolomite, shale massive quartz sandstone interbedded pyritic quartz sandstone, shale massive quartz sandstone dolomite, dolomitic shale with minor sandstone, siltstone, conglomerate sandstone, micaceous siltstone, quartz pebble conglomerate, shaly dolomite? micaceous siltstone, mudstone, shale, graphitic shale, fine sandstone
-CJiiconformityCoolbro Sandstone
up t o 4000
~
~~~
cross-bedded quartz sandstone with basal conglomerate
The culmination of this deformation occurs in the Rudall area where cores of high-grade Early Proterozoic(?) metamorphics are exposed below the lower greenschist facies Bangemall cover. The Rudall Metamorphics are exposed in major NW-SE-trending anticlinoria with folded Middle Proterozoic rocks preserved along the flanks and around the hinges as both plunging and closed folds. The dome-like anticlines have length : breadth ratios of 3.5 : 1 t o 4.5 : 1 , and hinge plunges of usually less than 30". The folds are commonly asymmetric and slightly overturned south-westwards, with the anticlines having steep southwest limbs and moderately dipping northeast limbs. In alternating sandstone-slate units the folds are tight t o isoclinal. A very strong slaty cleavage, dipping 45-80" northeastwards and parallel t o the fold trends, cuts finer grained rocks, while a less closely spaced fracture cleavage has developed in the quartzite units. Also parallel t o the fold trends are a number of northeastward-dipping thrust faults which dislocate the Middle Proterozoic rocks within the fold belt, and have locally caused the Rudall Metamorphics to structurally overlie the Middle Proterozoic units. The Paterson Fold Belt can be traced t o the northwest, and t o the east under younger Officer Basin sediments. A granite near Mt. Crofton in the Paterson Range area, dated by Trendall (1974) at about 600 Ma, provides the only indication of the possible age of the fold belt. Another granite about 240 km north-northwest of Mt. Crofton has also been dated by W. Compston as being of similar age (see discussion by Trendall, 1974, p. 95). The fold belt appears to correspond to the northwestern arm of the Anketell Regional Gravity Ridge (Fraser, 1976). This
175 can be traced t o the western end of the Musgrave Block (Fig. 3.2), where the Petermann Ranges Orogeny has also been dated at about 600 Ma (Forman, 1966; Wells et al., 1970). Late Proterozoic sediments in the Officer Basin along the southwestern margin of the Musgrave Block were also folded along WNW-trending axes prior t o the deposition of Early Palaeozoic sediments and volcanics (Daniels, 1975d; Playford et al., 1975). The recognition of such a previously unsuspected fold belt has important implications in the interpretation of the tectonic history of Australia (e.g., Austin and Williams, 1978), and particularly in its relation t o the development of the Phanerozoic Canning Basin whose southern margin it parallels. It is noteworthy that the northern margin of the Canning Basin is also marked by a parallel gravity ridge, and by Late Proterozoic tectonism (Bennett and Gellatly, 1970; Plumb and Gemuts, 1976). Such northwest trends form part of a continent-wide zone of major lineaments intermittently active over a period of at least 1500 Ma (see Williams and Goode, 1978).
Relationship between the Yeneena and Bangemall Groups The relationship between the Yeneena and Bangemall Groups is presently disputed. While Williams et al. (1976) regard the latter as unconformably overlying the former, Goode and Hall (in prep.) have tentatively suggested the two sequences to be correlatives. This is because of the similar nature of the sediments, their gradational structural and metamorphic boundary and the lack of any definite unconformable relationships between the two sequences. Furthermore, stromatolites from the Waltha Woora Formation, apparently the basal unit of the Yeneena Group on the northwest margin of the basin, are similar to those in the western Bangemall Group (Thom et al., 1973; Walter, 1972). Moreover, if the Paterson Fold Belt is 600 Ma old (see above) then it is difficult t o reconcile with the supposedly younger (undeformed) 1050-1100 Ma period of Bangemall sedimentation. Regional correlations On the basis of the age relationships outlined earlier, and the deceptively large volcanic contribution t o the eastern Bangemall clastic sediments, the Bangemall Group is correlated with the Bentley Supergroup on the eastern side of the Palaeozoic Officer Basin (Fig. 3.1). The Bentley Supergroup consists largely of terrestrial acid and basic volcanics with lesser quantities of clastics and minor carbonates, and rests unconfonnably on the high-grade basement of the Musgrave Block (Daniels, 1975d). Volcanics from part of the Bentley Supergroup have been dated at 1060 f 140 Ma by Compston and Nesbitt (1967). The Townsend Quartzite has also been included at the top of the Bentley Supergroup by Daniels (1975d), although it unconformably overlies the volcanics in several areas. Major (1975) has correlated the Townsend Quartzite with the Dean Quartz-
176 ite, the basal unit of the Late Proterozoic (Adelaidean) t o Palaeozoic Amadeus Basin. The Townsend Quartzite is conformably? overlain by Late Proterozoic glacial deposits typical of the Adelaidean, and it seems that the Bentley Supergroup should be redefined without the Townsend Quartzite. The general correlation of the eastern Bangemall Group with the Bentley Supergroup leads t o an interesting terrestrial t o marine facies model of the Bangemall Basin. Eastwards from the western Bangemall Basin the generalized facies are: (1)shallow-water dominantly marine carbonates, sandstones, shales and cherts (western Bangemall Group); (2) sandstones, including shallow-water marine and deltaic and possibly coastal aeolian varieties, tuffs and minor shallow-water carbonates (eastern Bangemall Group); ( 3 ) terrestrial volcanics with lesser sediments (Bentley SupergrouD). The Tollu Volcanics in the Bentley Supergroup have been variably interpreted as final fractionation products of the large basic stratiform Giles Complex intrusions within the Musgrave Block (Nesbitt, 1966), or as the extrusive equivalents of widespread acid intrusives of similar age t o the volcanics within the Musgrave Block (A.D.T. Goode, unpublished). A large number of granitic bodies within the Musgrave Block possess ages in the range 1000 t o 1100 Ma (e.g., Compston and Arriens, 1968).
Fig. 3.20. Aerial photograph of the Telfer Dome showing the alternation between sandstone and shale units. Road system (pre-development) within the dome outlines the approximate location of the stratabound gold mineralization.
177
Ore deposits The Telfer gold mine, one of the few opencut gold operations in the world, commenced production in 1977. The deposit contains about 4 million tonnes of ore grading 9.6 g/tonne, and is located in the northeastern part of the basin on the edge of the Great Sandy Desert. All economically significant gold mineralization at Telfer OCCUTS in a series of strata-bound quartz-limonite reefs in which the former pyrite mineralization has been oxidized t o limonite for a depth of 100 metres. The most important gold reef, the Middle Vale Reef (MVR), occurs at the base of a siltstone/shale horizon within the enclosing sandstone formations (Telfer Formation) of the Main Dome structure (Fig. 3.20). Additional ore of a similar grade occurs in two parallel reefs ( E l and E2) lying east of the Main Dome. The MVR varies from a few cm thick t o a maximum width of 2.5 m. The sandstone immediately underlying the MVR is intersected by narrow quartz veinlets carrying lower-grade gold values lying along bedding joints and fractures. Lower-grade mineralization has been recognized in reconnaissance drilling at the adjacent West Dome. This area is still being explored for future open pit reserves (R.T. Russell, pers. commun., 1978; see also Blockley, 1974). Minor stratiform and fault-controlled copper mineralization is known from the western Bangemall Basin (Gee, 197513). Minor sedimentary sequences
A number of small areas of little deformed and metamorphosed sedimentary rocks occur on or adjacent t o the western, southern and eastern margins of the Yilgarn Block (Low, 1975). They are either known or assumed t o be of Proterozoic age but in most cases their regional context is obscure. At least some of the sequences may be related t o those described previously from the major intracratonic basins. Badgeradda Group and Nilling Beds The Nilling Beds consist of arenites with a minor shale containing manganese nodules, and unconformably overlie Archaean migmatite (Perry and Dickens, 1960). The Badgeradda Group probably unconformably overlies the Nilling Beds, and comprises fine t o coarse clastic sediments (Bililly Formation, Woodrarmng Sandstone, Coomberarie Formation). It shows a strong facies similarity t o the Bangemall Group, according t o Low (1975). Wenmillia Formation The Wenmillia Formation comprises about 290 m of basalt, siltstone and shale, and is in fault contact with the Yilgarn Block. Yandanooka Group The Yandanooka Group is separated from the Yilgarn Block by the Dar-
178 ling Fault, but is thought to rest unconformably on the Mullingarra Inlier, and is therefore less than 1120 Maald (Wilson et al., 1960). The Yandanooka Group comprises a thick sequence (>go00 m) of clastic rocks containing variable amounts of volcanic detritus. The Group has been divided into the Arrowsmith Sandstone (base), Arrino Siltstone, Beaconsfield Conglomerate, Enokurra Sandstone and Mt. Scratch Siltstone (top).
Billeranga Group and Dudawa Beds The Billeranga Group comprises a succession of quartzites, sandstones, cherts, siltstones and interbedded intermediate lavas (Neereno Sandstone, Morawa Lavas, Oxley Chert, Campbell Sandstone). The sediments were laid down in shallow-water conditions in a somewhat tectonically unstable zone on Precambrian gneisses. The lavas have been imprecisely dated at 1400 Ma (Compston and Arriens, 1968). The Dudawa Beds (quartzite, stromatolitic? chert) disconformably overlie the Billeranga Group and contain the commercial Dudawa talc deposit.
Moora Group The Moora Group (Logan and Chase, 1961; Low, 1975) comprises chert, siltstone, sandstone and arkose, and unconformably overlies rocks of the Yilgarn Block. Much of the chert is probably silicified carbonate; stromatolite structures are common in parts of the Coomberdale Chert. Some units in the group are tuffaceous. The sequence has been divided into the Copalcarra Sandstone (base), Dalaroo Siltstone, Mokadine Formation, Coomberdale Chert, Noingara Siltstone and Winemaya Quartzite (top). Important talc deposits, such as the Coodawa or Three Springs deposit (Baker, 1976), occur in the Coomberdale Chert. The occurrence of stromatolitic units and associated talc deposits, as well as geographic proximity, suggest the equivalence of the Billeranga-Dudawa and Moora sequences. Similarly the presence of volcanoclastic units suggest the equivalence of these groups to the nearby Yandanooka Group, even though Low (1975) suggests the absence of dolerite dykes at Yandanooka may indicate a younger age t o the other sequences. Cardup Group The Cardup Group of coarse t o fine clastics (Whitby Sandstone, Neerigen Formation, Armadale Shale) lies unconformably on Archaean granite gneiss. The sequence is tentatively suggested to be between 560-590 and 700-750 Ma-old by Compston and Arriens (1968), on the basis of “events” registered by biotites in the adjacent Yilgarn Block.
‘‘Stir1ing-Barren Series” Shallow-water clastic sediments up t o 900 m thick occur in a discontinuous belt over 290 km long on the southern side. of the Yilgarn Block
179 (Clarke et al., 1954). The quartzites and phyllites have undergone low-grade metamorphism but preserve their primary textures and structures. They rest unconformably on Archaean rocks and possibly on c. 1100 Ma granites of the Albany-Fraser Province. However, Turek and Stephenson (1966) have suggested a minimum age of 1340 Ma for the metasediments. Woodline Beds The Woodline Beds consist of a thin sequence of shales with minor sandstone and basal conglomerate, lying on Archaean granite and metasediments. Turek (1966) has dated the shales at 1620 ? 100 Ma. The sequence may be equivalent t o kyanite and garnet bearing quartzites about 100 km along strike t o the northeast (Bunting et al., 1976), and t o the “Stirling-Barren Series” (Low, 1975). These metasediments all lie on the boundary between the Yilgarn Block and the younger Fraser-Albany metamorphic complex. The metamorphosed Mt. Ragged Beds in the southeastern Fraser Block may also represent equivalent sediments. They unconformably overlie migmatites but have indefinite relationships t o c. 1100 Ma-old granites (Doepel, 1975). They contain andalusite, sillimanite and garnet (Gee, 1979). Officer Basin The Phanerozoic Officer Basin contains basal Late Proterozoic tillitic and clastic units adjacent t o the Musgrave Block (e.g., Lowry et al., 1972). These beds have, however, not yet been recognized along the western margin of the basin adjacent to the Western Shield. Where penetrated by drilling, evaporitic-bearing sequences (Babbagoola Beds, Browne Beds) of possible Late Proterozoic age occur unconformably below Cambro-Ordovician basalts in this latter region (Lowry et al., 1972; Trendall and Peers, 1975; Playford et al., 1975). These evaporitic sequences locally form diapiric intrusions (Wells and Kennewell, 1974). Minor occurrences of Late Proterozoic? stromatolitic carbonates and clastics are known in the southwestern Officer Basin (Van de Graaff, 1974; Van de Graaff and Bunting, 1974). METAMORPHIC COMPLEXES
A1 bany-Fraser Province
The tectonic significance of the high-grade Albany-Fraser Province and its relationship to the Yilgarn Block has long been of interest. Despite some more detailed studies in recent years, however, there is still much t o be done before many of the questions relating t o this complex area can be answered. A major mapping programme in progress by the Geological Survey of Westem Australia will undoubtedly help resolve many of the problems in the region, particularly for the poorly known Albany Province. The 80,000 km2 province forms an arcuate belt around the southern and
180
Fig. 3.21. Generalized geological map of the Fraser Province (after Bunting et al., 1976). (With permission of the Director, Geological Survey of Western Australia).
eastern margins of the Yilgarn Block truncating older rock units (Wilson, 1969; Doepel, 1975; Bunting et al., 1976). The Fraser Range area is probably the most varied and well-known part of the province, and will be described in some detail (modified from Doepel, 1975; Bunting et al., 1976; Emerson, 1978; Gee, 1979). Bunting et al. (1976) and Gee (1979) divide the region into five linear tectonic subdivisions east of the Yilgarn Blockthese belts, together with the major folds and faults, all trend northeast (Fig. 3.21).
181 Transition Zone The Transition Zone forms a belt up t o 50 km wide in which Archaean remnants are affected by Proterozoic metamorphism and igneous intrusion. The Archaean remnants contain rocks typical of the greenstone belts of the eastern Yilgarn Block. Proterozoic intrusive rocks include the McKays Creek adamellite, dated at 1725 f 149 Ma (Bunting et al., 1976), and small intrusions penetrating the Archaean rocks. The deformed adamellite from McKays Creek shows peralkaline affinities. Western Gneiss and Granite Zone This zone contains Proterozoic igneous and metamorphic rocks with no recognizable Archaean remnants and with predominantly amphibolite facies metamorphism. It includes rocks dated at 1660 + 40 Ma (Arriens and Lambert, 1969) and 1592 & 36 Ma (Bunting et al., 1976). A younger isochron at 1289 ? 21 Ma is also recorded by Bunting et al. (1976). The zone is equivalent to the Mount Andrew Migmatite east of Norseman (Doepel, 1975). The area is characterized by a variety of felsic gneisses, including migmatite, which have been intruded by mafic and ultramafic rocks, rapakivi granite, and leucocratic granite. The predominant gneiss is a banded microgneiss which commonly contains alkali feldspar, quartz, minor biotite and/or garnet. The pre-metamorphic rock was probably a quartzo-feldspathic sediment (Bunting et al., 1976). Fraser Complex The Fraser Complex forms a linear belt of mafic granulite, with minor felsic granulite, gneiss, granite, and metagabbro (Wilson, 1969). Small intrusive mafic-ultramafic bodies within the complex contain disseminated cop per-nickel sulphide mineralization (Tyrwhitt and Orridge, 1975). The granulite-facies rocks have been dated by Arriens and Lambert (1969) at 1328 f 1 2 Ma, while muscovite from crosscutting pegmatites has given Rb/Sr and K/Ar ages of 1280 and 1210 Ma, respectively (Wilson et al., 1960). The complex is well exposed in the Fraser Range, but t o the northeast is largely obscured by Phanerozoic sediments. However, its subsurface extent is easily defined by an intense gravity ridge and strong magnetic relief. Structural data from the southern termination of the complex indicates that it is in a synformal position, and is therefore a floored body (Gee, 1979). The contact between the Fraser Complex and the Western Gneiss and Granite Zone is marked by the Fraser Fault, a zone of intense shearing and marked change in rock type and metamorphic grade. Geological and geophysical interpretations of the Fraser Fault indicate an upward thrusting of the southeastern high-grade rocks over those older rocks t o the west (Wilson, 1969; Anfiloff and Shaw, 1973; Mathur, 1974). However, the exact
182 nature of the Fraser Fault, its possible tectonic equivalence t o features such as the Grenville Front in Canada, and its northeastern extent and termination are still the subjects of investigation and controversy (see Gee, 1979). The Balladonia Fault forms the eastern margin to the complex (Tyrwhitt and Orridge, 1975). The origin of the Fraser Complex is also the subject of some discussion. Wilson (1969) has suggested that the mafic granulites were derived from a sequence of extrusive basalts and minor gabbro, with some of the felsic bands representing originally quartzo-feldspathic sediments. This has been supported by Bunting et al. (1976) who note the presence of thin quartzite layers, rounded zircons in felsic granulite hands, and possible amvqdales and pillows in mafic granulites. Doepel and Lowry (1970), however, found no evidence for supracrustal material and believed that the mafic granulites were derived from a uniform rock mass of basaltic composition. The presence of primary high pressure features in some of the intrusive mafic--ultramafic Fraser Complex units (D.N. Harley, unpublished) sunports a lower crustal origin for those particular younger units. The situation is very similar t o that in the Musgrave Block where the host metasedimentary granulites contain rounded zircons (Moore and Goode, 1978), while the intrusive mafic- --ultramafic rocks contain primary high pressure or lower crustal features (Goode and Moore, 1975). The petrological similarities between the Fraser and Musgrave Blocks has been long noted (e.g., Wilson, 1969), and it is also interesting t o note their comparable geochronologic histories (see Arriens and Lambert, 1969 Gray, 1978; Gray and Compston, 1978).
Eastern Gneiss and Migmatite Zone These rocks are similar t o those in the Western Gneiss and Granite Zone. Arriens and Lambert (1969) report a 1900 t 230 Pla isochron on gneisses from this zone. Eastern Granite Zone This poorly exposed zone contains a variety of porphyritic, leucocratic and mixed granitic rocks intrusive into gneisses. These intrusives have been dated at 1080 f 50 Ma and 1190 f 25 Ma at Balladonia and Esperence respectively (P.A. Arriens in Doepel, 1975), and are probably similar t o granites dated at 1100 t 50 Ma at Albany (Turek and Stephenson, 1966). K-Ar ages from hornblendes in amphibolites and mafic granulites in the Albany region also show comparable ages of 1060 t o 1160 Ma (Stephenson et al., 1977). The situation is once again similar t o the Fqusqrave Block where 1000-1100 Ma granites associated with a younger east-west fold belt crosscut the older high-pressure metamorphics and basic intrusions (e.g., see Gray, 1978). The Albany region appears t o be broadly similar t o the Fraser region,
183 although there seems t o be some doubt as t o their exact geological and geophysical relationship. It consists dominantly of migmatites (acid gneisses intruded by granitic material on all scales), gneissic granite and intrusive granites (Stephenson, 1973, 1974; Doepel, 1975). No major occurrences of basic rocks are known, however. Rare fayalite granite occurs near Albany (Stephenson and Hensel, 1978); similar unusual rocks are known from the Leeuwin Block (Peers, 1975a). The relationship of the sediments of the “Stirling-Barren Series” and Woodline Beds t o the high-grade metamorphic rocks is not well known. The sediments are at least slightly metamorphosed (correlatives northeast of the Woodline Beds contain kyanite and garnet) but lie on the stable Yilgarn Block marginal t o the orogenic zones. Limited geochronological data suggest they are of similar age or predate the metamorphics. It is possible that the Mt. Ragged Beds t o the east of the high-grade zone may be equivalent t o them, although Doepel (1975) has suggested this latter unit is younger. Until further detailed work is available it would seem best t o regard. the Fraser-Albany Province as a high-grade, tectonically reworked. zone affecting the margins of the Yilgarn Block, with a younger deformed sedimentary cover. Several metamorphic, deformational and intrusive events may have occurred in the period 1900-1200 Ma?, and were followed by a major postorogenic phase of granitic intrusion at about 1100 Ma (see also Bunting et al., 1976; Doepel, 1975). The period of uplift of the high grade rocks (mesumably by the Fraser Fault) is uncertain. In the Musgrave Block uplift took place along major E-W mylonite zones at about 800-950 Ma (see Goode, 1975). The present distribution of “older” sediments such as the Woodline Beds suggests that their depositional basin may have developed on a pre-existing but corresponding zone of weakness t o the present Fraser-Albany Province. The depositional ages of two groups of similar metasediments in the Musgrave Block are 1330 and 1550 Ma (Gray, 1978; Gray and Compston, 1978) suggesting that this pre-existing basin was very extensive and separated the Archaean cratons of the Yilgarn and Gawler Blocks. It is interesting t o note that the earliest fold trends accompanying the 1200 Ma high-grade metamorphism in the Musgrave Block (Nesbitt et al., 1970) are broadly parallel t o those in the Fraser Block. Undeformed acidic and basic volcanics of similar age also form a widespread cover sequence on the Gawler Craton (e.g., Blissett, 1975), and t o a lesser extent the Arunta Block (Page et al., 1976). Gascoyne Province The Gascoyne Province is a complex 40,000 km2 area of Proterozoic sedimentation, granite intrusion and polyphase metamorphism (Daniels, 1 9 7 5 ~ . Williams et al., 1978; Gee, 1979). Its relationships with the Yilgarn Block,
184 the Nabberu, Hamersley and Bangemall Basins, and the Ophthalmian, Stanley and Edmundian Fold Belts are still indefinite and obscure. The oldest rocks in the province are structurally complex migmatites (dominantly quartzo-feldspathic gneisses with granitic phases) which probably represent tectonically reworked Archaean rocks of the Yilgarn Block (Elias and Williams, 1977). These rocks predominate in the south within the Yarlaweelor Gneiss Belt (Fig. 3.1). Towards the north they are overlain unconformably by gneisses and schists tentatively correlated with the Wyloo granitic rocks occur throughout the provGroup (Daniels, 1 9 7 5 ~ ) Various . ince, and have been dated between about 1560 and 1780 Ma (Compston and Arriens, 1968; De Laeter, 1976; Williams et al., 1978). Pegmatitic muscovites from Yinnietharra have been dated by Rb-Sr and K-Ar methods at 9 4 0 4 8 0 Ma and 890-905 Ma, respectively (Aldrich et al., 1959). De Laeter (1976) and Williams et al. (1978) report mineral ages of 750-800 Ma from throughout the province which may represent a period of uplift, postdating Edmundian folding in the adjacent Bangemall Basin. The only notable mineral deposits known in the Gascoyne Province are the columbite- and beryl-bearing Yinnietharra pegmatites. Leeuwin Block The Leeuwin Block is separated from the main part of the Western Australian Shield by the southern part of the Perth Basin (Peers, 1975a). Fold trends vary from north t o northwest, and the western side of the block is highly deformed (cf. Fraser Fault, Wilson, 1969). Rock types are dominantly acidic gneisses with lesser basic rocks and fayalite granites. The southern areas are upper amphibolite facies in grade, while those in the north are low granulite facies (Wilson, 1969). The age of granulite facies metamorphism is 670 ? 25 Ma, the same age as overprinted mica and pegmatites along the western margin of the Yilgarn Block (Compston and Arriens, 1968). Northamp ton Block The Northampton Block is also separated from the main part of the Western Australian Shield, but forms the basement to the northern end of the Perth Basin (Peers, 1975b). Acidic granulites of dominantly greywacke composition are commonly garnet-bearing, but cordierite is also widespread (Wilson, 1969). Minor mafic and quartzitic granulites are also present. The granulites have been dated at 1040 ? 50 Ma, and were probably reworked from older crustal material (Compston and Arriens, 1968). They have been folded on north t o northwest trending axes. The block also contains intrusive granites with marginally developed migmatites. NNE-trending dolerite dykes are associated with minor base metal mineralization (Blockley, 1975d).
185 Mullingarra Inlier
The small Mullingarra Inlier to the west of the Darling Fault consists of biotite gneiss, granite, aplite pegmatite and quartz reefs (Cockbain, 1975). Pegmatitic muscovite has been dated at 1120 Ma (Wilson et al., 1960). The inlier is bounded to the west by the Urella Fault, and forms the basement to the Proterozoic Yandanooka Group t o the east. Rudall Inlier
The Rudall Metamorphics are exposed in an anticlinorium in the Paterson Fold Belt below the younger Middle Proterozoic cover sediments (Blockley and De la Hunty, 1975; Williams et al., 1976; Chin et al., 1979: A.D.T. Goode and W.D.M. Hall, unpublished). The rocks are high-grade, strongly deformed, and include possible tectonically reworked Archaean and Early Proterozoic rocks. No isotopic age data have been published on these rocks, and their relationship t o the Hamersley Basin or the Ophthalmian Fold. Belt is unknown. The Rudall Metamorphics have a strong NW t o NNW structural trend. A cleavage dipping steeply NE is very prominent, and cuts all earlier structures. Bedding is preserved only in the quartzites, and is often folded into isoclinal structures which are commonly steeply plunging and accompanied by a mineral lineation. The psammitic rocks have a well developed schistosity with only a moderately developed mineral lineation. The structural development involved the contemporaneous deformation and metamorphism of Archaean basement(?) and Early Proterozoic cover(?). Large-scale folds were formed, and are outlined by quartzite and hematitequartzite units. The isoclinal minor folds, mineral lineation and. schistosity all date from this period. The later deformation of the cover sequence and. its metamorphic basement resulted in the strong NW structural trend outlined by the cleavage, refolding of the earlier structures including the schistosity, and development of crenulation folds. The metamorphics are dominantly medium- t o coarse-grained, locally gneissic, quartzose psammitic schists with muscovite and chlorite. The schists are interbedded with quartzites with minor biotite and muscovite, and fine-grained quartz schist and quartzsericite schist. Thin but prominent bands of coarse-grained hematite-quartzite and quartz-magnetite gneiss (metamorphosed iron formation) with accessory garnet, amphibole, microcline and apatite occur in the metamorphics as does rare marble, altered ultrabasics, graphitic schists and minor quartz-plagioclase-hornblende (-garnet) amphibolite. Possible Archaean rocks within the metamorphics include medium- t o coarse-grained granite which grades into schistose muscovite-granite gneiss and fine- t o mediwn-grained quartz-plagioclasebiotite gneiss with K feldspar phenocrysts. Sheared and serpentinized me-
186
dium-grained peridotite cumulate may also be Archaean. The Rxdall metamorphics are cut by a number of leucocratic quartz-feldspar granite dykes. EARLY PROTEROZOIC DOLERITE DYKE SUITES
Both the Archaean Yilgarn and Pilbara cratons contain extensive suites of large undeformed basic dykes of approximately the same age but of markedly different trends (Fig. 3.1; see also Horwitz and Smith, 1978). Widgiemooltha D y k e Suite The Widgiemooltha Dyke Suite (Sofoulis, 1966) within the Yilgarn Block is characterized by its dominant cross-cutting east-west trend. The dykes are readily distinguished on the BMR regional aeromagnetic surveys; dykes of opposing magnetic polarities are interpreted as being of different ages (Evans, 1968). Some of the dykes have been dated at 2420 5 30 Ma (Turek, 1966). The largest dyke of the suite is the gabbroic Binneringie Dyke (McCall and Peers, 1971), which extends for approximately 320 km and has a maximum width of 3 km. The dyke has strongly chilled margins and shows abundant magmatic layering (both rhythmic and cryptic) near the contact. The layering has a vertical attitude, and has been compared to the marginal layered zone of the Skaergaard Intrusion (McCall and Peers, 1971). The Jimberlana Intrusion, the other major “dyke” in the suite, shows many similarities to the Great “Dyke” of Rhodesia (Campbell et al., 1970; McClay and Campbell, 1976; Campbell, 1977). It has a length of 180 km and a maximum width of 2.5 km. McClay and Campbell (1976) have demonstrated that the body has a V-shaped cross-section with a vertical depth of about 5.5 km. At several points along its length the intrusion widens into canoe-shaped complexes which are joined by a narrower connecting zone. The intrusion is subdivided into three layered series; the lower layered series, the upper layered series, and the marginal series (Campbell, 1977). The relationship between the three layered series and the geometry of their phase layering is shown in Fig. 3.22. (1)The lower layered series is made up of an alternating sequence of olivine cumulate and bronzite cumulate layers overlain by a thick plagioclaseaugite-hypersthene cumulate layer. Layering at the centres of the complexes is horizontal but steepens towards the edge of the intrusion, finally approaching vertical at the margin. (2) The marginal series occupies the lower part of the intrusion, surrounding the lower layered series at this level and filling the connecting zone between the “canoe” complexes. It shows reversed fractionation in that the olivine and bronzite cumulate layers overlie the plagioclase-augite-hypersthene cumulate layer and, in this respect, is similar t o the marginal series in
187 UPPER LAYERED SERIES
\.
.
.
_
_
:.'.':.'./ .
. . . . . . . . . . . . . .
Granophyric layel
-
Plagioclase augitehypersthene cumulate Bronzite cumulate Olivine cumulate
Fig. 3.22. An idealized cross-section of the Jimberlana intrusion showing the distribution of the major rock types and the relationship between the layered series. The diagram has been simplified by omitting the macro-rhythmic layering in the ultramafic zone of the upper layered series and by showing only two macro-rhythmic units in the lower layered series (after McClay and Campbell, 1976). (With permission of the Editor, Geological Magazine).
the Muskox Intrusion (Bhattacharji and Smith, 1964). (3) The upper layered series, which is found near the t o p of the intrusion, is like the lower layered series and consists of a repetitive macro-rhythmic sequence of olivine, olivine-bronzite and bronzite cumulate layers overlain by a plagioclaseaugite-hypersthene cumulate layer and a granophyric layer. Each of the macro-rhythmic units is associated with a reversal in mineral variation and in Ni-Cr trends and has a zone of high sulphide values at its base. These features are also found at the contact between the upper and lower layered series. The upper layered series is considerably smaller in volume than the lower layered series and its layering is nearer horizontal. The anomalous sulphide-rich zones (Travis, 1975) at the contacts between the upper and lower layered series and between the macro-rhythmic units are thought to have formed from a narrow supercooled zone which developed
188
for a brief period at the bottom of the magma chamber following each new magma pulse (Campbell, 1977). The olivines from these sulphide-rich zones are depleted in Ni, suggesting that Ni has been scavenged by the sulphides. This is only possible if the olivine grains crystallized with the sulphides in the restricted supercooled zone at the bottom of the magma chamber (Campbell, 1977). Black Range D y k e Suite
The Black Range Dyke Suite (Hickman and Lipple, 1975) within the Pilbara Block maintains a dominant n o r t h s o u t h trend (Fig. 3.1; see also Hallberg and Glikson, this volume, Fig. 2.9). The dolerites cross-cut both the Archaean granites and greenstones but seem preferentially developed within the granite batholiths. The Black Range Dyke, about 70 km long and 0.1 km wide, has been dated at 2329 f 89 Ma (Lewis et al., 1975). Lewis et al. (1975) consider that the dyke may have been a feeder for the lowermost volcanics of the Hamersley Basin’s Fortescue Group. Another dyke from the suite, Cajuput Dyke, is regarded as being slightly younger by Hickman and De Laeter (1977). The Black Range Dyke contains a core of coarse-grained gabbro with chilled margins. Granite country rocks have been remelted at the margin. The Coobina ultramafic intrusion (Bye, 1975) within the Sylvania Dome was suggested to belong t o the Early Proterozoic dolerite suites by Horwitz and Smith (1978), but its deformational history suggests it is older (W.D.M. Hall, pers. commun., 1979). Coobina is known for its important stratiform chromite deposits. CARBONATITES AND ALKALINE ROCKS
Carbonatites or associated alkaline rocks are as yet not well known in Australia (e.g., see Moore, 1973). However, at least one major Early Proterozoic carbonatite is known from the Yilgarn Block in Western Australia. The Mt. Weld carbonatite (sovite) near Laverton is about 3 km in diameter and contains sub-economic phosphate-magnetite-pyrochlore deposits (Fig. 3.1). K-Ar dating on biotite has indicated a 2064 rt 40 Ma age for the body (A.J. Cawsey and W.R. Appleby, unpublished). A number of small syenitic alkalic granites occur in the eastern Yilgarn Block, but little is known of their detailed distribution or age. The Fitzgerald Peaks Syenite (Fig. 3.21) near Norseman is 2360 rt 96 m.y. old (De Laeter and Lewis, 1978). Syenite with sodic pyroxenes is also known from within the Lake Teague Dome at the northern edge of the Yilgarn Block (Fig. 3.10; A.D.T. Goode and C.D. Arndt, unpublished). An unusual monazite-bearing riebeckitic amphibolite (metapyroxenite?) occurs within Archaean gneisses at the western end of the Nabberu Basin (A.D.T. Goode and W .D .M . Hall, unpublished).
189 In North America and southern Africa Proterozoic carbonatites are closely associated with important Archaean structures within the major Archaean cratons. In the Superior Province in Canada, Proterozoic carbonatites occur predominantly along the Kapuskasing High (Erdosh, 1979), a major gravity-magnetic anomaly crosscutting the dominant greenstone belt trend. In southern Africa, carbonatites and associated alkalic intrusives occur along the “Murchison Line”, corresponding t o the projection of the Murchison greenstone belt in the Kaapvaal Craton. Similar structural controls have not yet been established for the Western Australian situation, but it could be predicted that such controls may be NNW-trending, parallel t o the Archaean greenstone belts and t o various important long-lived lineaments such as the Celia and Keith-Kilkenny Lineaments (Williams, 1974). PROTEROZOIC CRUSTAL STRUCTURE
The Precambrian geology of Western Australia has long been recognized as similar in many aspects t o the Precambrian of eastern North America and southern Africa (e.g., Trendall, 1968; Anhaeusser et al., 1969: Button, 1976a). The common lithologic, structural and geochronologic development of these terrains is indicative of their comparable tectonic history of development (Fig. 3.23, Table 3.V111), and continued detailed studies of these regions will facilitate a better appreciation of the early history of the Earth. Each of the Western Australian, North American and southern African terrains are dominated by two or more Archaean granite-greenstone provinces of apparently different “age”. There is some indication that these provinces are separated by thin (200-300 km wide) strips of gneissic terrain with ancient and prolonged histories e.g., the Limpopo Belt in south Africa (Robertson and Du Toit, this volume, Chapter 10, part A), The Yarlaweelor Gneiss Belt (Elias and Williams, 1977) and its possible extension, the Southwestern Yilgarn Province * in Australia, and perhaps the Minnesota Valley terrain in North America (Morey and Sims, 1976: Sims, 1976). The boundary of these zones with the more classical granite-greenstone provinces appears t o be relatively sharp. The gneissic zones (or “mobile belts”) appear to be the loci, at least in part, fcr the development of the Early to Middle Proterozoic intracratonic basins and fold belts, such as those described earlier. These curvilinear Proterozoic belts (up t o 400 km wide) which separate the Archaean granitegreenstone provinces, show the repeated development of sedimentary basins, fold belts and igneous intrusion over a long period of time (c. 1300-
* The Southwestern Province of the Yilgarn Block is distinct from the remainder of the Block in containing high-grade metamorphic supracrustal sequences and in yielding the oldest ages in the Block (Hallberg and Glikson, this volume, Chapter 2).
WESTERN AUSTRALIA
h
SOUTHERN
...
7.. n..
in
\
\
Fig. 3.23. Generalized geological maps of the Precambrian of Western Australia, eastern North America and southern Africa.
191 TABLE 3.VIII Comparison of t h e Precambrian geology and important mineral deposits of Western Australia, North America and Southern Africa North America
Southern Africa
Western Australia
Archaean granitegreenstone terrains
Superior-Ungava ( 2 6 0 0 Ma) Churchill C u - Z n , Ni deposits
Rhodesian ( 2 6 0 0 Ma) Yilgarn ( 2 6 0 0 Ma)
Archaean gneissic belts
Kaapvaal (3000 Ma) Ni, Cr deposits
Pilbara ( 3 0 0 0 Ma) Ni, C u - Z n deposits
Minnesota Valley ( 3 5 5 0 Ma)
Limpopo ( 3 6 0 0 Ma)
Yarlaweelor-Southwestern Yilgarn (3000-3400 Ma)
Late Archaean basins
Huronian (22002600? Ma) U deposits
Witwatersrand (2700-2800 Ma) Ventersdorp ( 2 6 0 0 Ma) A u - U deposits
Major dyke swarms
Nipissing ( 2 1 5 0 Ma) Matachewan ( 2 6 9 0 Ma )
Great “Dyke” ( 2 5 2 0 Ma) C r deposits
Widgiemool tha ( 2 4 0 0 Ma) Black Range ( 2 3 0 0 Ma)
Early Proterozoic basins
Animikie BasinLabrador Trough etc. (> 2000 Ma) Fe, Ni, C u - Z n deposits
Transvaal (2100-2300 Ma) Fe, M n deposits
Hamersley (20002400 Ma) Nabberu Fe, M n deposits
Basic intrusions
Sudbury ( 1 9 5 0 Ma) Cu-Ni deposits
Bushveld ( 2 1 0 0 Ma) Pt, Cr deposits
Early Proterozoic fold belts
HudsonianPenokean (1700-1800
Ubendian (1800-2000
Ophthalmian ( 1 7 0 0 Ma) Stanley (15501 7 0 0 ? Ma)
Ma)
Ma) Waterberg (1700 Ma)
Sedimentation
Bresnahan (< 1 7 0 0 Ma) e.g., Mt. Weld ( 2 0 5 0 Ma ) P, N b deposits
e.g. Cargill ( 1 1 0 0 , 1700 Ma) P deposits
e.g., Palabora ( 2 0 5 0 , 1 2 5 0 Ma) P , C u , F , rare earth deposits
Middle Proterozoic Keeweenawan basins-volcanism (1100-1200 Ma) C u deposits
Umkondo ( 1 2 0 0 Ma)
Bangemall ( 1 1001 2 0 0 Ma)
Late Proterozoic zones of tectonic reworking
Namaqualand-Natal (1000 Ma) Mozambique (1000-1200 Ma)
Musgrave (10001 2 0 0 Ma) Fraser (10001 2 0 0 Ma)
Carbonatites, alkalic complexes
Grenville (10001 1 5 0 Ma)
-
Ages (in brackets) are for major granitic intrusions. Ages (in brackets) are oldest reported. Ages (in brackets) are for major granitic intrusions and/or metamorphism.
192 1700 Ma). Between continents the similarity of each basin and fold belt of a particular age is most striking. The apparently greater degree of erosion of Proterozoic cover in Vorth America (Fig. 3.23) allows a better appreciation of the underlying Archaean structures, which in turn can be applied t o the Western Australian and southern African situations. The boundary (suture) between the granite-greenstone Superior Province and the gneissic Minnesota Valley Province corresponds to the break between the overlying undeformed Early Proterozoic (Animikie) sequences and those involved in the Penokean orogeny to the south (Sims, 1976). Such a break also seems consistent with available evidence from the Australian and African regions. In the Australian case, this immediately raises an important problem, that of the “duality” of the mirror-image symmetry of the Pilbara-Hamersley and Yilgarn-Nabberu regions. This may be related to the sutures on both sides of the Yarlaweelor Gneiss Belt (i.e. the boundaries with the Pilbara and Yilgarn basements). The Yarlaweelor Belt extending under the Bangemall Basin cover may represent a tectonically active basement core (see p. 168) which controlled the development and deformation of the Proterozoic basins. Both suture zones correspond to independent assessments of the position of the margins of the Pilbara and Yilgarn cratons (Brake1 and Nuhling, 1976; Wellman, 1978; Horwitz and Smith, 1978; Gee, 1979). Wellman (1978) has suggested that the Australian strips corresponding to the gneissic belts are isostatically compensated and have an anomalous mean density ( A p = O . l ) , particularly near the suture zone, whereas Gibb and Thomas (1976) have proposed that the equivalent Canadian belts are consistently thinner (by 5.5 km) and slightly denser than the adjoining granite-reenstone provinces. There seems t o be no geophysical or geological evidence that the basement beneath the Bangemall Basin is of basaltic material formed during a period of crustal separation as suggested by Smith and Horwitz (1975). Obviously the understanding of the gneissic terrains and their boundary sutures is critical in resolving the problems surrounding our understanding of the evolution of the Archaean crust. It is interesting to note that in all regions the projection of the sutures suggests that they wrap around the Archaean granite-greenstone provinces, and that they both parallel and crosscut the dominant structural trends in these provinces. In all(?) cases, the Early Proterozoic basins developed with more distal facies away from the Archaean granite-greenstone province, and the Early Proterozoic fold belts show directed but decreasing deformation on t o them. These features indicate the greater stability of the granite-greenstone provinces relative t o the gneiss belts. It is also important t o note that mineral deposits associated with these Precambrian sutures account for the bulk of the world’s iron ore (e.g., Hamersley), manganese (e.g., Kalahari), gold (e.g., Witwatersrand), platinum,
193 chromite (e.g., Bushveld) and sulphide nickel (e.g., Sudbury, Thompson) resources, as well as major base metal (e.g., Michigan copper district) and uranium (e.g., Elliot Lake district) reserves, see Table 3.VIII. The Fraser-Musgrave “mobile belt” which crosscuts the eastern margin of the Yilgarn Block and presumably also the intracratonic HamersleyNabberu basin strongly resembles the Grenville Province in Canada (Doig, 1977) and the Mozambique and possibly the Namaqualand-Natal Belts in southern Africa. These intracratonic belts show evidence of a complex and repeated interplay between tectonic reactivation of older basement, sedimentation, highgrade metamorphism, and intrusion. All are markedly transgressive t o older trends, all are elongate and relatively narrow structures (400(?)km wide) and all have similar geochronologic histories. They appear t o be younger analogues of the Limpopo-style belts, and mag similarly separate Archaean provinces of different ages. For example, the Fraser Belt and its possible easterly extension under Phanerozoic sediments (see Wellman, 1978, fig. 13) appears t o separate the 2600 Ma Yilgarn Block from the “younger” 2400 Ma Gawler Block (Cooper et al., 1976; Webb and Thomson, 1977). REFERENCES Adeleye, D.R., 1975. Derivation of fragmentary oolites and pisolites from desiccation cracks. J. Sediment. Petrol., 45: 794-798. Aldrich, L.T., Wetherill, G.W., Bass, M.N., Compston, W., Davis, G.L. and Tilton, G.R., 1959. Mineral age measurements. Carnegie Inst. Washington Yearb., 58?’237-250. Allchurch, P.D. and Bunting, J.A., 1976. The Kaluweerie Conglomerate: a Proterozoic fluviatile sediment from the northeast Yilgarn Block, Western Australia. West. Aust. Geol. SUIT. Annu. Rep., 1975: 83-87. Anfiloff, W. and Shaw, R.D., 1973. The gravity effects of three large uplifted granulite blocks in separate Australian shield areas. Proc. Symposium on Earth’s Gravitational Field and Secular Variations in Position. Univ. N.S.W., Sydney, pp. 273-289. Anhaeusser, C.R. and Button, A., 1976. A review of southern African stratiform ore deposits - their position in time and space. In: K.H. Wolf (Editor), Handbook of Strata-Bound and Stratiform Ore Deposits, Vol. 5. Elsevier, Amsterdam, pp. 257-319. Anhaeusser, C.R., Mason, R., Viljoen, M.J. and Viljoen, R.P., 1969. A reappraisal of some aspects of Precambrian Shield geology. Geol. Soc. Am. Bull., 80: 2175-2200. Arriens, P.A. and Lambert, I.B., 1969. On the age and strontium isotopic geochemistry of granulite-facies rocks from the Fraser Range, Western Australia, and the Musgrave Ranges, central Australia. Spec. Publ. Geol. Soc. Aust., 2: 377-388. Austin, P.M. and Williams, G.E., 1978. Tectonic development of Late Precambrian to Mesozoic Australia through plate motions possibly influenced by the Earth’s rotation. J. Geol. SOC.Aust., 25: 1-21. Ayres, D.E., 1971. The hematite ores of Mount Tom Price and Mount Whaleback, Hamersley Iron Province. Australas. Inst. Min. Metall. Proc., 238: 47-58. Baker, G.F.U., 1976. Talc - Coodawa Deposit at Three Springs, Western Australia. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 4. Industrial Minerals and Rocks. Australas. Inst. Min. Metall. Monogr., 8: 398-401. Baldwin, J.T., 1975. Paraburdoo and Koodaideri iron ore deposits, and comparisons
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201 Richards, J.M., 1972. Proterozoic uraniferous conglomerates, Pilbara, Western Australia. Geol. Soc. Aust., Joint Special Groups Meeting (Abstr.), pp. C 1 8 - C l 9 . Rutland, R.W.R., 197‘- 3rogenic evolution of Australia. Earth Sci. Rev., 12: 161-196. Serebryakov, S.N. ar,,, Semikhatov, M.A., 1974. Riphean and recent stromatolites: a comparison. Am. J. Sci., 274: 556-574. Sims, P.K., 1976. Precambrian tectonics and mineral deposits, Lake Superior region. Econ. G w l . , 71: 1092-1127. Sm;+h, R.E., 1976. Exploration significance of a large hydrothermal alteration system in the Hamersley Basin. CSIRO Min. Res. Lab. Rep., FP14. Smith, R.E., 1979. Interpretation of volcanic relations and low-grade metamorphic alteration, Maddina Volcanics, Western Australia. CSIRO Min. Res. Lab. Rep., FP21. Smith, R.E. and Horwitz, R.C., 1975. Synthesis for the deposition of the Wyloo and Bangemall Groups. Geol. Soc. Aust., 1st Aust. Geol. Convention, “Proterozoic Geology” (Abstr.), p. 90. Sofoulis, J., 1966. Explanatory notes on the Widgiemooltha 1 : 250,000 Geological Sheet. West. Aust. Geol. Surv. Stephenson, N.C.N., 1973. The petrology of the Mt. Manypeaks adamellite and associated high-grade metamorphic rocks near Albany, Western Australia. J. Geol. SOC. Aust., 19: 413-439. Stephenson, N.C.N., 1974. Petrology of the Albany and Torbay adamellite plutons, near Albany, Western Australia. J. Geol. SOC. Aust., 21: 219-246. Stephenson, N.C.N. and Hensel, H.D., 1978. A Precambrian fayalite granite from the south coast of Western Australia. Lithos, 11: 209-218. Stephenson, N.C.N., Russell, T.G., Stubbs, D. and Kalocsai, G.I.Z., 1977. Potassiumargon ages of hornblendes from Precambrian gneisses from the south coast of Western Australia. J.R. Soc. West. Aust., 59: 105-109. Thom, R., Hickman, A.H. and Chin, R.J., 1973. Nullagine 1 : 250,000 Geological Sheet (preliminary edition). West. Aust. Geol. Surv. Travis, G.A., 1975. Nickelcopper sulphide mineralization in the Jimberlana Intrusion. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5: 75-78. Trendall, A.F., 1965. Pisolitic tuffs in Western Australia. West. Aust. Geol. Surv. Annu. Rep., 1964: 51-55. Trendall, A.F., 1968. Three great basins of Precambrian banded iron formation deposition: a systematic comparison. Geol. SOC.Am. Bull., 79: 1527-1544. Trendall, A.F., 1972. Revolution in Earth history. J. Geol. SOC. Aust., 19: 287-311. Trendall, A.F., 1974. The age of a granite near Mt. Crofton, Paterson Range Sheet. West. Aust. Geol. Surv. Annu. Rep., 1973: 92-96. Trendall, A.F., 1975a. Precambrian. Introduction. West. Aust. Geol. Surv. Mem., 2: 2532. Trendall, A.F., 1975b. Hamersley Basin. West. Aust. Geol. Surv. Mem., 2: 119-143. Trendall, A.F., 1975c. The Hamersley Basin -regional geology. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall., Monogr., 5: 411-413. Trendall, A.F., 1975d. Preliminary geochronological results from two Pilbara porphyry bodies. West. Aust. Geol. Surv. Annu. Rep., 1974: 103-106. Trendall, A.F., 1976a. Geology of the Hamersley Basin. 25th Int. Geol. Contr., Excursion Guide 43A. Trendall, A.F., 1976b. Striated and faceted boulders from the Turee Creek Formation - evidence for a possible Huronian glaciation on the Australian continent. West. Aust. Geol. Surv. Annu. Rep., 1975: 88-92. Trendall, A.F., 1979. Revision of the Mt. Bruce Supergroup. West. Aust. Geol. Surv. Annu. Rep., 1978: 63-73.
202 Trendall, A.F. and Blockley, J.G., 1968. Stratigraphy of the Dales Gorge Member of the Brockman Iron Formation, in the Precambrian Hamersley Group of Western Australia. West. Aust. Geol. Surv. Annu. Rep., 1967: 48-53. Trendall, A.F. and Blockley, J.G., 1970. The iron formations of the Precambrian Hamersley Group, Western Australia. Geol. Surv. West. Aust. Bull., 119. Trendall, A.F. and Peers, R., 1975. Precambrian rocks beneath Phanerozoic basins. West. Aust. Geol. Surv. Mem., 2: 217-219. Trendall, A.F. and Pepper, R.S., 1977. Chemical composition of the Brockman Iron Formation. West. Aust. Geol. Surv. Rec., 1976/25. Turek, A., 1966. RubidiumStrontium Isotopic Studies in the Kalgoorlie-Norseman Area, Western Australia. Ph.D. Thesis, Aust. National University (unpubl.). Turek, A. and Stephenson, N.C.N., 1966. The radiometric age of the Albany Granite and the Stirling Range Beds, southwest Australia. J. Geol. SOC.Aust., 13: 449-456. Tyrwhitt, D.S. and Orridge, G.R., 1975. Regional geology and mineralization of the Fraser Range orogenic belt, Western Australia. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5 : 405-408. Van de Graaff, W.J.E., 1974. Explanatory notes on the Mason 1 : 250,000 Geological Sheet. West. Aust. Geol. Surv. Van de Graaf, W.J.E. and Bunting, J.A., 1974. Explanatory notes on the Neale 1: 250,000 Geological Sheet. West, Aust. Geol. Surv. Rec., 1974/9. Walker, R.G. and Middleton, G.V., 1977. Facies models 9. Eolian sands. Geosci. Can., 4: 182-190. Walter, M.R., 1972. Stromatolites and the biostratigraphy of the Australian Precambrian and Cambrian. Spec. Pap. Palaeontol., No. 11, Palaeontol. Assoc. London. Walter, M.R., Goode, A.D.T. and Hall, W.D.M., 1976. Microfossils from a newly discovered Precambrian stromatolitic iron formation in Western Australia. Nature, 261: 22 1-223. Ward, D.F., Coles, I.G. and Carr, W.M.B., 1975. Jimblebar and Western Ridge iron ore deposits, Hamersley Iron Province. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5: 916-924. Webb, A.W. and Thomson, B.P., 1977. Archaean basement rocks in the Gawler Craton, South Australia. Search, 8: 34-36. Wellman, P., 1976. The gravity field of the Australian basement. BMR J.Aust. Geol. Ceophys., 1: 287-290. Wellman, P., 1978. Gravity evidence for abrupt changes in mean crustal density a t the junction of Australian crustal blocks. BMR J. Aust. Geol. Geophys., 3: 153-162. Wells, A.T. and Kennewell, P.J., 1974. Evaporite exploration in the Officer Basin, Western Australia at the Woolnough Hills and Madley Diapirs. Bur. Miner. Resour. Aust. Rec., 1974/194. Wells, A.T., Forman, D.J., Ranford, L.C. and Cook, P.J., 1970. Geology of the Amadeus Basin, central Australia. Bur. Miner. Resour. Aust. Bull., 100. Williams, G.E. and Goode, A.D.T., 1978. Possible western outlet for an ancient Murray River in South Australia. Search, 9: 442-447. Williams, I.R., 1968. Explanatory notes on the Yarraloola 1 : 250,000 Geological Sheet. West. Aust. Geol. Surv. Williams, I.R., 1974. Structural subdivision of the Eastern Goldfields Province, Yilgarn Block. West. Aust. Geol. Surv. Annu. Rep., 1973: 53-59. Williams, I.R., Brakel, A.T., Chin, R.J. and Williams, S.J., 1976. The stratigraphy of the eastern Bangemall Basin and the Paterson Province. West. Aust. Geol. Surv. Annu. Rep., 1975: 79-83.
Williams, S.J., Elias, M. and De Laeter, J.R., 1978. Geochronology and evolution of the eastern Gascoyne Province and the adjacent Yilgarn Block. West. Aust. Geol. Sum. Annu. Rep., 1977: 50-56. Wilson, A.F., 1969. Granulite terrains and their tectonic setting and relationship to associated metamorphic rocks in Australia. Spec. Publ. Geol. SOC.Aust., 2: 243-258. Wilson, A.F., Compston, W., Jeffery, P.M. and Riley, G.H., 1960. Radioactive ages from the Precambrian rocks in Australia. J. Geol. Soc. Aust., 6: 179-196.
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205
Chapter 4
THE PROTEROZOIC OF NORTHERN AUSTRALIA
*
K.A. PLUMB, G.M. DERRICK, R.S. NEEDHAM and R.D. SHAW
Introduction This chapter summarizes the Proterozoic evolution of that part of Australia which lies to the north of the Musgrave Block and southern Canning Basin, and to the west of the Tasman Orogenic Province (Fig. 4.1). Following an overall synthesis of the tectonics, characteristics, structure, and evolution of the region as a whole (Synthesis), detailed descriptions are given of those better exposed and studied areas which best illustrate the various stages of the region’s evolution (Regional Geology). The region encompasses the North Australia Craton bounded by younger marginal mobile belts - the Northeast Orogens and Central Australian M o bile Belts (Fig. 4.2) (Plumb, 1979a). The basement of the North Australian Craton was cratonized during the Early Proterozoic and then overlain by Mid-Proterozoic platform covers, at the same time that the Northeast Orogens and Central Australian Mobile Belts were reworking the edges of the Craton. The Late Proterozoic and Phanerozoic platform covers of the Central Australian Mobile Belts are thicker and more deformed than the stratigraphically equivalent covers of the North Australian Craton. Plumb (1979b) has synthesized the structure and tectonics of northern Australia, and his terminology for structural units is adopted here. Most of the structural units are defined by Plumb and Derrick (1975), Blake and Hodgson (1975), Crohn (1975), Oversby et al. (1975), Shaw and Stewart (1975a) and Smart et al. (1975), but these syntheses have generally been superseded by later work, both published and unpublished, which is incorporated into this chapter. Detailed descriptions of most areas are provided by Bulletins or Reports of the Bureau of Mineral Resources (BMR). The Arunta Inlier, Georgetown Inlier, Mount Isa Orogen, McArthur Basin, and Pine
* Published with the permission of the Director, Bureau of Mineral Resources, Geology and Geophysics.
206
Fig. 4 . 1 . Principal tectonic elements, northern Australia. (Adapted from Plumb, 1979b; Plumb et al., 1979).
Creek Inlier are all subjects of intensive study at present, so the interpretations and conclusions described here may be modified in the near future.
Is0 topic ages Most isotopic ages used in this paper are based on the Rb-Sr or U-Pb zircon methods. Published Rb-Sr ages which used the 1.39 X lo-" yr-' decay constant of 87Rbhave been adjusted to agree with the 1.42 X lo-'' yr-' constant recommended by the IUGS Subcommision for Geochronology
207
Fig. 4.2. Major crustal subdivisions of Australia (After Plumb, 1979a).
(Steiger and Jager, 1977). The ages quoted for many events will therefore be some two per cent lower than in previous regional syntheses of northern Australia, or in the source papers of the original isotopic data.
Time subdivision Recent geochronological studies have shown that the selected chronostratigraphic type sections d o not adequately meet all the criteria laid down for the present chronostratigraphic subdivision of the Australian Precambrian (Dunn et al., 1966); the stratotypes are only of regional, but rarely interregional significance. An informal local time subdivision is therefore used in this paper, more in accord with the concepts of the Subcommission on Precambrian Stratigraphy (James, 1978). The Archaean-Proterozoic boundary is 2500 Ma (James, 1978). Early Proterozoic, from 2500 Ma t o 1700 5 30 Ma, encompasses the deposition, orogenesis, transitional tectonism, and cratonization of the basement rocks of the North Australian Craton. Mid-Proterozoic, from 1700 2 30 Ma t o c. 1000 Ma, encompasses the next major phase of cratonization, that of the polymetamorphic Central Australian Mobile Belts. The interval c. 1000 Ma to base Cambrian is Late Proterozoic. The chronostratigraphic units of Dunn et al. (1966) retain regional value only for those stratigraphic sequences which may be positively correlated with their stratotypes. Carpentarian is now defined as encompassing only the McArthur Basin succession (including the Roper Group) and its correlatives in the adjacent Mount Isa region and Birrindudu Basin; the previously defined (Dunn et al., 1966) base - the transitional tectonic Cliffdale Volcanics -- are part of the basement t o the McArthur Basin. The base of the Carpentarian sequences is diachronous, and its age in some areas may be older than that of some of the youngest transitional domains elsewhere. Adelaidean is restricted t o those successions which may be positively correlated with the stratotype in the Adelaide “Geosyncline” (Thomson, 1966) the Amadeus and Ngalia Basin successions (Preiss et al., 1978), and the late Proterozoic glacial successions on the North Australian Craton.
Tectonic terminology and definitions The structural units of Plumb (1979b) (e.g., Pine Creek Inlier, Mount Isa Orogen, etc., Fig. 4.1), which represent the exposed portions of the Orogenic Domains of the Tectonic Ma-,, are here classified and grouped according to the Tectonic Map of Australia and New Guinea (GSA, 1971). This terminology and classification differs, in detail, from that of Rutland (this volume, Chapter 1). “An Orogenic Province is a group of broadly contemporaneous orogenic domains (individual structural units; author’s note) of similar tectonic history and style that, together with associated transitional domains, form the youngest basement t o an immediately succeeding platform cover.” They incorporate the period of “intense tectonism involving the development and deformation of orthogeosynclines or metamorphic and igneous complexes
209 prior to cratonization” (GSA, 1971). Transitional Domains (not separated on Fig. 4.1)are zones of late t o postorogenic tectonism associated with the last stages of cratonization; they are “transitional in time, place and style between orogenic and cratonic tectonism”. The “deposits unconformably overlie their immediately preceding orogenic domain and are unconformably overlain by immediately succeeding Platform Cover”. Diagnostic characteristics include “molasse-like sediments; abundant volcanic and plutonic rocks; moderate deformation and rare metamorphism of fill” (GSA, 1971). “A formally named Platform Cover is a group of sedimentary basins whose developments are similar, and broadly contemporaneous. Its deposits overlie the immediately preceding Orogenic Province, and associated transitional domains, and spread across older Orogenic Provinces and Platform Covers” (GSA, 1971). They are epicratonic units. Synthesis PRINCIPAL TECTONIC UNITS
The principal tectonic elements are outlined on Fig. 4.1; their age relationsh IS are summarized on Fig. 4.3. KIMBERLEYS -~
KIMBERLEYS TO AUSTRAL
T O CAPE YORK
CFNTQAL
I
RUM JUNGLE
HALLS
CREEK INLIER
1
~
NANAMBU
COMPLEXES
BLOCK
1
I
THE GRANITES - TANIIMI, TENNANT C R INLlERS
Fig. 4.3. Time relationships of Precambrian tectonic elements, northern Australia. (Adapted from Plumb, 1979b).
210 Northern Australia is underlain by various domains of the Early Proterozoic North Australian Orogenic Province, which is bounded by younger Proterozoic mobile belts of the Central Australian Orogenic Province along its eastern and southern margins - partly exposed in the Georgetown and Coen Inliers (the Northeast Queensland Precambrian Province of Withnall et al., 1980), in the Mount Isa Orogen, and in the southern Arunta Inlier. In the north, inliers of basement rocks belonging to the Archaean West Australian Orogenic Province - The Rum Jungle and Nanambu Complexes - underlie the North Australian Orogenic Province, but the subsurface extent of Archaean basement is unknown: it may possibly underlie the whole of the North Australian Craton (Plumb, 1979a,b). The region is bounded t o the east by the Phanerozoic Tasman Orogenic Province. The mainly Mid-Proterozoic North Australian Platform Cover was deposited on the cratonized North Australian Orogenic Province at the same time that sedimentation, deformation, and metamorphism was continuing in the adjacent Central Australian Orogenic Province. Late Proterozoic and Palaeozoic covers, parts of the Central Australian Platform Cover, overlie both the North and Central Australian Orogenic Provinces and the North Australian Platform Cover. The Mesozoic-Cainozoic Trans-Australian Platform Cover was deposited on both the Tasman Orogenic Province and the various older domains farther to the west. The mega-structural units of Plumb (1979a) (Fig. 4.2) are crustal blocks which each incorporate several of the Tectonic Map units above: The North Australian Craton incorporates all of the block containing the North Australian Orogenic Province plus the overlying North Australian and younger Platform Covers; the Northeast Orogens comprise the Mount Isa Orogen, the Northeast Queensland Precambrian Province, and their subsurface extensions; the southern Arunta Inlier, the Musgrave Block, and their overlying platform covers are part of the Central Australian Mobile Belts. Because of their spatial relationships and tectonic setting, it has been suggested that the Northeast Orogens, particularly the Mount Isa Orogen, might be more properly assigned t o the North Australian Orogenic Province (Rutland, 1973, 1976; Plumb, 1979a; Plumb et al., 1980). The Tectonic Map classification will be retained in this chapter. SUMMARY DESCRIPTIONS
West Australian Orogenic Province (Archaean basement inliers) The oldest rocks identified in northern Australia are small inliers of Archaean granite, gneiss, and metasediments -the Rum Jungle, and Nanambu Complexes, and possibly the Litchfield Complex (Fig. 4.1) - which form basement to the main sequence of the Pine Creek Inlier. The metasediments and gneisses were intruded by massive granites about 2450 Ma ago. The com-
211 plexes unconformably underlie the Early Proterozoic succession of the Pine Creek Inlier, but were later reactivated and emplaced into the Pine Creek Inlier rocks as mantle gneiss domes during c. 1800 Ma-old deformation and metamorphism.
North Australian Orogenic Province (Early Proterozoic orogenic domains) These domains all contain Early Proterozoic geosynclinal successions which were deformed, metamorphosed, and finally cratonized before the end of the Early Proterozoic. The domains all have similar ages (Fig. 4.3), rock types, and tectonic styles, and final cratonisation took place via a welldefined stage of transitional tectonism; most domains went through their principal history of geosynclinal sedimentation, orogenesis, and transitional tectonism within total time spans of about 250 Ma (Plumb, 1979b). The exposed domains are only samples of much larger structures (geosynclines?), whose full extent and interrelationships are unknown; Fig. 4.4 attempts to extrapolate the domains into the subsurface, with the aid of gravity anomalies or similar data. Stratigraphic successions of the orogenic domains can exceed 7 km in thickness. Deposits range from quartz-rich turbidities, through shallowmarine and evaporite-earbonate complexes, t o fluviatile facies. Most, if not all, of the geosynclines are interpreted as ensialic structures. The strata are invariably tightly to isoclinally folded, slaty cleavage or schistosity is ubiquitous, and superposed folding is common. Most metamorphic rocks lie within the greenschist facies, but reach amphibolite and low-granulite facies locally. Metamorphism is mostly of the low pressure-intermediate type. Both basic and acid intrusive activity is common. The transitional domains are characterized by widespread mildly deformed acid volcanics and late- or post-tectonic granite plutons. The North Australian Orogenic Province may be subdivided into older domains and younger domains (Plumb, 1979b), which are now named following Rutland (this volume, Chapter 2). The McClintock Subprovince (from McClintock Hills, near Halls Creek) comprises all those rocks which underwent regional metamorphism about 1950-1900 Ma ago, and this was succeeded, locally, by transitional tectonism between 1900 and 1800 Ma (Rutland’s Halls Creek Subprovince is only one domain of our McClintock Subprovince). The Palmerston Subprovince (from the original settlement at Darwin) comprises all those domains which were metamorphosed at about 1840-1800 Ma ago, and followed by transitional tectonism between about 1780 and 1680 Ma. Rocks metamorphosed as part of the McClintock Subprovince are widespread right across the North Australian Orogenic Province, but the related transitional tectonism was restricted to the area west of the Halls Creek Mo-
212
Fig. 4.4. Inferred distribution of orogenic and transitional domains, northern Australia. (Adapted from Plumb, 1979b; Plumb et al., 1980). Age ranges refer t o cratonization: i.e. deformation and metamorphism of orogenic domains, and subsequent transitional tectonism.
213
bile Zone. The Palmerston Subprovince is restricted to the area t o the east of the Halls Creek Mobile Zone, where it apparently forms a series of belts lying between and above McClintock Subprovince rocks.
North Australian Platform Cover (Early t o Mid-Proterozoic platform covers) Following cratonization of the constituent domains of the North Australian Orogenic Province, widespread mildly deformed, unmetamorphosed sequences of quartz-rich arenites and lutites, shelf carbonates, and flood basalts were deposited on the North Australian Craton, at the same time that intense tectonism continued in the Central Australian Orogenic Province (Fig. 4.3). The component basins of the covers show very similar deposits and tectonic styles, but show significant differences in age. The late Early Proterozoic Kimberley Basin overlies the Halls Creek Inlier and developed to the west of the Halls Creek Mobile Zone between about 1815 and 1760 Ma ago, at the same time as the Palmerston Subprovince was undergoing its final orogenesis and transitional tectonism t o the east. The next group of covers - the Carpentarian McArthur, South Nicholson, and Birrindudu Basins, and the Lawn Hill Platform - developed diachronously to the east of the Halls Creek Mobile Zone between about 1700 and 1300 Ma ago, principally on a basement of Palmerston Subprovince rocks, and at the same time that stratigraphically equivalent and lithologically similar successions were being deposited, deformed and metamorphosed in the adjacent Mount Isa Orogen. A feature of these basins is broad stable shelves (e.g., Arnhem and Wearyan Shelves; Fig. 4.10) separated by narrow, commonly fault-bounded zones of relatively intense tectonism (e.g., Batten Fault Zone, Fig. 4.10; Halls Creek Mobile Zone, Fig. 4.1). The shelves are characterized by shallowly dipping sequences with thicknesses of around 5 km over wide areas. The intervening mobile zones have sequences up to 1 2 km thick and are moderately intensely deformed, mostly in response to movements on the major faults; some of these intracratonic mobile zones (Batten Trough) may be compared with aulacogens. Successions in both the shelves and troughs comprise accumulations of regularly alternating very shallow-marine, intertidal, supratidal, and fluviatile facies ; evaporite-carbonate complexes are abundant. The basins are interpreted as various types of intercontinental and continental-margin marine embayments. The late Mid-Proterozoic (about 1200-900 Ma) Victoria River Basin has been assigned to the Central Australian Platform Cover (Figs. 4.1 and 4.3) (GSA, 1971) but, in view of its age relationship to the final cratonization of the Central Australian Mobile Belts, it is appropriate to describe it with the other Mid-Proterozoic covers. In fact, its overall tectonic style and position on top of the North Australian Craton is more related t o the North Australian Platform Cover, and reclassification may be appropriate.
214 Halls Creek and King Leopold Mobile Zones. The development and deformation of the various platform covers in the Kimberley and adjacent regions has been dominated by two intersecting linear belts of repeated intracratonic deformation - the Halls Creek and King Leopold Mobile Zones - which flank the stable Kimberley and Sturt Blocks (Figs. 4.1. and 4.13). These mobile zones are among the most important lineaments of northern Australia. The Halls Creek and King Leopold Mobile Zones must not be confused with the Halls Creek Inlier: the Halls Creek Inlier is an orogenic and transitional domain which happens t o be exposed within the Halls Creek and King Leopold Mobile Zones; the mobile zones are much longer-lasting, younger, essentially intracratonic structures. The mobile zones have been the site of repeated deformation from at least the Early Proterozoic t o the Permian, and each zone has been deformed independently of the other, and at different times (Fig. 4.6). The major movements which distinguish the mobile zones from the adjacent stable blocks postdate the cratonization of the rocks of the Halls Creek Inlier. Although the principal structures within the mobile zones appear t o have originated from at least the time of the metamorphism in the Halls Creek Inlier, their earlier influence on the development of the Inlier remains enigmatic (see later discussion). The Halls Creek Mobile Zone was deformed by left-lateral strike-slip displacements, on several anastornosing major faults, with a total cumulative displacement of at least 110 km (Plumb and Gemuts, 1976); large vertical movements are also apparent. The King Leopold Mobile Zone was deformed by reverse faulting and folding, overturned towards the northeast; a shortening of 80 km has been calculated across the exposed part of the zone (Sofoulis et al., 1971). Where the zones meet, major structures of each truncate structures of the other; they do not form an orocline. The mobile zones have been the site, at various times, of alternate uplift and subsidence. Stratigraphic equivalents of both the Kimberley and Victoria River Basins have been locally metamorphosed t o low-greenschist facies along major shears in the Halls Creek Mobile Zone; alteration ages indicate movements around 1650 and 1000 Ma (Bofinger, 1967). Kimberley Basin rocks were locally metamorphosed around 1500 and 600 Ma ago in the King Leopold Mobile Zone (Bennett and Gellatly, 1970); the 600 Ma metamorphism locally reached amphibolite facies (Sofoulis et al., 1971). Central Australian Orogenic Province (Early to Mid-Proterozoic orogenic d o ma ins) The three Early t o Mid-Proterozoic mobile belts which lie along the eastern and southern margins of the North Australian Craton lack the unity of age, spatial and tectonic setting, and tectonic style which typifies the North Australian Orogenic Province: the Mount Isa Orogen and Northeast
21 5 Queensland Precambrian Province have been separated out as the Northeast Orogens; the southern Arunta Inlier, and the Musgrave Block farther t o the south, are part of the Central Australian Mobile Belts (Plumb, 1979a). Only the Arunta Inlier and Musgrave Block can be positively related t o the main Central Australian Orogenic Province, of central and southern Australia. The Mount Isa Orogen is situated along the eastern edge of the North Australian Craton (Figs. 4.2 and 4.4). Early Proterozoic rocks, continuous with those of the North Australian Orogenic Province t o the west, were deformed, metamorphosed and intruded by granite, and they then formed basement to a succession of late Early Proterozoic and Carpentarian rocks which were deposited, and intensely deformed, metamorphosed, and intruded by granite, along with the basement, all at the same time that Carpentarian strata of the North Australian Platform Cover (McArthur Basin-Lawn Hill Platform) were being deposited on the Craton to the west. Following cratonization, a younger Carpentarian element of the North Australian Platform Cover (South Nicholson Basin) was deposited on top of the Lawn Hill Platform. The orogen is interpreted as having developed as a rifted continental margin. Up to 20 km of shallow-marine, intertidal, fluviatile, and turbidite facies and continental tholeiitic basalts were deposited in a developing sequence of rift basins, shallow shelves, and parageosynclines. Facies resemble those of the adjacent platform covers, into which the orogen grades, but those of the orogen are thicker and contain more volcanic rocks. Several periods of deformation and metamorphism are apparent. Metamorphism ranges up t o amphibolite facies, of the low pressure-high temperature type. The Northeast Queensland Precambrian Province is a polymetamorphic terrain, situated at the eastern edge of the Australian Precambrian Craton, in which Early or Mid-Proterozoic sedimentary and volcanic rocks underwent repeated cycles of deformation, metamorphism, and granite emplacement during the Mid and Late Proterozoic (1570-960 Ma), and then again during the Palaeozoic under the influence of the adjacent Tasman Orogenic Province. The province is separated from the Tasman Orogenic Province by the Palmerville Fault. It is unconformably overlain by Mesozoic rocks of the Trans-Australian Platform Cover, which obscure its relationships with Proterozoic domains farther to the west (Fig. 4.1). 87Sr/86Srinitial ratios of metasediments suggest that the earliest sediments in the Georgetown Inlier are not very much older than 1600 Ma. They may have accumulated in a rifted marine basin or marginal sea. Metamorphism ranges from greenschist to high-amphibolite and marginal granulite facies, of the low pressure-intermediate type. The southern Arunta Inlier is the exposed part of a polymetamorphic belt adjacent to the southern margin of the North Australian Craton (Figs. 4.1 and 4.2), in which three sequences of Early Proterozoic rocks, separated by unconformities and apparently continuous with sequences of the North
216 Australian Orogenic Province t o the north, were repeatedly metamorphosed, deformed, and intruded by granite throughout the Early and Mid-Proterozoic (>1800-1000 Ma), overlain by the Adelaidean-Palaeozoic Central Australian Platform Cover (Amadeus-Ngalia-eorgina Basins), and then further reactivated by mid-Palaeozoic overthrusting and retrogressive metamorphism. Tectonic zones, each with a distinctive evolutionary history, are recognized. Metamorphic histories are complex, but most regional metamorphism appears to belong to the low pressure - intermediate facies series; grades range up t o granulite facies. The Musgrave Block (Fig. 4.1) has many similarities to the Arunta Inlier, and it is quite likely that they link up beneath the Amadeus Basin, to form a single intracontinental mobile belt between the North Australian and Gawler Cratons (Figs. 4.2 and 4.4). It is a typical example of the intracontinental ensialic mobile belts so characteristic of Proterozoic terrains, with all of their problems of evolution, origin, and interpretation. Central Australian Platform Cover (Late Proterozoic (Adelaidean) platform covers) Following final cratonization (c. 900 Ma) of the younger domains of the Central Australian Orogenic Province (Arunta Inlier and Musgrave Block) the principal elements of the Central Australian Platform Cover in northern Australia were deposited. The Late Proterozoic sequences may be correlated with the type Adelaidean succession of South Australia (Rutland et al., this volume, Chapter 5). These sequences then pass up, with only minor breaks, into the Palaeozoic, when the Central Australian Platform Cover was terminated at various times, in different places, during the Early, Middle, and Late Palaeozoic. The first covers were shallow-marine to fluviatile quartz-rich arenites and subordinate lutites, and carbonate rocks, deposited on the Arunta Inlier and Musgrave Block in the Amadeus, Ngalia, and southern Georgina Basins. These basins are only the structurally preserved remnants of more extensive, probably interconnected, sedimentary covers which existed prior t o Palaeozoic deformation. During the Adelaidean, the Amadeus Basin was probably also connected with the Adelaide Geosyncline and Officer Basin to the south. During the late Adelaidean the extent of sedimentation expanded, with glacigene sequences known from central Australia, the Kimberley region, and the southern end of the Mount Isa Orogen. The covers expanded even further, over most of the North Australian Craton, during the Palaeozoic (Plumb, 1979a,b), but these latter covers are not the subject of this chapter. The tectonic style of the various elements of the Central Australian Platform Cover can be directly related to their basements: those elements which lie within the North Australian Craton are generally much thinner and less deformed than those within the Central Australian Mobile Belts. In the
217
. .
.
141W 0
200
400
600km I
Fault. type unknown High angle fault showmg downlhrown block
-
High angle reverse faull
-
L
Thrust fault. showing upper plate
+--
Antic/ine
+
a
Syncline
Downwarp
Upwarp
Strike slip faulr
Fig. 4.5. Principal structures, northern Australia. (Adapted from Plumb, 1979b).
latter, up to 10 km of Adelaidean and Palaeozoic sediments were deposited, with several breaks, and the margins of the present basins were deformed into thrust-nappe complexes, thrust sheets, and reverse faults; firstly by the localized Peterman Ranges Orogeny at the end of the Proterozoic, and then by the more extensive Carboniferous Alice Springs-Mount Eclipse Orogeny. In the Kimberley region, the glacial successions were deformed within the Halls Creek and King Leopold Mobile Zones.
218 REGIONAL STRUCTURE
The major structures trend roughly northerly (NNW to NE) and northwest throughout the North Australian Craton and Northeast Orogens, but the trends change in the Central Australian Mobile Belts and in the Tasman Orogenic Province (Fig. 4.5). Plumb (1979b) has postulated that a fundamental basement fracture pattern was established throughout northern Australia by no later than the end of the Early Proterozoic, and the subsequent tectonic development of the region then resulted from the repeated reactivation of such major fractures. Because of the geometric regularity of the pattern it was suggested that a uniform crust, perhaps an Archaean continental crust, underlies the whole region. The northerly trending structures include many major faults with large strike-slip movements, and oblique folds and reverse faults are associated with them: in the west, individual faults of the Halls Creek Mobile Zone show left-lateral displacement of 40 km or more (Dow and Gemuts, 1969; Plumb and Gemuts, 1976); in the east, right-lateral displacements of up to 25 km are an important component of faults in the Mount Isa Orogen and Batten Fault Zone (Wilson and Derrick, 1976; Plumb et al., 1980). The northwest-trending faults throughout the region commonly show left-lateral strike-slip displacements and, in the east, these may be interpreted as forming a conjugate set with the right-lateral northerly trending faults (Plumb et al., 1980). The westerly trending faults of the region are commonly thrust and reverse faults (e.g., King Leopold Mobile Zone; Sofoulis et al., 1971) or thrust-nappe complexes (e.g., Central Australia; Wells et al., 1970). An interrelationship between movements on the northerly trending, strike-slip faults and the west-trending folds and thrusts has been demonstrated in the Kimberley region (Plumb and Gemuts, 1976) and in the McArthur Basin (Plumb et al., 1980). A simple model t o explain all these structural interrelationships has been suggested in which the central block of northern Australia, roughly between the Kimberleys and Mount Isa, has moved northwards relative to the blocks on either side, and locally the horizontal displacements have been absorbed by crumpling against or between rigid blocks (Plumb, 1979). However, this model is probably a gross simplification. The cumulative displacement across the Halls Creek Mobile Zone is at least 110 km, whereas displacements on the fault zones in the east were considerably less. The regional geological relationships appear to restrict possible “plate” displacements to very modest proportions.
219 MINERAL DEPOSITS
Early Proterozoic
Uranium deposits constitute the economically most important reserves of the North Australian Orogenic Province, principally in the Alligator Rivers and Rum Jungle fields of the Pine Creek Inlier. The principal deposits are stratabound in carbonaceous pelites, they are situated adjacent t o Archaean basement domes, and they show long histories of epigenetic concentration. The South Alligator Valley deposits are stratabound in carbonaceous pelites, where they are in contact with acid volcanics of the overlying transitional domain. Stratabound base-metal deposits, commonly volcanogenic (e.g., Ag-Pb-Zn, Cu, Cu-Zn) or supergene (Fe), and hydrothermal base-metal deposits (Cu-AuB; Au) are scattered through most domains, but most are subeconomic. Small hydrothermal deposits of Au, Sn, Cu, Ta, and W are associated with many granites in the transitional domains.
Mid-Proterozoic The Carpentarian of the McArthur Basin-Lawn Hill Platform-Mount Isa Orogen is noted for lead-zinc deposits, particularly the large stratiform shalehosted deposits - Mount Isa, McArthur River, Hilton, and Lady Loretta. These may be correlated stratigraphically; all are hosted by similar dolomiticcarbonaceous shales with tuffs; all occur in similar tectonic situations depressions adjacent t o syndepositional faults in major rift zones. Carbonatehosted (Mississippi Valley-type) Pb-Zn deposits are also common, but are mostly small. Recently discovered stratiform Pb-Zn deposits in slightly older strata of the Mount Isa Orogen (Pegmont-Squirrel Hills type) resemble the deposit at Broken Hill, N.S.W. Stratabound copper deposits in brecciated sedimentary rocks are widespread; most noteworthy is the huge Mount Isa deposit, which is spatially associated with, but of different origin to, the Mount Isa Pb-Zn deposit. Other potentially significant Cu deposits are associated with volcanic breccia pipes and dykes (Redbank type). Small uranium deposits are scattered through various parts of the Carpentarian, or slightly older rocks, particularly in sequences associated with basic volcanics. The major deposit of the region is the massive pyrometasomatic skarn deposit at Mary Kathleen, in the Mount Isa Orogen. Stratigraphically equivalent, subeconomic iron deposits, consisting of pisolitic hematitesiderite-chamosite, occur in the South Nicholson and McArthur Basins.
220 Regional Geology ARCHAEAN BASEMENT INLIERS (Principal references: Rhodes, 1965; Stephansson and Johnson, 1976; Needham and Stuart-Smith, 1980)
The small inliers of Archaean granite, gneiss, and metasediments, which unconformably underlie the main sequence of the Pine Creek Inlier (Table 4.11), and which were later reactivated and emplaced into the Pine Creek Inlier rocks as mantled gneiss domes, have been assigned to the West Australian Orogenic Province (Fig. 4.1). I n the west, the Rum Jungle and Waterhouse Complexes form domes, about 15 km across, of coarse, medium, and porphyritic granite, lesser aplitic and garnetiferous granite, some migmatite, gneiss, and schist, and minor quartzite, banded ironstone, metadiorite, and acid volcanics, in easterly striking bands which are discordant t o the adjacent, concentrically disposed, Early Proterozoic metasediments. Rhodes (1965) showed that the Rum Jungle Complex was uplifted as a mantled gneiss dome during regional folding of the overlying Early Proterozoic succession. Stephansson and Johnson (1976) suggested that the Rum Jungle and Waterhouse Complexes were diapirically emplaced during intrusion of late Early Proterozoic (post-tectonic) granites, although no granites of this age have been identified from the complex. Rhodes’ interpretation is more in accord with data from the Nanambu Complex. The Litchfield Complex may contain some Archaean elements, but this remains to be confirmed by further work; it almost certainly contains younger elements as well. To the east, in the Nanambu Complex, the oldest rocks exposed are 2470 Ma-old leucogranites. Coarse pegmatitic leucogneiss and interlayered quartzite and schist in the Complex are interpreted as Early Proterozoic psammitic sediments, which were accreted t o the Archaean granite mass during regional folding, doming, and amphibolite-grade metamorphism, at about 1800 Ma. In the nearby South Alligator Trough the Stag Creek Volcanics, which were interpreted as Archaean basement by Walpole et al. (1968), are now shown to be interbedded with the Early Proterozoic Namoona Group (Fig. 4.8) (Foy and Miezitis, 1977). Massive granites which intrude the metamorphic rocks of the Rum Jungle and Waterhouse Complexes broadly conform t o a Rb-Sr isochron about 2400 Ma old (Page, 1976a), supported by U/Pb zircon discordia about 2450 Ma (Page, 1976b), but stratigraphically older gneisses from the Waterhouse Complex yield a 1755 ? 26 Ma isochron with high initial 87Sr/86Sr ratio, related to c. 1800 Ma reactivation (Page, 1976a). A crystallization age for leucogranite in the Nanambu Complex is determined by a Rb-Sr isochron of 2470 ?I. 32 Ma (Page et al., 1980), and by similar (2400 ?I. 50 Ma) zircon
221 ages (Page, 1976b). Other rocks from the complex yield secondary isochron ages between 1750 Ma and 1980 Ma, reflecting incomplete isotopic homogenization during the c. 1800 Ma regional metamorphism (Page et al., 1980); 2000 Ma (cf. Page, 1974; Plumb, 197913) is now not recognized as a discrete event. EARLY PROTEROZOIC OROGENIC DOMAINS
McClintock Subprovince Halls Creek Inlier (Principal references: Dow and Gemuts, 1 9 6 9 ; Gemuts, 1 9 7 1 ; Gellatly et al., 1 9 7 4 )
The most commonly exposed domain of the McClintock Subprovince is the Halls Creek Inlier, now exposed over about 45,000 km2 within the zones of repeated younger deformation, the Halls Creek Inlier and King Leopold Mobile Zones. A thick Early Proterozoic geosynclinal-facies succession (Halls Creek Group) was deposited, and then these rocks were intruded by a variety of igneous rocks and converted to high-grade metamorphic rocks, collectively known as the Lamboo Complex. Basic intrusive rocks, which were emplaced either just before or during intense polyphase folding and metamorphism, and the high-grade metamorphic rocks and syntectonic granites, together comprise the Early Lamboo Complex. The subsequent transitional tectonism is represented by the Late Lamboo Complex: the late- or post-orogenic acid volcanics and cogenetic discordant granites which unconformably overlie or intrude the Early Lamboo Complex. These rocks are all unconformably overlain by platform covers of the Early Proterozoic Kimberley Basin and younger basins. The Halls Creek Goldfield was the first goldfield discovered in Western Australia (in 1884) but production waned rapidly before 1900. The area is now the centre of active base-metal exploration, mostly for copper, zinc, and nickel . The stratigraphy of the Halls Creek Inlier is summarized in Table 4.1, and the Proterozoic evolution of the Kimberley region is summarized in Fig. 4.6.
Tectonic setting. The Halls Creek Inlier is only exposed within the Halls Creek (East Kimberley) and King Leopold (West Kimberley) Mobile Zones. No basement is exposed beneath the Halls Creek Group. The Mobile Zones possess intense linear gravity anomalies, typical of mobile belts surrounding older cratons, and the most intense anomalies are spatially related t o the zones of most intense metamorphism and basicultrabasic intrusives within the Inlier. These anomalies, and offshore structural data, show that the high-grade metamorphic rocks exposed within the Halls Creek Mobile Zone may be traced northwards beneath the sea, while those exposed within the King Leopold Mobile Zone continue southwards
to to
to
600
1000
", z
L?
a Y
w 0
w
m I400 rn
a < w
0 Lc
v1
0 z J
=' s
moo
w
4 0
2000
2200
Fig. 4.6. Diagrammatic relation of rock units and summary of geological history, Kimberley region. (After Plumb and Gemuts, 1976, with permission of International Geological Congress.) NB: Isotopic ages here are original data (1.39 X lo-" yr-' decay constant of 87Rb), and are c . 2% older than recalculated ages in text.
223 TABLE 4.1 Summary of Early Proterozoic Stratigraphy, Halls Creek Inlier Unit
Rock types and thickness
Remarks
LAMBOO COMPLEX Latc L a m b o o C o m p l e x (transitional domains -late o r post-orogenic igneous actiuity) Bow River, Lennard, etc. Coarse granite, adamellite; Discordant plutons. Intrude Granites s o m e granodiorite, tonalite Whitewater Volc. and older units Castlereagh Hill, Bickleys, etc. Porphyries
Porphyritic microgranite, quartz-feldspar porphyry
Whitewater Volcanics
Rhyolitic to andcsitic ashflow tuff and lava. Minor tuff, agglomerate. U p t o 12.000 m
Comagmatic. Intrude o r unconformable o n Halls I
Gp.3
k'arly L a m h o o C o m p l e x (orogenic domains - intensely d e f o r m e d and metamorphosed gcosynclinal r o c k s ) East Kimberly. Anatectic Mabel Downs Granodiorite Foliated granodiorite, equivalent of Tickalara tonalite Mets. Tickalara Metamorphics
Schist, paragneiss, granulite, amphibolite, calcsilicates
East Kimberly. High-grade equivalent of Halls Cr. Gp.
McIntosh Gabbro
Gabbro, troctolite, dolerite Pyroxenite, peridotite, anorthosite, schist
East Kimberly. Genetically related. Intrude Halls Cr. Gp. Antedate Tickalara metamorphism
Dykes and sills of uralitized dolerite
Intrudes, folded and metamorphosed with Halls Cr. GP.
Kongorow Granite
Foliated granite, granodiorite; locally gneissic
Complex. Partly anatectic; partly intrudes Lennard Gr.
Womharella Quartz Gabhro
Orthopyroxene quartz gabhro, norite and tonalite
Intrudes Kongorow Gr. Intruded by Kongorow Gr., Lennard Gr.
Little Gold River Porphyry
Orthopyroxene feldspar porphyry
Related t o Womb. Qtz. Gahh.? Intrudes Whitewater Volc.
Alice Downs Ultrahasics Woodward Dolerite
l t g c uncertain
HALLS CREEK GROUP (geosynclinal s e d i m e n t s ) Undivided Phyllitic siltstone and greywacke, schist. Some acid volcanics. U p to 6000 m
Tentatively correlated with Olympio F m . Unconformably overlain by Whitewater Volc.
224 TABLE 4.1 (continued) Unit
Rock types and thickness
Remarks
HALLS CREEK GROUP (continued) Olympio Formation
Subgreywacke, siltstone; minor dolomite, acid volcanics. 3500 m+
Turbidites. Unconformably overlain by Whitewater Volc.
Biscay Formation
Tholeiitic basalt; tuffaceous greywacke, siltstone; carbonate rocks. 1503000 m
Pillow lava. Prominent carbonate marker beds
Saunders Creek Formation
Quartz sandstone, conglomerate. 0-200 m
Fanglomerate deposit. Possible hiatus at base
Ding Dong Downs Volcanics
Amygdaloidal basalt, tuff, tuffaceous greywacke. Minor acid volcanics. 300 m+
Acid rocks may be intrusive o r extrusive
beneath Palaeozoic cover (Fig. 4.4). Deformation related t o the King Leopold Mobile Zone may be extrapolated southeastwards t o the GranitesTanami region, and ultimately into the Arunta Inlier. Although the major structures of the Mobile Zones appear t o have originated from at least the time of the Tickalara metamorphism, their earlier influence on the Halls Creek Inlier rocks remains enigmatic. Gellatly (1971) speculated that the geosyncline within which the Halls Creek Group was deposited was confined essentially to the Halls Creek and King Leopold Mobile Zones, adjacent t o an Archaean craton beneath the Kimberley Basin (Kimberley Block). However, the alternative, that the Halls Creek Group underlies most of the Kimberley Block, cannot be ruled out, and the linear mobile zones may have simply been the loci of the most intense igneous and metamorphic activity (Plumb and Gemuts, 1976). Age. Deposition of the Halls Creek Group commenced after about 2150 Ma (Page, 1976c), and ceased before the main (Tickalara) metamorphism, dated as 1920 27 Ma (Rb-Sr, total-rock isochron; Bofinger, 1967). The Whitewater Volcanics, of the Late Lamboo Complex, have yielded a rather imprecise Rb-Sr isochron of 1900 ? 110 Ma (Bennett and Gellatly, 1970), while the post-tectonic granites which intrude them are better dated at about 1840-1815 Ma old (various Rb-Sr isochrons; Bofinger, 1967; Bennett and Gellatly, 1970). +_
Sedimentation. The Halls Creek Group is more than 7000 m thick (Table 4.1). The volcanic units low in the succession (Ding Dong Downs Volcanics,
225
Biscay Formation) may represent an ancient island arc or eugeosynclinal facies, and these pass up into a thick monotonous turbidite sequence (Olympio Formation). A possible hiatus at the base of the Saunders Creek Formation (above Ding Dong Downs Volcanics) cannot be confirmed by an angular discordance. Gold mineralization is associated with dolerite dykes intruding the Biscay Formation, and volcanogenic copper mineralization occurs in basal turbidities of the Olympio Formation. Early Lam boo Complex. The basic intrusive rocks are probably genetically related. Long, narrow basic dykes and sills (Woodward Dolerite) are tightly folded with the Halls Creek Group. In the East Kimberley, competent broadly warped, differentiated layered basic-ultrabasic sills (Alice Downs Ultrabasics, McIntosh Gabbro), many hundred of metres thick and up to 30 km across, are metamorphosed at the margins, but may postdate the earliest deformation of the Halls Creek Group. The age relationship of the Wombarella Quartz Gabbro and petrographically similar Little Gold River Porphyry, in the West Kimberley, is obscure; they could include both syntectonic and post-tectonic members. The Halls Creek Group rocks are invariably tightly to isoclinally folded and are generally metamorphosed t o low-greenschist facies. In the East Kimberley the higher-grade metamorphic equivalents are known as the Tickalara Metamorphics, but in the West Kimberley all metasediments, up t o amphibolite facies, are still mapped as Halls Creek Group. The syntectonic Mabel Downs Granodiorite is believed to have formed by partial melting of the Tickalara Metamorphics, but the equivalent unit in the West Kimberley (Kongorow Granite) is a complex body, including both anatectic gneiss and elements of the Late Lamboo Complex. There appears to be a spatial relationship between the area of higher grade metamorphism and the abundance of basic intrusives. In the East Kimberley, Gemuts (1971) has subdivided the Halls Creek Group and Tickalara Metamorphics into three metamorphic zones, ranging from low-greenschist facies up to granulite facies, in pelitic and arenaceous, calcareous, and basic-igneous rocks. The metamorphic rocks belong to the low pressure-intermediate facies series. Superposed metamorphisms have been identified locally (Nevill, 1974), but these have yet t o be identified throughout the region. The Halls Creek Group and Tickalara Metamorphics have been intensely deformed by at least two episodes of very tight and isoclinal folding; associated axial-plane cleavages of both episodes trend northnortheast to northeast, parallel to the Halls Creek Mobile Zone. The main “Tickalara” metamorphism coincided with the second episode, and was accompanied by movements on major faults. In the West Kimberley, the Halls Creek Group rocks mostly lie within the low-greenschist facies, but higher-grade rocks form a belt along the southwestern edge of the Inlier. Two distinct episodes of metamorphism are recog-
nized (Gellatly et al., 1974): an early low-pressure, moderately high-temperature metamorphism (Buchan type), and a later higher-pressure metamorphism (Barrovian type). At least three periods of intense folding predate sedimentation in the Kimberley Basin. North-northeast-trending cleavage (i.e. parallel to Halls Creek Mobile Zone) and open folds accompany the earliest (Buchan) metamorphism. A multiphase period of southeast-trending cleavage (i.e. parallel t o King Leopold Mobile Zone) accompanied the major (Barrovian) metamorphism and granitization. The Whitewater Volcanics and post-tectonic granites have extensive zones of southeasterly trending cleavage and mineral lineations (see below).
Late Lam boo Complex. The acid Whitewater Volcanics unconformably overlie the Early Lamboo Complex, and are closely associated with petrographically similar intrusive equivalents (e.g., Castelreagh Hill, Bickleys Porphyries). These are in turn intruded, over a considerable time span, by a wide variety of coarse-grained granites and adamellites in composite batholiths (e.g., Bow River, Lennard Granites). In the East Kimberley, the granites are massive and undeformed, and the Whitewater Volcanics are only folded into broad, open folds. However, the equivalent rocks in the West Kimberley were intensely sheared and folded in the King Leopold Mobile Zone, before the development of the Kimberley Basin (Fig. 4.6). It is possible that the major (Barrovian) metamorphism occurred in the King Leopold Mobile Zone at this time, rather than at the same time as the “Tickalara” metamorphism. Other domains In the Tennant Creek Inlier pelitic gneisses, which were metamorphosed to the amphibolite facies 1920 Ma ago (Black, 1977), and which possibly form basement to the low-grade Warramunga Group, are assigned to the McClintock Subprovince. Pelitic rocks in the eastern Arnhem Inlier were metamorphosed t o high-amphibolite and low-granulite facies about 1905 Ma ago (Plumb and Derrick, 1975); the age of post-tectonic granites and volcanics within this Inlier are still imprecisely known (McDougall et al., 1965), but they appear to be closest t o those of the transitional domain of the adjacent Pine Creek Inlier. Division 1 of the Arunta Inlier (Table 4.VII) has been tentatively correlated with the Halls Creek Group, and rocks of this age have also been inferred to underlie the Kalkadoon-Leichardt Block (Plumb, 1979b). McCulloch and Wasserburg (1978) have shown that sediments in the Mount Isa Orogen were derived from an approximately 2070 Ma-old terrain, presumably somewhere t o the west. These combined data all suggest that rocks belonging t o the McClintock Subprovince are widespread throughout the North Australian Orogenic Province, but only the Halls Creek Inlier was cratonized by transitional tectonism during this cycle.
227
Palmerston Su bprovince Pine Creek Inlier (Principal references: Walpole et al., 1968; Crick e t al., 1980; Needham et al., 1980; Needham and Roarty, 1980; Needham and Stuart-Smith, 1980; Stuart-Smith e t al., 1980)
The typical domain of the Palmerston Subprovince is the Pine Creek Inlier, an area of about 66,000 km’, t o the south and east of Darwin (Fig. 4.1), in which Early Proterozoic metasediments, granites, basic intrusives, and acid volcanics surround small basement inliers of Archaean rocks. The Inlier abuts the possibly Archaean t o Early Proterozoic Litchfield Complex in the west (the Chilling Platform, Fig. 4.7). The Inlier is unconformably overlain by the Carpentarian McArthur Basin and the Palaeozoic Daly River Basin t o the south, and by Mesozoic sediments of the Bathurst Terrace t o the North. About 1 4 km of Early Proterozoic continental, supratidal, and shallowmarine deposits, passing up t o more unstable and deeper-water turbidites at the top of the sequence, accumulated in an intracontinental, ensialic geosyncline. Tholeiitic dolerites were then emplaced into the sequence, followed by a period of intense polyphase folding, metamorphism, and migmatization, which reached a peak at about 1800 Ma ago. Subsequent transitional tectonism is represented by late- and post-orogenic discordant granites and mildly deformed, comagmatic acid volcanics, which intrude or unconformably overlie the metasediments of the orogenic domain. Finally, layered sills of tholeiitic dolerite were emplaced, after which the platform cover of the overlying McArthur Basin was deposited. The Inlier is of special economic importance because of the discovery, since 1970, of large and high-grade uranium deposits in the northeast -the “Alligator Rivers Uranium Field”. The worked-out Rum Jungle and South Alligator Valley Uranium Fields, in the west and southeast, respectively, were mined during the 1950’s-early 1960’s. Several small base-metal and numerous precious metal (mostly gold and tin) deposits have been mined, mostly during the late 19th century, but most have been mined out. The stratigraphy of the inlier is summarized in Table 4.11. The major tectonic elements are shown on Fig. 4.7, and the relationship of stratigraphic units in Fig. 4.8,
Tectonic setting. The Pine Creek Inlier is part of an intracontinental ensialic structure, in which Early Proterozoic rifting of Archaean basement formed a northerly trending trough, the South Alligator Trough, flanked by the Batchelor and Kakadu Shelves (Fig. 4.7). The Inlier lies t o the east of the Halls Creek Mobile Zone (Figs. 4.1 and 4.4): the Litchfield Complex, which the Inlier abuts t o the west, lies either within a subsidiary divergent arm of the Halls Creek Mobile Zone (Plumb, 1979b), or within the northern extension of the Mobile Zone itself (Traves, 1955; Rod, 1966). The geosynclinal
228
Fig. 4 . 7 . Major tectonic elements, Pine Creek Inlier. (After Stuart-Smith et al., 1980).
structure, of which the Inlier is part, probably continues beneath the McArthur Basin for a considerable distance t o the southeast (Fig. 4.4). Age. The metasediments of the Pine Creek Inlier unconformably overlie the c. 2450 Ma granites of the Rum Jungle and Nanambu Complexes, but sedimentation is not thought to have commenced until about 2200-2000 Ma ago, and the end of sedimentation is dated by 1870 t 20 Ma-old U/Pb zircon discordia in intrusive gneissic granites of the Nimbuwah Complex (Page et al., 1980). The peak of the subsequent period of metamorphism and migmatization is reflected in an 1803 t 20 Ma Rb-Sr total-rock isochron from these same and other gneisses in the Complex, and from other widespread Rb-Sr total-rock and mica isochrons and K-Ar mica ages around 1800 Ma, throughout the Kakadu Shelf (Page et al., 1980). Several of the late and post-orogenic granites of the transitional domain
C H ILL1NG PLATFORM +r
BATCHELOR SHELF
NANAMBU HIGH
S O U T H ALLIGATOR J-4- r TROUGH
KAKADU SHELF
A
7
I
I
I
FINNISS W L E Q S R O J F
?
Fig. 4.8. Diagrammatic relation of rock units, Pine Creek Inlier. (Adapted from Needham et al., 1980).
I
230 TABLE 4.11 Summary of Early Proterozoic Stratigraphy, Pine Creek Inlier Unit
Rock types and thickness
Remarks
Transitional domains - Late or post-orogenic igneous activity Oenpelli Dolerite Layered tholeiitic dolerite Undeformed lopoliths Kakadu Shelf. About 1688 and gabbro Ma Edith River Volcanics
Acid lava, tuff, ignimbrite; minor basalt, breccia, sandstone, conglomerate. Up to 1200 m
Unconformable on metasediments. Comagmatic with granites. About 1720 Ma
Cullen, Jim Jim, etc. Granites
Biotite granite, adamellite, syenite, granodiorite
Numerous mainly discordant plutons. About 1780-1720 Ma
Orogenict domains -- Intensely deformed and metamorphosed geosynclinal rocks Most intense metamorphism Granitoid migmatite, Nimbuwah Complex and anatexis of metasedigranite, gneiss, schist ments - NE Kakadu Shelf. Regional met. about 1800 Ma Zamu Dolerite
Geosynclinal sediments FINNISS RIVER GROUP
Layered continental tholeiitic dolerite sills, minor dykes
Intruded before folding and metamorphism
Siltstone, shale, greywacke, slate, arkose, quartzite, sandstone. 1500-5000 m
Widespread turbidite (“flysch”) facies. Unconformable on underlying units
SOUTH ALLIGATOR GROUP Pyritic black shale, chertyhematitic siltstone; stromatolitic dolomite, banded iron formation, tuff, jaspilite. Up t o 5000 m
Shallow-water fine terrigenous-carbonate facies. Thickest in South Alligator Trough. Unconformable on units below. Sth. Alligator U-Au
MOUNT PARTRIDGE GROUP Sandstone, siltstone, arkose, shale, conglomerate, quartz greywacke. Up to 5000 m
W-thinning fluviatile wedge - Kakadu Shelf. Mostly unconformable on units below
NAMOONA GROUP
Pyritic-carbonaceous calcareous) shale, limestone, lithic greywacke, calcareous sandstone, tuff, agglomerate. Up to 3000 m
Widespread, E-thinning, turbidite-carbonate facies, Conformable with units below. Rum Jung., Alligator R. U
KAKADU GROUP
Sandstone, arkose, siltstone, conglomerate, quartzite. Up t o 6000 m
Basal fluviatile sequence Kakadu Shelf. Unconformable on Nanambu Cplx. (basement)
231 TABLE 4.11 (continued) ____
Unit.
Rock types and thickness
Remarks
BATCHELOR GROUP
Conglomerate, arkose, siltstone, sandstone, stromatolitic dolomite, magnesite. Up t o 1500 m
Basal fluviatile sequence Batchelor Shelf. Unconformable on Rum Jungle and Waterhouse Cplx. Equivalent t o Kakadu Gp.
Porphyritic adamellite, biotite-muscovite granite, migmatite, gneiss, schist, meta-diorite
Granites emplaced 25002400 Ma ago; reactivated 1800 Ma
Arclinean basement inliers Rum Jungle, Waterhouse, and Nanambu Complexes
(Cullen Granite, etc.), and the coeval acid volcanic rocks (Edith River Volcanics), fit a combined Rb-Sr isochron about 1720 Ma-old (Compston and Arriens, 1968; Walpole et al., 1968); more recently, 1780-1760 Ma Rb-Sr isochrons have been determined for these same or related granites (Page et al., 1980: Riley, 1980). Finally, the layered sills of Oenpelli Dolerite, emplaced before deposition commenced in the McArthur Basin, is dated by a 1688 1 3 Ma-old Rb-Sr isochron (Page et al., 1980).
*
Sedimentation. The metasediments in the Pine Creek Inlier were shale and siltstone (mainly carbonaceous), sandstone, quartzite, and subordinate greywacke and carbonate. These rocks and minor volcanics are all intruded by dolerite sills. The units form a dominantly layered-stratigraphic sequence (Fig. 4.8) (Crick et al., 1978), in contrast to the laterally interfingering facies model of Walpole et al. (1968). A basal fluviatile sequence (Batchelor, Kakadu Groups) is unconformable on Archaean basement. Interbedded carbonates in the west (Batchelor Group) contain massive supratidal evaporite deposits, now replaced by carbonate (Crick and Muir, 1980). These beds are succeeded by a thicker, fine turbidite-carbonate facies (Namoona Group), which was deposited over most of the Batchelor Shelf and South Alligator Trough, and thins onto parts of the Kakadu Shelf; further massive evaporites, replaced by carbonate, are found in the lower Cahill Formation (of the Namoona Group). This stage closed with minor submarine basic to intermediate volcanism (Stag Creek Volcanics) in the South Alligator Trough. Movement along the South Alligator Hinge Zone, after this volcanism, rejuvenated source areas to give rise t o a wedge of quartz-rich, fluviatile psammites and rudites (Mount Partridge Group) over the Kakadu Shelf, thinning westwards across the South Alligator Trough and Batchelor Shelf.
232 Mild tectonism caused folding, uplift, and peneplanation, which was followed by a return t o the deposition of a shallow-marine t o supratidal, fine terrigenous-carbonate-evaporite facies (South Alligator Group) in the South Alligator Trough; these facies thin westwards onto the Batchelor Shelf, where some units wedge out in the Rum Jungle area. An episode of subaerial felsic volcanism was centred in the South Alligator Trough at this time. Subsequent faulting and volcanism in the west marked a shift in tectonism from the South Alligator Trough t o the Western Fault Zone. Tectonic reactivation produced an eastward-thickening sheet of turbidite (“flysch”) facies (Finniss River Group), which extended across the Batchelor Shelf and into the South Alligator Trough. Sandstone was then deposited in the Western Fault Zone, onlapping onto Archaean basement in the Chilling Platform.
Igneous intrusion, folding, and metamorphism. Continental tholeiitic dolerites (Zamu Dolerite) were emplaced predominantly into the South Alligator and Finniss River Groups after sedimentation ceased, mostly as conformable sills which were later folded and metamorphosed during regional orogenesis; most are now amphibolites (Ferguson and Needham, 1978). Four phases of moderate t o tight folding, including two isoclinal folding events, have been recorded in the metasediments of the Pine Creek Inlier (Needham and Stuart-Smith, 1976). Tight t o isoclinal folds with steep limbs, in the central area, grade through overturned tight t o isoclinal folds across the South Alligator Trough, into very complex folding over the Kakadu Shelf. An eastward increase in metamorphic grade is also apparent. Metamorphism to lower greenschist facies was coincident with folding in the western and central parts of the inlier, but reached amphibolite facies of Barrovian type (Smart et al., 1975; Ferguson, 1980) during the final (1800 Ma) fold phase on the Kakadu Shelf, Where the metasediments were migmatized to form the gneissic outer parts of the Nimbuwah Complex, or accreted as gneiss onto Archaean basement around the Nanambu Complex mantled gneiss dome. Two-pyroxene tonalite is evident locally within the Nimbuwah Complex (Ferguson, 1980), this may, however, be pristine igneous material and need not be interpreted as the high-grade end member of the 1800 Ma metamorphism. 1870 Ma-old zircons in the granitoid inner parts of the Nimbuwah Complex indicate some syntectonic granitoid emplacement, before the peak of the 1800 Ma-old regional metamorphism. Transitional domains. Numerous, mostly discordant, plutons of massive biotite granite, adamellite, and minor syenite and granodiorite were emplaced about 1780-1760 Ma ago, after the polyphase folding and regional metamorphism; in the lessdeformed and metamorphosed areas, these bodies domed (Stephansson and Johnson, 1976) and hornfelsed the surrounding
233 metasediments. The granites are mostly subsolvus, I-types (Ferguson et al., 1980). Typical is the Cullen Granite, in the middle of the inlier, in which several plutons hav xcreted t o form the largest of the composite batholiths. The mildly deformed Edith River Volcanics, which are mainly valley-fill rhyolite, dacite, and pyroclastics, are associated with hypersolvus granites. The Volcaiiics were regarded by Walpole et al. (1968) as the basal unit of the ovc-lying McArthur Basin succession. Structural relationships with the overlying platform-cover rocks range from paraconformable to highly unconformable. A more recent tectonic analysis (GSA, 1971) indicates that they are more properly related t o the final cratonization of the Early Proterozoic orogenic domain. The youngest event recognized in the Pine Creek Inlier, before platformcover deposition commenced in the McArthur Basin, is the intrusion of the 1688 Ma-old Oenpelli Dolerite. Extensive layered lopoliths of continental tholeiite were emplaced about 1-2 km below the top of the enclosing metasediments (Stuart-Smith and Ferguson, 1978), then were exposed by erosion to form basement highs, before the basal beds of the unconformably overlying Mc Arthur Basin succession were deposited. Post-tectonic phonolite dyke swarms (c. 1320 Ma) and dolerite dykes (c. 1200 Ma) were emplaced into the Pine Creek Inlier, at about the same time as dolerite sills were intruded into the overlying McArthur Basin.
Mineral deposits. About 2500 tonnes of U 3 0 s have been produced from the Rum Jungle and South Alligator fields, while present measured reserves of the “Alligator Rivers Uranium Field” total 250,000 tonnes of U30s. The U and U-Au deposits of the Alligator Rivers and Rum Jungle fields are stratabound in brecciated carbonaceous pelitic rocks, associated with carbonate rocks after massive evaporites, in the correlative Masson and lower Cahill Formations (Fig. 4.8), and are situated immediately adjacent to basement domes of the Rum Jungle and Nanambu Complexes. A long history of epigenetic concentration followed probable original sedimentary concentration of the metals. The South Alligator Valley U and U Au deposits are-mainly stratabound in carbonaceous shales of the Koolpin Formation, where they are in contact with the Edith River Volcanics, and they contain minor Cu. Small stratabound base metal deposits (Ag-Pb-Zn, Cu), such as Woodcutters (Roberts, 1973), occur at the same stratigraphic level of the Masson Formation as the basement-associated uranium deposits, but away from the basement outcrops; both volcanic exhalative and biogenic processes are indicated in their origin (T.H. Donnelly and W.M.B. Roberts, pers. commun., 1979). Other small stratabound deposits (Ag-Pb-Zn, Cu, Au) are associated with a distinctive suite of carbonate, iron-rich, pelitic, and volcanic rocks in the South Alligator Group (Gerowie Tuff, Kapalga Formation; Fig. 4.8). Similar iron-enriched beds are hosts to supergene Fe deposits (Frances Creek: Wildman Siltstone; Mount Bundey : Koolpin Formation; Fig. 4.8).
234 Numerous small hydrothermal deposits of Au, Sn, Cu, Ta, and W deposits are concentrated close t o post-tectonic granite plutons.
Other domains In the Murphy Znlier, a synorogenic phase of the Nicholson Granite Complex was emplaced about 1820 Ma ago, towards the close of folding and metamorphism of pelitic and quartzo-feldspathic sediments of the Murphy Metamorphics (Gardner, 1978). Post-tectonic phases of the granite were emplaced during the subsequent transitional tectonism, together with the cogenetic 1730 Ma-old Cliffdale Volcanics (Webb, 1973; Mitchell, 1976). The Tennant Creek Znlier (Crohn and Oldershaw, 1965; Dunnet and Harding, 1967; Crohn, 1975) contains a metasedimentary sequence of greywacke, siltstone, and acid volcanics (Warramunga Group). This sequence was deposited possibly after the 1920 Ma-old metamorphism (which is provisionally assigned to the McClintock Subprovince) and the sequence is inferred to have undergone its major deformation about 1810 Ma ago (Black, 1977). Subsequent transitional tectonism is represented by arenites, lutites, and basic and acid volcanics (Hatches Creek Group) in the Davenport Znlier, to the south (Smith, 1964, 1970), and by the emplacement of 18001500 Ma-old granites into both inliers (Black, 1977); the younger (16001500 Ma) granites apparently reflect events in the adjacent Arunta Inlier. The Warramunga and Hatches Creek Groups appear t o merge southwards into Divisions2 and 3, respectively, of the Northern Zone of the Arunta Znlier. The Division 2/Warramunga Group rocks continue westwards, through the Arunta Inlier, into the metasedimentary and metavolcanic Tanami Complex of The Granites-Tanami Znlier (Blake et al., 1975, 1979), which Plumb (197913) has assigned to the Palmerstone Subprovince (Blake et al. have, alternatively, correlated the Tanami Complex with the Halls Creek Group). Intense superposed folding of the Tanami Complex, about axes parallel to the nearby Halls Creek and King Leopold Mobile Zones, was followed by transitional tectonism : 1770 Ma-old acid volcanics (Mount Winnecke Formation) and some 1765-1685 Ma-old granite bodies (Page et al., 1976). Gravity trends suggest that the Murphy, Tennant Creek, northern Arunta, and The Granites-Tanami Inliers may be part of a single orogenic belt (Fig. 4.4). Mineral deposits associated with these orogenic domains include the CuAu-Bi-bearing quartz-hematite and quartz-magnetite lodes in production at Tennant Creek, and small Au-bearing quartz veins at The Granites and Tanami. Transitional domain associations are W deposits in the Davenport Inlier, and small Cu, Sn, and W deposits associated with the late-tectonic granites, and U in shears in acid volcanics, in the Murphy Inlier.
235 EARLY TO MID-PROTEROZOIC PLATFORM COVERS
Kimherley Basin (Principal references: Gellatly et al., 1970, 1975; Plumb and Gemuts, 1976)
The Kimberley Basin is a relatively undisturbed structural basin, covering 160,000 km2 in area, which overlies the Kimberley Block. It is flanked by the Halls Creek and King Leopold Mobile Zones (Fig. 4.1), from which most of the basin succession has been eroded. Up t o 5 km of quartz-rich arenite, lutite, flood basalt, and minor carbonate rocks were deposited on top of the Halls Creek Inlier rocks, after the 1840-1815 Ma-old transitional domain rocks had been folded and eroded. The upper age limit t o the basin is provided by the 1760 25 Ma-old Hart Dolerite (Rb-Sr, total-rock isochron; Bofinger, 1967), which intrudes the whole of the basin succession. The basin was mildly deformed several times during the Proterozoic and Palaeozoic, and deformation increases suddenly adjacent to, and within, the Halls Creek and King Leopold Mobile Zones. The stratigraphy is summarized in Table 4.111 and Fig. 4.9.
*
Stratigraphy The basin succession is divided into three main groups: the Speewah Group, the generally conformably overlying Kimberley Group, and then the conformably overlying Bastion Group (or the stratigraphically equivalent Crowhurst Group). The Revolver Creek and Moola Bulla Formations within the Halls Creek Mobile Zone, and the Red Rock Beds flanking the eastern edge of the Mobile Zone, are correlated with all or part of the Kimberley Basin succession. The intrusive Hart Dolerite, Fish Hole Dolerite; and Wotjulum Porphyry are probable correlatives of each other. A planar unconformity has been identified throughout the basin in the middle of the King Leopold Sandstone, at a point marked by a sudden change in palaeocurrent trends (Plumb and Gemuts, 1976); the “lower” King Leopold Sandstone is described with the Speewah Group in this discussion (Fig. 4.9). Pa laeogeograp h ic e uo lu tion A feature of the basin is its uniform stratigraphy in the central part (Fig. 4.9) and several indicators suggest that the extent of the original sedimentary basin may not have been appreciably greater than the presently exposed structural basin. Facies changes accompany the onlap of units onto the Halls Creek Mobile Zone. The Speewah Group thins out across the King Leopold Mobile Zone, but the overlying Kimberley Group transgresses it with little change, until covered with Phanerozoic rocks farther south. Speewah Group. The Speewah Group is only exposed along the upturned southeastern and southwestern margins of the basin; its extent beneath the
236 TABLE 4.111 Summary o f Early Proterozoic stratigraphy, Kimberley Basin and Equivalents ______ Unit
Rock types and thickness
Remarks
Hart Dolerite
Tholeiitic dolerite, gabbro, granophyre. Up t o 3000 m
Sills intrude Kimberley and particularly, Speewah Gps.
Wotjulum Porphyry
Quartz-feldspar Up t p 6 0 0 m
Intrudes Kimberley Gp. in west. Same age as Hart Dol.
Fish Hole Dolerite
Epidotized dolerite. Up t o 900 m
Sills intrude Red Rock Beds
CROWHURST GROUP
Purple and green shale; siltstone; dolomite; sandstone. 1 6 0 m +
Equivalent t o Bastion Gp. Stromatolites, m u d cracks ripple marks, carbonate nodules
BASTION GROUP
Green shale, siltstone; alternates with quartz sands t o n e ; minor dolomite. 1 3 5 0 m+
Conformable o n Kimberley Gp. Cross-beds, ripple marks, mud cracks, carbonate nodules
KIMBERLEY GROUP
Quartz and feldspathic sandstone; alternates with red siltstone, tholeiitic basalt; minor dolomite, glauconitic sandstone. 1000-3500 m
SPEEWAH GROUP
Feldspathic and quartz sandstone, arkose, and granule sandstone; interbedded micaceous and chloritic siltstone, shale; minor ashflow tuff. Up t o 1700 m
Abundant cross-beds, ripple marks, shale pellets; some stromatolites. Laterally consistent units. Palaeocurrents from north. Unconformable o n Speewah GP. Abundant cross-beds, ripple marks, m u d cracks, shale pellets. Only exposed southern edge of basin. Thins t o southwest. Palaeocurrents from northeast. Unconformable o n “Late L a m b o o Cplx.”
Revolver Creek Formation
Basalt, quartz sandstone, siltstone. Up t o 1200 m
Unconformable between “Late L a m b o o Cplx.” and Carr Boyd Gp. Equivalent to Kimberley Gp.
Moola Bulla Formation
Arkose, greywacke, conglomerate, siltstone. 3000 m +
Unconformable between Halls Cr. Gp. and Kimberley Gp. Probably equivalent to Speewah Gp.
Red Rock Beds
Quartz sandstone, red siltstone, conglomerate. 2000 m +
Unconformable between Halls Cr. Gp. and Mt. Parker Sandstone. Equivalent to Speewah o r Kimberley Gps.
_____
porphyry.
I
p
Pentecost Sandstone
e
Elgee Siltstone
1
Teronis Member
' w Warton Sandstone
_-. Conglornerote
* - ~
k
Undivided
Bedded hemotite K i n g Leapold Sandstone
L u m o n Siltstone
H'
k
Lansdowne Arkose
lveroge poloeocurrent
rop of cofurnnor section eroded Bore of co/urnnor section not exposed
Valentine Siltstone Tunganary Formation
Kimberley Group
-
Fault
O'Donnell Formation Undivided Speswah Group
,--.
Unconformity
Fig. 4.9. Diagrammatic stratigraphy and palaeocurrents, Kimberley Basin (Adapted from Plumb and Gemuts, 1976).
238
main part of the basin is unknown. Red feldspathic arenites are a feature of the sequence. The group thins in the southeast and west, and in the Halls Creek Mobile Zone the Kimberley Group lies directly on basement. Palaeocurrents are predominantly from the northeast and east, with reversals in places (Fig. 4.9). The group is interpreted as a broad transgressive-regressive cycle, with sediment derived from an elevated and tectonically active Halls Creek Mobile Zone: fluvial sands (“lower” O’Donnell Formation) pass up through alternating or interfingering fluvial and shallow-marine facies (“upper” O’Donnell Formation to Luman Siltstone), which then pass back into fluvial sands (“lower” King Leopold Sandstone). The Red Rock Beds and Moola Boola Formation are fluvial deposits, locally derived from and deposited on the other side of the Halls Creek Mobile Zone.
Kirnberley Group. The characteristic feature of the Kimberley Group is the maturity and lateral uniformity of most formations. Contacts between successive formations are usually gradational. The Bastion and Crowhurst Groups follow the same depositional patterns. Mature, laterally persistent, quartz sandstone units alternate with tholeiitic basalt (Carson Volcanics), massively bedded and deep-red siltstone (Elgee Siltstone), or green shale (Wyndham Shale). Palaeocurrents are remarkably consistent from the north to northwest, except for some changes in trend and reversals in the west, where local unconformities are found (Fig. 4.9). Sections thin markedly in the southeast, northeast, and west, and thicken locally in the west. It is interpreted that, after erosion of the Speewah Group, the Kimberley Basin took the form of a broad semi-enclosed, shallow-marine basin, which persisted throughout Kimberley and Bastion Group times. Shores t o the west, and perhaps northeast, supplied sediment for dispersal by strong unidirectional, perhaps longshore currents flowing towards the south-southeast. Sediments lapped onto downcurrent highs in the Halls Creek Mobile Zone and, perhaps, in the King Leopold Mobile Zone. A basin similar to the present North Sea is envisaged. Hart Dolerite The Hart Dolerite underlies the whole 160,000 km2 of the Kimberley Basin and, where now upturned along the Halls Creek and King Leopold Mobile Zones, a composite thickness of up to 3000 m of dolerite is exposed in several sills. It mostly intrudes the Speewah Group, where it encloses rafts of sedimentary rocks up t o several kilometres across. Compositions range from olivine dolerite, through quartz dolerite, to granophyre. Granophyre sheets up to 250 m thick cap the thickest sills. Compositions and differentiation trends are identical t o those of continental tholeiites. Structure Deformation of the Kimberley Basin rocks was in response to several peri-
239 ods of movement in the marginal mobile zones (Fig. 4.6). The central part of the basin is deformed into broad basins and domes, which reflect superposed movements on basement fractures parallel t o the adjacent mobile zones. Deformation in and adjacent t o the Halls Creek Mobile Zone was in response to north-northeast-trending left-lateral strike-slip displacements on numerous major faults (Fig. 4.5); individual horizontal displacements up t o 30 km and vertical displacements up t o 5 km have been demonstrated; associated thrust faults and open folds are oblique t o the major faults; slaty cleavage is developed locally, adjacent to and between major faults (Plumb, 1968; Dow and Gemuts, 1969). Deformation along the King Leopold Mobile Zone was by northwest-trending asymmetric t o overturned folds and moderately to steeply dipping reverse faults, parallel t o the axial planes of folds; overturning was mostly towards the northeast; Kimberley Basin rocks were locally metamorphosed up t o amphibolite facies about 600 Ma ago (Bennett and Gellatly, 1970; Sofoulis et al., 1971).
McArthur Basin (Principal references: Plumb and Brown, 1973; Plumb and Derrick, 1975; Walker et al., 1977; Plumb et al., 1980; Williams, 1980).
The McArthur Basin is exposed over an area of about 200,000 km’, and is the principal element of the North Australian Platform Cover (Fig. 4.1). The basin includes the type section of the Carpentarian, as redefined here (see Introduction). A feature of the succession is the vast thickness (12 km) of shallow marine, intertidal, supratidal, and fluvial deposits which have accumulated. Evaporitic-carbonate sequences comprise a significant component of the sequence, and very precise comparisons may be made with modem sedimentary environments (e.g., Muir, 1979a; Muir et al., 1980). The basin is well known for an abundance of stromatolites (e.g., Cloud and Semikhatov, 1969; Walter, 1972; Walter et al., in prep.), and an abundance and diversity of well-preserved microfossils (Muir, 1974, 1976; J.H. Oehler, 1977; J.H. Oehler and Logan, 1977; D.Z. Oehler, 1978; Peat et al., 1978). The rocks show a remarkable lack of detectable metamorphism, considering the thickness of section which has accumulated. The huge McArthur River Pb-Zn-Ag deposit is well known as a model for stratiform ore formation. The principal tectonic elements are shown in Fig. 4.10, and the stratigraphy is summarized in Table 4. IV and Fig. 4.11.
Tectonic setting The McArthur Basin lies near the eastern edge of the North Australian Craton, adjacent t o the penecontemporaneous continental margin belt - the Mount Isa Orogen (Figs. 4.4 and 4.10). The McArthur Basin, Lawn Hill Platform, and Mount Isa Orogen, all have similar stratigraphic successions and parallel evolutionary histories (Fig. 4.11). The basin is bounded by, and un-
240 ~~~
~
AFURA
SEA
~~~~~
~~
GULF
~
~
OF
CARPEN TA R I A
I 1 1
\
NICHOLSON BASIN
OUNT
ISA
1
L A W N HILL PLATFOR MI Gordon
Foul1
MI I s 0 F ~ u l t LEICHHARDT
3R"3
Fig. 4.1 0 . Principal tectonic elements, McArthur Basin-Mount Isa Orogen. (After Plumb et al., 1980, with permission of Geological Society of Australia, Queensland Division).
conformably overlies, the Early Proterozoic Pine Creek, Arnhem, and Murphy Inliers. The Murphy Inlier separates the McArthur Basin from the Lawn Hill Platform (Fig. 4.1),and formed a ridge throughout sedimentation. In the north, south, and east, the limits of the basin are covered by unconformably overlying covers of the Palaeozoic Arafura, Georgina, and Daly
241
Fig. 4.11. Inferred time relationship of principal rock units, McArthur Basin to Mount Isa Orogen. (After Plumb et al., 1980; with permission of Geological Society of Australia, Queensland Division.)
TABLE 4.IV Summary of Carpentarian (Mid-Proterozoic) stratigraphy, McArthur Basin Unit and distribution
Rock types and thickness
Remarks
Dolerite sills
Tholeiitic dolerite and gabbro
Intrude Roper, Mt. Rigg, Malay Rd Gps.
ROPER GROUP and equivalents
Blanket quartz arenites alternate with lutites
Regionally unconformable o n McArthur G p . and equivs
ROPER GROUP (Throughout basin)
Quartz sandstone; micaceous siltstone and (glauconitic) sandstone; black shale; some carbonate, ironstone. 500-5000 m
Abundant cross-beds, ripple marks, flute casts, mud cracks, shale pellets etc. Upward-fining cycles, red beds. Roper R. Fe
MALAY ROAD GROUP (Caledon Shelf)
Quartz sandstone; micaceous siltstone and (glauconitic) sandstone; black shale 1550 m+
Overlies Wilberforce Beds. Equivalent to lower Roper Gp; upper section concealed
M c A R T H U R GROUP and equivalents
Dominantly carbonate rocks
Regionally unconformable on Tawallah G p and equivs
Dolarenite and dolomitic sandstone, dololutite and dolomitic siltstone; red siltstone and sandstone; conglomerate. Stromatolites, oolites, evaporites. Volcanics locally. 5001150 m Dolomitic siltstone and sandstone, stromatolitic dololutite, chert, quartz sandstone, conglomerate. Evaporites. Up to 1500 m
Unconformable on Batten Sub-Gp. Gypsum, anhydrite, halite casts. Persistent stromatolite markers. Intraclasts, flake breccias, ripple lamination, cross-beds, mud cracks
McARTHUR GROUP Smythe Sst.-Balhirini Do1.-Dungaminnie Fm. (Batten Trough, Urapunga Tectonic Ridge)
Batten Subgroup (Batten Trough)
Barney Cr. Fm.-Reward DoI.- “Lower Lynott Fm.” and equivs (Batten Trough)
Dolomitic, tuffaceous, carbonaceous, and pyritic shales; dolarenite, dololutite, dolomite breccia. Stromatolites. 50-1000 m
Umbolooga Subgroup (Batten Trough, Bauhinia Shelf)
Stromatolitic dolomite; dololutite; red and green ferruginous and dolomitic siltstone; dolarenite and dolomitic sandstone; quartz sandstone. Abundant evaporites. Chert, oolites. 1100-2800 m
Unconformable on Umbolooga Sub-Gp. Gypsum, anhydrite casts. Ripple lamination, cross-lamination, lenticular bedding, mud cracks, teepee structures, etc. Upward-fining cycles Presently included in Umbolooga (Barn., Rew.) and Batten (“L.Lyn.”) Sub-Gps. Marked facies change: Turbidites near Emu Fault pass westwards into thin stromatolitic shelf facies. McArthur R. Zn-Pb. Unconformable on Tawallah Gp. Gypsum, anhydrite, halite, barite casts. Abundant and varied stromatolites. Intraclasts, flake breccias, ripple marks, mud cracks, crossbeds, etc.
TABLE 4.IV (continued) Unit and distribution
Rock types and thickness
Remarks
____-
McARTHUR GROUP (continued) Karns Dolomite
Stromatolitic dolomite; dolomitic siltstone, sandstone; chert. Evaporites. 150 m
Overlies dissected Mastert o n F m . Similar facies, probably equiv. t o (lower?) Umbolooga Sub-Gp.
HABGOOD GROUP (Northern Batten Trough)
Dolomitic siltstone and sandstone, dolomite, siltstone, sandstone. Stromatolites. 4 0 0 0 m
Northern equivalent of McArthur Gp. Base and t o p n o t exposed.
MOUNT RIGG GROUP (Arnhttm Shelf)
Stromatolitic dolomite; dolomitic siltstone and sandstone; quartz sandstone, conglomerate; chert, oolites. 700 m
Probably equivalent only t o Umbolooga Sub-Gp. Unconformable o n Katherine R . Gp. Bulman Pb-Zn
Mount Bonner SandstoneWilberforce Beds (Caledon She1f )
Micaceous dolomitic siltstone, shale, and fine sand stone; quartz sandstone, conglomerate. 1600 m
Section incompletely exposed. Unconformable between Malay R d . Gp. and basement
T A WALLAH GROUP and cquivalents
Alternating cycles o f quartz arenites and subordinate volcanics, carbonates, and lutites
R egio null y u n c o n f orm a bl c o n basemenf
TAWALLAH GROUP (Bauhinia Shelf, Batten Trough, Wearyan Shelf)
Quartz and feldspathic sandstone, conglomerate. Basic and minor acid volcanics; dolomite, dolomitic siltstone, micaceous siltstone. Some glauconite. 4000-5000 m
Cross-beds, ripple-marks; mud cracks, shale pellets, halite casts, stromatolites. Redbank Cu. Westmoreland U.
PARSONS RANGE GROUP (Batten Trough)
Quartz, feldspathic, and ferruginous sandstones, siltstone, dolomite. 6000 m
Northern equivalent of Tawallah Gp. Ubiquitous cross-beds, ripple marks
KATHERINE RIVER GROUP (Arnhem Shelf)
Quartz, feldspathic, and ferruginous sandstones, conglomerate; basic volcanics; dolomite, dolomitic siltstone, micaceous siltstone. Some glauconite. 1800-2700 m
Several unconformities in section. Cross-beds, ripple marks, shale pellets, stromatolites.
Groote Eylandt Beds (Caledon Shelf)
Quartz and argillaceous sandstone, shale, conglomerate. 9-600 m
Lateral equivalent of uncertain part o r all of Tawallah Gp.
Mount Reid Beds (Urapunga Tectonic Ridge)
Porphyritic rhyolite overlain by sandstone and conglomerate. 60 m
N o exposed base. Correlation uncertain. Volcanics could be part of Tawallah Gp. o r basement
( Wearyan Shelf)
244
River Basins, and the Mesozoic Carpentaria Basin, respectively. However, palaeogeographic trends indicate that a basin margin existed not very far east of the present exposed McArthur Basin rocks, suggesting that a land barrier existed between the McArthur Basin and the postulated northerly extension of the Mount Isa Orogen (Figs. 4.1 and 4.4). Its palaeogeographic setting was that of a linear marine embayment, comparable t o either the Persian Gulf or Red Sea (intracontinental marine embayments) or to the Gulf of California (parallel to coast). Stratigraphy Three major subdivisions are recognized within the McArthur Basin succession (Fig. 4.11, Table 4.IV). The Tawallah Group and equivalents consist of quartz-rich arenites and subordinate volcanics, carbonates, and lutites up to 6 km thick. They are overlain by a dominantly carbonate sequence -the McArthur Group and equivalents - up t o 5.5 km thick; on which rest with regional unconformity the Roper and Malay Road Groups, which consist of alternating quartz arenites and micaceous lutites up to 5 km thick. A regional break beneath the uppermost sandstone of the Masterton Formation (top of the Tawallah Group) and Fleming Sandstone (Parsons Range Group), extrapolated from the Mount Isa Orogen by Plumb and Sweet (1974), has recently been confirmed (Rod, 1978; Jackson, 1979). The Masterton and Fleming Sandstones are here interpreted as the start of the new sedimentary cycle which encompasses the overlying McArthur Group and its equivalents (e.g., Figs. 4.11, 4.12).
Age Age control within the McArthur Basin succession is sparse, and depends somewhat on correlations with the Mount Isa Orogen (Fig. 4.11). In the southeast, the base unconformably overlies the 1730 Ma-old Cliffdale Volcanics of the Murphy Inlier, while in the northwest, a thinner basal sequence unconformably overlies the 1688 Ma-old Oenpelli Dolerite of the Pine Creek Inlier; the base may be diachronous. The younger limit t o the succession is provided by a minimum age of 1280 Ma for dolerites intruding the Roper Group (McDougall et al., 1965). Very preliminary data from the Australian National University suggest that the Roper Group may be older than 1400 Ma (W. Compston, pers. commun., 1978). The top of the Tawallah Group is fixed by the Packsaddle Microgranite/Hobblechain Rhyolite Member; its 1575 Ma Rb-Sr age (Plumb and Sweet, 1974; Plumb and Derrick, 1975) is now superseded by a 1680 Ma U-Pb zircon age for the correlative Carters Bore Rhyolite/Fiery Creek Volcanics, at Mount Isa (Page, 1978). The Mount Isa Group, which correlates with part of the McArthur Group, is now thought to be between 1620 and 1670 Ma old (Page, 1978). The final deformation and regional metamorphism of the Mount Isa Orogen is dated as 1490-1460 Ma old (Page, 1978), but this may overlap deposition of the Roper Group, rather than predate it as previously thought (e.g., Fig. 4.10).
245 Tectonic elements A number of tectonic elements are recognized within the McArthur Basin (Fig. 4.10). The palaeogeography of the basin was dominated by the northerly-trending Batten Trough : a palaeotectonic feature, controlled by syndepositional faults, in which up t o 12 km of shallow-water sediments accumulated, compared to 1.5-4 km on the stable Arnhem, Bauhinia, Caledon, and Wearyan Shelves, either side. A very thin succession accumulated on the Urapunga Tectonic Ridge, which separates the Arnhem and Bauhinia Shelves, and may also have divided the Batten Trough into two separate elements (Plumb et al., 1980). The Batten Trough and Urapunga Tectonic Ridge were then deformed into the presently observed tectonic features, the Batten Fault Zone and Urapunga Fault Zone. Large vertical uplifts have reversed the Batten Trough into a horst-like feature, with basement rocks locally exposed within it. Deformation of the stable shelves is very mild. The Batten Trough was an intracratonic graben with a history of evolution very similar to that of an aulacogen (e.g., Hoffman, 1973). However, lack of outcrop obscures its spatial relationship to a continental margin: it lies parallel to the Mount Isa Orogen but other margins, possibly perpendicular to it, are not exposed. Palaeogeographic evolution The palaeogeographic evolution of the McArthur Basin (Plumb et al., 1980) is summarized in Fig. 4.12 and Table 4.IV. Tawallah Group and equivalents (Fig. 4.12A). The Batten Trough was clearly defined in the north, where the Parsons Range Group is very much thicker than the equivalent Katherine River Group and Groote Eylandt Beds, but in the south the Tawallah Group is of similar thickness across both the Batten Trough and Wearyan Shelf. Volcanic rocks are common on the Wearyan and Arnhem Shelves, but decrease significantly in the southern Batten Trough and are, surprisingly, absent from the fault-bounded Parsons Range Group. Two cycles are recognized. The “lower succession” on the marginal shelves comprises basal fluviatile arenites, derived from the southeast and northwest, respectively, passing up into flood basalt and marginal-marine stromatolitic carbonates. In the northwest, locally thickened sections were deposited in marginal depressions (Mount Callanan, Edith Falls Basins, Waterhouse Syncline). The basal fluviatile arenites grade into mainly shallow-marine arenites in the Batten Trough. The “upper succession” comprises sandstone, red siltstone, stromatolitic carbonate, and acid to basic volcanics, deposited in a complex of fluviatile, lagoonal, and marginal-marine environments. Unconformities are common on the Arnhem Shelf, and the succession is less extensive than the “lower succession”.
246 A
TAWALLAH
GROUP
A N D EQUIVALENTS
,
ion&,
6. UMBOLOOGA SUBGROUP A N D EQUIVALENTS
P
L BATTEN SUBGROUP A N D EQUIVALENTS BALBlRlNl DOLOMITE ETC A N D EQUIVALENTS 19
Fig. 4.1 2. Palaeogeographic evolution of the McArthur Basin. (After Plumb et al., 1980; with permission of Geological Society of Australia, Queensland Division).
247
Umbolooga Subgroup and equivalents (Fig. 4.12B). The Batten Trough is now clearly defined in both the north and south. Following a regional unconformity, a basal sheet sandstone (upper Masterton Formation and equivalents) passes up into a carbonate-evaporite-red silt and sand complex, with varied and abundant stromatolites and pseudomorphs after evaporites. Most lithostratigraphic units in the trough show striking continuity over more than 10,000 km2. Many units may contain significant components of reworked aeolian material. Alternating cycles of transgression and regression gave rise t o sequences of intertidal, shallow subtidal, and supratidal deposition, in a variety of restricted marginal-marine and lagoonal environments. A desert climate is indicated. Very close analogies may be made in parts of the sequence with the sabkhas of the Trucial Coast (Muir, 1979a). Successions on adjoining shelves are much thinner, but of similar facies, and probably only the lower part of the McArthur Group was deposited there. On the Wearyan Shelf the very thin Karns Dolomite was deposited in isolated lagoons on a highly eroded terrain, between hills of Masterton Formation up to 100 m high. Towards the end of Umbolooga Subgroup time, tectonism related to syndepositional movements on the Emu and other major faults radically changed the palaeogeographic framework. Locally thick sequences of Barney Creek Formation and Reward Dolomite accumulated in sub-basins separated by areas of much thinner sequences. Massive talus breccia, slump breccia, and turbidite fanned out from the Emu Fault into deposits of pyritic carbonaceous siltstone, tuff, and dolomite, passing eventually into more stable shelf sequences farther to the west. The huge McArthur River (H.Y.C.) and related stratiform Pb-Zn deposits accumulated in the Bulburra Depression, adjacent to the Emu Fault (Walker et al., 1978). Batten Subgroup and equivalents, and younger units (Fig. 4.12C). Following tectonism at the end of Umbolooga Subgroup time a marked shift in depositional pattern occurred ; the unconformably overlying Batten Subgroup and equivalents are entirely confined to the Batten Trough. Local unconformities occur within the subgroup near the basin margins. A flood of fine terrigenous and carbonate detritus, derived by erosion of the exposed shelves t o the east and west, was deposited again in mainly intertidal and supratidal environments, along with stromatolites and local evaporites, in a broadly regressive sequence. Very close analogies may be made between parts of the sequence and Holocene supratidal deposits accumulating in Spencers Gulf (K.A. Plumb, unpublished data) and the Coorong Lagoons (Muir et al., 1980). Wetter climates and greater groundwater discharge than the Umbolooga Subgroup is indicated. The Batten Subgroup is followed by regional unconformity. A basal lenticular sandstone and conglomerate (Smythe Sandstone and equivalents)
248 is followed by further subtidal to intertidal carbonate and terrigenous sediments, with stromatolites and evaporites (Balbirini Dolomite-Dungaminnie Formation and equivalents), deposited on a marginal marine shelf - a higherenergy environment than the Umbolooga Subgroup. Acid t o basic volcanics were extruded locally on the Urapunga Tectonic Ridge.
Roper and Malay R o a d Groups (Fig. 4.12D). Post-Mount Isa Group deformation in the Mount Isa Orogen is reflected in a regional unconformity throughout the McArthur Basin, and a change in the pattern of sedimentation. The Batten Trough ceased t o exist, and the thickest sequence was deposited on top of the Bauhinia Shelf, Hinge zones existed at the Hells Gate Hinge Line in the north, and the Emu Fault in the east. The sequences are characteristic of a sub-stable shelf association (or molasse facies?), and probably reflect the immediately preceding or contemporaneous tectonism in the Mount Isa Orogen. The succession is provisionally interpreted as indicating a series of transgressive cycles in which fluviatile sand, derived from a southerly source, passes up into marginalmarine silt and mud, with a distal carbonate shelf in the north. Banded pisolitic iron ores (Roper River Fe) (Cochrane and Edwards, 1960) were deposited in marginal marine environments. Associated black shales have yielded an important new assemblage of microfossils, and the organic matter indicates remarkably little or no thermal metamorphism (Peat et al., 1978). Dolerite sills Following deposition of the Roper Group, sills of tholeiitic dolerite were emplaced into the western part of the McArthur Basin, before the basin was deformed. Structure Deformation of the basin has been mainly in response t o complex block faulting, particularly along the Batten and t o lesser extent the Urapunga Fault Zones (Fig. 4.10); deformation of the adjoining shelves is mild (Fig. 4.5) (Plumb et al., 1980). Fault movements in the northerly trending Batten Fault Zone are apparently dominated by large vertical block displacements of u p to 7.5 km, but right-lateral strike-slip displacements are also apparent. Deformation of the Urapunga Fault Zone was in response to westii-ending reverse faults. Major northwest-trending faults of the Arnhem and Wearyan Shelves (Bulman and Calvert Faults) locally show small left-lateral strike-slip displacements. The well-known joint pattern in the Katherine River Group, on the Amhem Shelf, has originated by the classic ordered wrench fault tectonics of Moody and Hill (1956), controlled by major shears parallel to the Bulman Fault and Batten Fault Zone. Considerable preMcArthur Basin strike-slip displacements may be demonstrated on all the major faults which extend into the basement.
249 The structural development of the McArthur Basin was controlled by strike-slip movements on major pre-existing basement faults of continental proportions. Large secondary vertical displacements are common features of such faults. The Batten Fault Zone overlies a major zone of weakness and has dominated most of the structure of the basin. The Urapunga Fault Zone is situated over a basement ridge which resisted the n o r t h s o u t h horizontal movements of the surrounding blocks, and so produced the reverse faulting on westerly trending faults. Mineral deposits Mineralization in the McArthur Basin is dominated by the huge shalehosted stratiform McArthur River (H.Y .C.) Pb-Zn-Ag deposit, and similar smaller deposits nearby. Reserves of the single deposit are around 190 million tonnes, assaying 9.5% Zn, 4.1% Pb, and 44 g/tonne Ag (Murray, 1975; Walker et al., 1978), but metallurgical problems have prevented its commercial exploitation. Finely banded galena and sphalerite is hosted by laminated, turbiditic, pyritic-carbonaceous-dolomitic shales of the Barney Creek Formation, which were deposited in a locally subsiding basin (Bulburra Depression), adjacent t o the syndepositional Emu Fault and within the overall carbonatesabkha terrain of the Umbolooga Subgroup. Because of the lack of metamorphism and recrystallization of the ore, and the well-preserved microfossils associated with the sulphides, the deposit is the subject of considerable research into syngenetic and possibly biogenic ore formation (e.g., Croxford, 1968; Lambert and Scott, 1973; Smith and Croxford, 1973,1975; Croxford et al., 1975; Lambert, 1976; J.H. Oehler, 1977; Oehler and Logan, 1977; Williams, 197813,1979). Small carbonate-hosted stratabound (Mississippi Valley-type) Pb-Zn desposits are also found in the McArthur Group, such as Coxco, Cooley, and Ridge (Walker et al., 1978; Williams, 1978a), and at Bulman (Patterson, 1965) in the equivalent Dook Creek Formation on the Arnhem Shelf. Several occurrences of stratiform Cu deposits occur in dolomites and red siltstones of the McArthur and Tawallah Groups, while small Cu-breccia deposits are associated with recently identified unconformities and karstic surfaces in the McArthur Group (Muir, 197913). A distinctive type of Cu mineralization occurs in a number of volcanic breccia pipes, associated with upper Tawallah Group volcanism at Redbank (Orridge and Mason, 1975; Knutson et al., 1979). Minor U mineralization is associated with altered basic dykes intruding basal Tawallah Group arenites at Westmoreland (Hills and Thakur, 1975). The subeconomic Roper River Fe deposits consist of interbanded pisolitic hematitesiderite--chamosite deposited in a marginal-marine environment (Cochrane and Edwards, 1960; Canavan, 1965; Peat et al., 1978).
250 Lawn Hill Platform The Lawn Hill Platform contains a very similar succession t o the McArthur Basin (Fig. 4.11). It is separated from the McArthur Basin by the Murphy Inlier, and it grades eastwards into the Mount Isa Orogen (Fig. 4.10). Its evolution will be discussed with the Mount Isa Orogen. Birrindudu Basin (Principal references: Dow and Gemuts, 1969; Sweet e t al., 1974a; Blake e t al., 1975, 1979; Sweet, 1977).
The Birrindudu Basin (Fig. 4.1,13)is the mildly deformed basin, about 120,000 km2 in area and overlying the Sturt Block, within which the stratigraphically equivalent Birrindudu Group, Limbunya Group, and the Mount Parker Sandstone/Bungle Bungle Dolomite were deposited (this definition (Plumb and Derrick, 1975) differs from that of Blake et al., 1975, 1979). Small discontinuous inliers of sandstone suggest that the basin may extend
TABLE 4.V Summary of Carpentarian (Mid-Proterozoic) stratigraphy, Birrindudu Basin Unit and distribution
Rock types and thickness
Remarks
BIRRINDUDU GROUP (Sturt Block-The GranitesTanami region)
Argillaceous sublithic and quartz sandstones, conglomerate, dolomite, chert, siltstone. Stromatolites; some glauconite. Up to 6000 m
Unconformable between basement and Redcliff Pound Gp. Abundant crossbeds; ripple marks, mud cracks. Marked thickness variations
LIMBUNYA GROUP (Sturt Block-Victoria River region)
Dolomite, siltstone, shale, sandstone, chert, dolarenite, conglomerate. Abundant varied stromatolites. 1500 m+
Correlated with Birrindudu Gp. Unconformable between Bunda Grit and Victoria R. Basin succession. Marked thickness variations in basal sandstone
Mount Parker Sandstone/ Bungle Bungle Dolomite (Sturt Block -adjacent to Halls Cr. Mobile Zone)
Dolomite, dolomitic siltstone, quartz sandstone, black shale, micaceous shale. Abundant and varied stromatolites. 1750 m+
Precise correlations with Limbunya Gp. Unconformable between Red Rock Beds and Victoria R. Basin succession
Colombo Sandstone (Intersection Halls CreekKing Leopold Mobile Zones)
Quartz sandstone, chertpebble breccia. 1 0 0 m+
Unconformable on Kimberley Basin succession. Possibly correlates with Mt. Parker Sst.
251 eastwards to include the Tomkinson Creek Beds (Randal and Brown, 1969), to the north of the Tennant Creek Inlier. The stratigraphy of the basin is summarized in Table 4.V. Age The Birrindudu Basin succession, of arenites passing up into carbonates, broadly resembles the Tawallah and McArthur Groups of the McArthur Basin. The succession unconformably overlies the 1765-1685 Ma-old transitional granites of The Granites-Tanami Inlier, and was moderately deformed before the unconformably overlying Victoria River Basin succession, with minimum ages of around 1200 Ma (Webb and Page, 1977), was deposited. Glauconite from the Birrindudu Group gives minimum ages (c. 1560 Ma; Page et al., 1976) similar to those obtained from the upper Tawallah Group (McDougall et al., 1965). Further support t o the correlations between the Birrindudu and McArthur Basins is provided by microfossil assemblages in the Bungle Bungle Dolomite, which closely resemble those in the McArthur Group (Diver, 1974). Palaeogeograp hy The Birrindudu Basin succession shows lateral uniformity over wide areas, except the lower arenite unit which reflects irregularities in the basement surface, particularly in the Birrindudu area and adjacent t o the Halls Creek Mobile Zone. The Tomkinson Creek Beds may be up to 15 km thick, but the form of the structure in which they were deposited is unknown. Analogies with the McArthur Basin indicate that the Birrindudu Basin succession was deposited as a basal, largely fluviatile sand, passing up into a wide marginalmarine shelf, where carbonates were deposited in subtidal, intertidal, and probably supratidal environments (Plumb and Gemuts, 1976; Sweet, 1977). Structure Above the central Sturt Block, of the Victoria River region (Fig. 4.13), the Limbunya Group is only gently t o moderately folded and faulted (Sweet, 1977). Adjacent t o the Halls Creek Mobile Zone, the Bungle Bungle Dolomite is broadly folded and intensely faulted by thrust faults developed obliquely to the strike-slip Halls Creek Fault (Fig. 4.5) (Plumb and Gemuts, 1976). Southwards, in the The Granites-Tanami region, the Birrindudu Group is folded into moderately tight folds between basement domes, in response to superposed movements on trend with the nearby Halls Creek and King Leopold Mobile Zones (Blake et al., 1979). Victoria River Basin and equivalents (Principal references: Dow et al., 1964; Sweet et al., 1974a,b,c; Plumb and Gemuts, 1976; Sweet, 1977).
252
Fig. 4 . 1 3 . Diagrammatic stratigraphy and tectonic setting, Victoria River Basin and equivalents.
253 Up to 3.5 km of virtually undisturbed stable-shelf carbonate and terrigenous sequences accumulated in the Victoria River Basin, which covers an area of about 160,000 km2 on the Sturt Block. At the same time, very much thicker (up to 9 km) unstable-shelf terrigenous sequences accumulated in the adjacent Mobile Zones: the Carr Boyd Group in the Halls Creek Mobile Zone and the Fitzmaurice Group in the Fitzmaurice Mobile Zone. These groups are up to six times thicker than their immediate stratigraphic equivalents on the Sturt Block (Fig. 4.13), and have been moderately to intensely deformed by major faulting in the Mobile Zones. Small outliers of Glidden Group (Roberts et al., 1972), to the west of the Halls Creek Mobile Zone, are correlated with part of the Carr Boyd Group. The stratigraphy is summarized in Table 4.VI. The tectonic setting and correlations are summarized in Fig. 4.13.
S tra tigrap h ic re la tionsh ips The Victoria River Basin succession and equivalents unconformably overlie the Birrindudu Basin, the Kimberley Basin and equivalents, and basement rocks of the Halls Creek and Pine Creek Inliers. The succession is unconformably overlain by late Adelaidean glacial successions, and by Palaeozoic rocks of the Ord, Bonaparte Gulf, Wiso, and Daly River Basins. A feature of the successions in the Victoria River Basin and equivalents is the presence of several unconformities. The key to the correlations (Fig. 4.13) is direct field correlation between the Golden Gate Siltstone, Goobaieri Formation, “Undivided Auvergne Group”, and Helicopter Siltstone, and between the Pincombe and Legune Formations (Sweet, 1977), and the recognition of an unconformity above the Mount John Shale Member (Plumb and Gemuts, 1976). A major part of the Victoria River Basin succession was deposited before sedimentation commenced in the Mobile Zones and an apparent interval of non-deposition in the Victoria River Basin is reflected in the derivation of sediment from the immediately adjacent Sturt Block, in the Lissadell and Glenhill Formations of the Carr Boyd Group. Age The preferred general age range of the successions is provided by stratigraphically consistent shale whole-rock Rb-Sr isochrons from the Carr Boyd Group (Bofinger, 1967); the Golden Gate Siltstone 1 1 5 8 k 1 2 3 Ma; the Glenhill Formation 1057 k 80 Ma; and the Pincombe Formation 891 k 149 Ma. The Mount John Shale Member at 1104 k 110 Ma, and the Glidden Group at 1019 k 51 Ma are consistent with the correlations. Shale data from the Victoria River area show a wide scatter. K-Ar and Rb-Sr ages of glauconite from the Wondoan Hill Formation (1125 k 30 Ma; Webb and Page, 1977) provide the best agreement. These data therefore put the successions somewhere in the interval 1200-900 Ma, with emphasis towards the older rather than the younger limits.
254 TABLE 4.VI Summary of Mid-Proterozoic stratigraphy, Victoria River Basin and equivalents Unit
Rock types and thickness
remarks
Ferruginous and feldspathic sandstone; conglomerate. 320 m
Unconformable o n Auvergne Gp. Ubiquitous cross-beds
AUVERGNE GROUP
Quartz sandstone, siltstone, shale, dolomite, dolomitic siltstone and sandstone. Glauconite. 500-1000 m
Unconformable o n S t u b b and Wondoan Hill Fms. Cross-beds, ripple marks, m u d cracks, m u d flakes, halite casts, etc.
Wondoan Hill-Stubb Formations
Quartz sandstone, grey shale and siltstone, glauco nite. U p t o 260 m
Unconformable o n Bullita GP.
BULLITA GROUP
Dolomite, dolomitic siltstone and sandstone, siltstone, sandstone. 500-800 m
Conformable o n Wattie Gp. Stromatolites, oolites, halite casts, mud cracks
WATTIE GROUP
Quartz sandstone, siltstone. 350-600 m
Unconformable o n Limbunya Gp. Ubiquitous cross-beds, ripple marks
Wade Creek Sandstone (upper) - Helicopter Siltstone
Quartz sandstone; laminated siltstone and shale. 250 m
Unconformable between Duerdin Gp. and Mt J o h n Sh. Mbr. Equivalent to Auvergne Gp.
Wade Creek Sandstone (lower) - Mount J o h n Shale Member
Quartz sandstone; black and grey shale. Up to 300 m
Unconformable o n Bungle Bungle Dol. Equivalent t o Wondoan Hill-Stubb Fms.
TOLMER GROUP
Quartz sandstone, siltstone, dolomite. U p to 1000 m
Unconformable between Pine Cr. Inlier and Cambrian volcs. Equivalent t o Wattie-Bullita Gps.
Cyclic alternations of quartz sandstone and siltstone, shale, fine sandstone. Some black (pyritic) shale, ferruginous o r feldspathic sandstone, conglomerate. U p to 6000 m
Unconformable between Revolver Cr. F m . o r Lamboo Cmplx. and Cambrian volcs. Unconformities separate major cycles. Marked facies changes. Equivalent in part to Auvergne Gp.-Bullo R. Sst.
Victoria River Basin Succession Bullo River Sandstone
Iialls Creek-Fitzmaurice Mobile Zone CARR BOYD GROUP
255 TABLE 4.VI (continued) Unit
Rock types and thickness
Remarks
Halls Creek-Fitzmaurice Mobile Zone (continued) FITZMAURICE GROUP
K i m b e r l e y Block GLIDDEN GROUP
Alternating quartz sandstone and siltstone, shale, fine sandstone. Some feldspathic sandstone, conglomerate. Up t o 10,000 m
Unconformable on Lamboo Cplx. and Pine Cr. Inlier. Unconformities; facies changes. Equivalent to parts Carr Boyd Gp. and t o Auvergne Gp.-Bullo R. Sst.
Quartz sandstone, siltstone, black shale. 550 m
Unconformable between Kimberley and Kuniandi Gps. Equivalent t o part of Carr Boyd Gp.
Palaeogeograp hy Victoria River Basin. Lateral continuity of thin lithostratigraphic units over wide areas, evidence of subaerial exposure, evaporite minerals, and abundant columnar stromatolites are features of much of the succession. Several cycles of transgression and regression have resulted in paralic sedimentation on a broad stable shelf, interspersed with periods of subaerial erosion. Major cycles comprise basal arenites grading up into lutites and carbonates (Fig. 4.13). Depositional environments include marginal-marine, lagoonal, intertidal, shallow sub-tidal, evaporitic, and fluvial (Sweet, 1977). Halls Creek/Fitzmaurice Mobile Zones. Distinctive features include marked changes in thickness and facies changes across the mobile zones, syndepositional fault movements, truncation of syndepositional faults by unconformities, and source areas in both the Kimberley and Victoria River Basin areas. Palaeocurrent directions vary throughout the sequence. A particular example is the Golden Gate Siltstone, which changes, in only 10 km, from 270 m of near-shore sands, silts, and ironstone in the west, t o over 2000 m of pyritic black shale in the east, adjacent t o the Halls Creek Fault. In comparison, the Glenhill Formation shows an easterly derived, alluvial-fan deposit in the east (adjacent to a major fault zone), passing into a glauconitic marine succession farther west. The Carr Boyd Group comprises four 1500 m-thick unconformable cycles of basal sandstone grading up into lutite-rich sequences: the product of repeated cycles of marine transgression and regression, produced by periodic subsidence and uplift of the narrow fault-bounded Halls Creek Mobile Zone
256 and its margins; only two of the cycles are preserved in the Fitzmaurice Group. Ideally, basal fluvial sands pass up into marine (tidal-bay, tidalchannel, and subtidal complexes), thence back t o marginal-marine (tidal-flat) deposits, followed by erosion and the next cycle. The multi-cycle source has produced mature sediments in a tectonically active area (Dow et al., 1964; Plumb and Gemuts, 1976).
Structure The strata of the Victoria River Basin are flat-lying and virtually undisturbed. Deformation increases suddenly adjacent to the mobile zones (Fig. 4.5; Sweet, 1977). The Carr Boyd Group has been deformed by the major north-northeasterly trending strike-slip faults of the Halls Creek Mobile Zone, and associated broad folding (Fig. 4.5). Many of the displacements were syndepositional, but post-Carr Boyd Group left-lateral displacements of at least 30 km have been identified. Complex vertical displacements, commonly reversed, are up to 2-3 km. Subsidiary east-northeasterly trending thrust fults diverge from the strike-slip faults. The Carr Boyd Group is unmetamorphosed, except for local development of slaty cleavage within or between the larger shear zones (Plumb, 1968; Plumb and Gemuts, 1976). The Fitzmaurice Mobile Zone has been considered to be an extension of the Halls Creek Mobile Zone (Traves, 1955; Rod, 1966), but it is considered by Plumb (197913) to be a subsidiary feature, diverging to the east from the main Halls Creek Mobile Zone, which swings northwards beneath the sea (Figs. 4.1 and 4). Left-lateral strike-slip displacements in the Fitzmaurice Fault Zone are not as large as in the Halls Creek Mobile Zone, and the eastern boundary fault (Victoria River Fault) has a major thrust component in the south, where it strikes northeast t o east-northeast (Fig. 4.5; Sweet, 1977). EARLY TO MID-PROTEROZOIC OROGENIC DOMAINS
Mount Isa Orogen -Lawn Hill Platform (Principal references: Carter e t al., 1961; Plumb and Derrick, 1975; Glikson et al., 1976; Wilsonand Derrick, 1976; Derrick et al., 1977; Wilson, 1978; Plumb et al., 1980; Williams, 1980)
The Mount Isa Orogen and Lawn Hill Platform are exposed over about 80,000 km2 (Fig. 4.1). Geophysical evidence shows that the Orogen extends in the subsurface, without structural break, for at least another 70,000 km2 to the south. To the north it may extend as far as the western edge of the Cape York Peninsula (Fig. 4.4; Plumb, 1979b). The region contains the massive shale-hosted stratiform Pb-Zn-Ag deposits of Mount Isa, Hilton, and Lady Loretta, the huge copper deposit at Mount Isa, and Australia’s only
257 TABLE 4.VII Summary of Early Proterozoic and Carpentarian (Mid-Proterozoic) stratigraphy, M o u n t Isa Orogen, Lawn Hill Platform a n d S o u t h Nicholson Basin Unit and distribution
Rock types and thickness
Remarks
SOUTH NICHOLSON G R O U P Blanket quartz arenites and equivalents alternate with lutites
Mildly deformed Platform Cover, unconformable on Mt. Isa G p . equivs.
SOUTH NICHOLSON GROUP (South Nicholson Basin)
Quartz sandstone, micaceous siltstone and (glauconitic) sandstone, oolitic ironstone, conglomerate, calcareous greywacke. 1600-9000 m
Cross beds, ripple marks, scour-and-fill, mud cracks, mud pellets. Constance Ra. F e
Pilpah Sandstone (West of Sybella Granite)
Ferruginous and feldspathic sandstone, siltstone, shalepellet and chert-fragment conglomerate. 300 m Syntectonic, meridional d y k e swarms in Leichhardt R . Fault Zone; post-tectonic dykes in Mary Kathleen Fold Belt
Dolerite dykes
Regional metamorphism
Greenschist grade in Mt. Isa area
Accompanies regional folding and faulting
M O U N T 1.5.4 GROUP and equivalents
Dominantly carbonates and lutites
Unconformable on Surprise Ck. Fm. Quilalar Fm., Haslingden G p . and Sybella Granite
MOUNT ISA GROUP (Leichhardt River Fault Trough)
Dolomitic and pyritic siltstone, black shale, dolomite, K- and Na-rich tuffites, chert, quartzite; siliceous and dolomitic breccia. Up to 5000 m
Fine clastics mainly flatlaminated; minor algal laminae; some slumping, grading, and Bouma cycles; evaporite pseudomorphs near t o p of Gp. and in orebody. Mt. Isa Ag-Pb-Zn; Cu in silica-dolomite breccia; Hilton Ag-Pb-Zn
McNAMARA GROUP (Lawn Hill Platform)
Dolomite, dolomitic and pyritic siltstone, black shale, tuffite, quartzite, lithic sandstone, quartz wacke, siltstone, shale. Up t o 7500 m
Stromatolites abundant; gypsum and anhydrite pseudomorphs, teepee structures, m u d cracks; Lady Loretta Ag-Pb-Zn with baryte; Kamarga Z n prospect; Lawn Hill P b in veins
258 TABLE 4.VII (continued) Unit and distribution
Rock types and thickness
Remarks
MOUNT ISA GROUP and equivalents (continued) FICKLING GROUP (Lawn Hill Platform)
Dolomite, dolomitic siltstone, chert, pyritic black shale, sandstone, conglomerate. 950 m+
Oolites, stromatolites; cross-beds, ripples, mudflakes, halite casts; glauconite in sandstone; minor Pb Mount Albert Group contains amphibolite facies rocks
Regional metamorphism in east
Arenites grading upwards to lutites
Regionally unconformable on older units
Surprise Creek Formation (West of Kalkadoon-Leichhardt Block)
Conglomerate, sandstone, micaceous and carbonaceous siltstone, shale, minor dolomite. 400-2500 m
Cross-beds, ripple marks; flaser bedding, convolutions, Bouma cycles in fine clastics
MOUNT ALBERT GROUP (Mary Kathleen Shelf)
Conglomerate, quartzite, siltstone, phyllite, slate, limestone, garnetstaurolite schist. 1000-3700 m
S URPRISE CREEK FORMA.
TION and equivalents
Dominantly felsic and minor basic igneous actiuity
Vulcanicity, granite intrusion and regional metamorphism Fiery Creek Volcanics (Lawn Hill Platform)
Rhyolite, trachyte, trachybasalt, basalt, agglomerate tuff, conglomerate, sandstone, dolomite, siltstone. Up t o 700 m
Regionally unconformable on Quilalar Fm. and Haslingden Gp. May postdate Weberra Gr.
Peters Creek Volcanics (Lawn Hill Platform)
Basalt, rhyolite, rhyodacite, trachyte, andesite; tuff, shale, siltstone, sandstone, conglomerate, dolomite. 1500-2000 m
Equivalent t o Tawallah Gp. volcanics in McArthur Basin
Carters Bore Rhyolite (Leichhardt River Fault Trough)
Rhyolite, quartz-feldspar porphyry. 1 5 0 m
Weberra Granite (Lawn Hill Platform)
Leucogranite, microgranite
Calc-silicate hornfels in Quilalar Fm.
Sybella Granite (Leichhardt River Fault Trough)
Biotite granite, gneissose granite, microgranite, pegmatite, diorite
Intrudes Haslingden Gp; overlain by Mt. Isa Gp. and equivs. Rare-earths and mica in pegmatites
259 TABLE 4.VII (continued) Unit and distribution
Rock types and thickness
Remarks
Burstall, Wonga, Williams, Wimberu, Tommy Creek, Naraku Granites (Mary Kathleen Fold Belt)
Hornblende-biotite granite, alkali granite, rhyolite, aplite, minor gneissose granite
Intrude Tewinga, Malbon, Mary Kathleen and Soldiers Cap Gps.; syn- to posttectonic; U-F at Mary Kathleen (Burstall Gr.), where skarn replaces regional metamorphics
Lunch Creek Gabbro and related basic intrusives (Mary Kathleen Fold Belt)
Olivine gabbro, two-pyroxene-biotite gabbro, hornblende gabbro, diorite, dolerite. 500 m+
Syn- to post-tectonic sills and dykes, K-rich tholeiite t o calc-alkaline geochemistry
M A RY KATHLEEN GROUP and equivalents
Transgressive, shallo w-shelf arenite-carbonate association
Conformable or disconformable on Malbon, Sol diers Cap and Haslingden Gps.; unconformable on basement Tewinga Gp.
MARY KATHLEEN GROUP (Mary Kathleen Shelf)
Metamorphosed conglomerate, sandstone, siltstone, shale, limestone; calc-silicate granofels, marble, schist, basalt, amphibolite, tuff, porphyry 3000 m
Dugald River Zn-Pb in black shale basin in carbonate platform; some associated anhydrite pseudomorphs
Quilalar Formation (West of Kal kadoon-Leichhardt Block)
Conglomerate, sandstone, dolomite, dolomitic sandstone, siltstone and mudstone (red beds), minor tuff and trachybasalt. 1800 m
Bevelled ripple marks, mud cracks, cross-beds, oolites, halite casts, possible anhydrite pseudomorphs
HASLINGDEN GROUP and equivalents
Dominantly arenite-basalt association in rifts and marginal basins
Unconformable and conformable on Tewinga G p .
HASLINGDEN GROUP (West of KalkadoonLeichhardt Block)
Most basalt subaerial Arkose, tuff, greywacke ; quartzite, basalt; sandstone, (6000 m); ripples, crossbeds, mud-cracks, halite dolomitic red-bed siltstone casts in west. Mammoth and dolomite. 6000 m Cu in Myally Subgp. Cu in basalt flow-tops
MALBON GROUP (Mary Kathleen Shelf)
Basalt, quartzite, siltstone, limestone, stromatolitic jaspilite. 4000 m
Basalt subaerial t o shallow submarine; jaspilite, with anhydrite pseudomorphs, transitional to Mary Kathleen Gp.
260 TABLE 4.VII Unit
Rock types and thickness
Remarks
SOLDIERS CAP GROUP (Mary Kathleen Shelf)
Schist, greywacke, quartzite, basalt, amphibolite, siltstone, BIF, limestone, chert. 6000 m+
Base n o t exposed; basalt chemically less continental than Haslingden Gp. basalts. Pegmont Pb-Zn, Monakoff Cu-Zn-Ba-F, Squirrel Hills Cu-Zn
Granite intrusion and regional m P l a ni orph ism
Western edge of Kalkadoon-Leich. B l o c k ; predates basalt in Haslingden GP.
Ewen Granite and equivalents
Coarse to medium granite and aplite
T B W I N G A G R O U P and cyuivalents
Dominantly subaerial acid volcanics, m i n o r basalt and sediments
Basement t o overlying sequences
Argylla Formation, Magna Lynn Metabasalt and equivalents (KalkadoonLeichhardt Block)
Rhyolite, dacite, ash-flow tuff, basalt, quartzite, siltstone; gneiss, schist. 1700 m+
Metamorphosed by Wonga and Ewen Gr. May overlie Leichhardt Mmics. unconformably
Leichhardt Metamorphics (Kalkadoon--Leichhardt Block)
Dacite, rhyodacite, ashflow tuff, rhyolite, schist, gneiss, migmatite. 4 0 0 0 m +
Base n o t exposed; unconformity with upper Tewinga Gp. possibly related t o intrusion of Kalkadoon Gr.
Kalkadoon-Leich.
Granite intrusion and regional i n e tam o rp h is m Kalkadoon Granite
Block
Granite, granodiorite, aplite, pegmatite
Batholithic pluton with metamorphosed roof pendants of ?Leichhardt Mmics.
Phyllite, schist, gneiss, migmatite, rhyolite. 1500 m
Intruded by Big Toby Gr. (Rb-Sr 1 7 4 0 Ma); overlain by Haslingden Gp.
Basement rocks of u n k n o w n affiliation Yaringa Metamorphics (West of Leichhardt River Fault Trough)
currently (1978) producing uranium mine at Mary Kathleen. The tectonic setting and principal tectonic elements are shown in Figs. 4.1, 4.4 and 4.10. Stratigraphy is summarized in Table 4.VII and Fig. 4.11.
Tectonic setting The Mount Isa Orogen is an Early to Mid-Proterozoic mobile belt, situated
261 along the eastern edge of the North Australian Craton. Early Proterozoic rocks, continuous with those of the North Australian Orogenic Province to the west, were deformed, metamorphosed, and intruded by granite, to form basement to a succession of late Early Proterozoic and Carpentarian rocks which were deposited, intensely deformed, metamorphosed, and intruded by granite, at the same time that the North Australian Platform Cover was being deposited on the craton t o the west. The Lawn Hill Platform is part of the North Australian Platform Cover, and is transitional between the Orogen to the east and the mildly deformed McArthur Basin t o the northwest. The boundary between the Lawn Hill Platform and the Mount Isa Orogen is placed at the Mount Gordon Fault Zone (Fig. 4.10). The Murphy Inlier lies between the Lawn Hill Platform and the McArthur Basin. The South Nicholson Basin unconformably overlies the rocks of the Lawn Hill Platform, and contains a sequence of similar age and type t o the Roper Group of the McArthur Basin. The sequence was deposited either after or during the final deformation of the Mount Isa Orogen. Palaeozoic cover of the Georgina Basin overlies the Mount Isa Orogen, Lawn Hill Platform, and South Nicholson Basin in the west; to the east they are overlain by Mesozoic cover of the Carpentaria and Eromanga Basins, and these covers obscure the relationship of the Orogen t o the Northeast Queensland Precambrian Province and to the Amnta Inlier. Palaeogeographic and stratigraphic studies suggest that the Mount Isa Orogen was formerly a rifted continental margin (Glikson et al., 1976), similar to parts of the Red Sea intracontinental marine embayment, the Appalachianrtluachita structural system of eastern North America (Thomas, 1977), and the Coronation Geosyncline (Hoffman, 1973) along the western margin of the Archaean Slave Province in North America. Stratigraphy Seven major stratigraphic subdivisions are recognized in the Mount Isa Orogen and Lawn Hill Platform (Fig. 4.11; Table 4.VII). The Tewinga Group, at least 5.7 km thick, is a dominantly acid and basic volcanic suite which contains a major regional unconformity within it. It forms basement to overlying sequences. Arenite and basalt of the Haslingden Group and equivalents overlie the Tewinga Group, both conformably and unconformably. The Haslingden Group is up t o 16 km thick in the west, and includes 6 k m of mainly continental tholeiitic basalt; equivalent sequences in the east are from 4 to 5 km thick, and contain lesser volumes of basalt. The Mary Kathleen Group and equivalents were deposited as a transgressive clastic-carbonate sheet, up to 3 km thick, across older groups; contacts with Haslingden Group and equivalents are mainly conformable, and contacts with Tewinga Group rocks are unconformable. This transgressive phase was terminated by folding, metamorphism, granite intrusion, uplift, and deposition of up t o 700 m of
262 subaerial Fiery Creek Volcanics and equivalents, mainly in the west of the Orogen and on the Lawn Hill Platform. The volcanics rest with marked unconformity on older rocks. A coarse fining-upwards clastic-siltstone sequence up t o 3.7 km thick, the Mount Albert Group and equivalents, rests unconformably on the Fiery Creek Volcanics and equivalents, and passes conformably or unconformably upwards t o the Mount Isa Group and equivalents, up t o 7.5 km thick, which are present only in the west of the Mount Isa Orogen and on the Lawn Hill Platform. The latter contain mainly fine clastics and carbonate rocks, and are host to the major stratiform Pb-Zn-Ag deposits at Mount Isa, Hilton, and Lady Loretta. The South Nicholson Group, a mainly shallow-water, coarse- t o fine-grained arenaceous and lutitic sequence 1.6 t o 9 km thick, overlies the Mount Isa Group and equivalents unconformably, and occurs only above the Lawn Hill Platform. It is host to large but subeconomic deposits of ironstone. Age and correlations Earlier Rb-Sr and K-Ar geochronology in the Mount Isa Orogen (e.g., Richards et al., 1963; Richards, 1966) has been largely superseded by recent U-Pb zircon geochronology (Page, 1978). Rhyolite and rhyodacite from the Tewinga Group give U-Pb zircon ages (Page, 1978) of 1865 3 Ma for the basal formation (Leichhardt Metamorphics) and 1777 2 7 Ma for the topmost formation of the Group (Argylla Formation). A regional unconformity is thought to separate these two formations. Kalkadoon Granite intruding Leichhardt Metamorphics gives a U-Pb zircon age of 1862 ? 24 Ma (Page, 1978), which suggests that these two units are comagmatic. The Ewen Granite, with a minimum age of about 1775 Ma (K-Ar biotite; Richards et al., 1963) or 1742 ? 20 Ma (whole-rock Rb-Sr; McDougall et al., 1965) intrudes the upper Tewinga Group and is overlain unconformably by basalt within the Haslingden Group. This age for the base of the Haslingden Group is consistent with the 1777 Ma-old conformably underlying Argylla Formation. Burstall Granite, intruding the Mary Kathleen Group in the east of the Mount Isa Orogen, gives a U-Pb zircon age of about 1720 Ma (Page, 1979). Correlations (Fig. 4.11) indicate that this age is also a minimum age for the top of the Haslingden Group. A major unconformity separates the Carters Bore Rhyolite (and the equivalent Fiery Creek Volcanics) from the older Rocks. The Carters Bore Rhyolite yields a U-Pb zircon age of 1678 f 1Ma (Page, 1978). This age supersedes previously published Rb-Sr ages (1611--1504 Ma; Wilson and Derrick, 1976) for the Sybella Granite, which intrudes Haslingden Group and underlies, or is partly comagmatic with, the Carters Bore Rhyolite. In the east of the Mount Isa Orogen, the syntectonic Wonga Granite yields a U-Pb zircon age of 1671 k 8 Ma, and widespread metamorphic overprints throughout the Mount Isa Orogen produce disturbed Rb-Sr isochrons
*
263 between 1670 and 1620 Ma (Page, 1978). The Mount Isa Group (partly equivalent t o the McArthur Group) contains tuff which has a U-Pb zircon age of between 1650 and 1670 Ma (Page, 1979). This depositional age compares with model lead ages of between 1600 and 1500 Ma for the Mount Isa orebodies (Richards, 1963; Ostic et al., 1967; Cooper et al., 1969). Post-Mount Isa Group metamorphism has a minimum age of 1490--1470 Ma (Rb-Sr biotite; Page, 1978). The South Nicholson Group overlies Mount Isa Group equivalents unconformably. The Mullera Formation gave a shale total-rock isochron of 1480 120 Ma (R.R. Harding, BMR, unpublished), consistent with the possibly greater than 1400 Ma age for the correlative Roper Group in the McArthur Basin. New data since Plumb et al. (1980) suggest some modification t o the correlations shown in Fig. 4.11, but further confirmation is desirable. Field studies now suggest that the Peters Creek Volcanics may be equivalent only to the Fiery Creek Volcanics (Cavaney, 1975; I.P. Sweet, pers. commun., 1979), and not to the Haslingden Group as well (cf. Fig. 4.11). Since the Peters Creek Volcanics probably correlate with almost all of the Tawallah Group of the McArthur Basin (Plumb and Sweet, 1974), it follows that the Haslingden Group is probably older than the Tawallah Group (cf. Plumb and Derrick, 1975; Plumb et al., 1980). The latest isotopic data support these revisions: the maximum age for the base of the Tawallah Group appears t o be around 1730--1688 Ma, compared to an age range for the Haslingden Group of between 1777 and 1720 Ma. A consequence of all of this is that the Haslingden Group would no longer be part of the Carpentarian.
*
Tectonic elements Three major palaeogeographic elements are recognized within the Mount Isa Orogen: a core of basement rocks, flanked to the west and east by belts of younger sediments and volcanics. Early Proterozoic basement volcanic and granitic rocks are exposed in a north-trending median ridge, 300 km long and about 30 km wide - the Kalkadoon--Leichhardt and Ewen Blocks. The Malbon Block is an inlier of equivalent rocks to the east. To the west of the median ridge over 6 km of basalt and 20 km of sedimentary rocks accumulated during the late Early Proterozoic and Carpentarian in the Leichhardt River Fault Trough, a north-trending rift or aulacogen (Glikson et al., 1976; Dunnet, 1976a), formed between the Quilalar Arch (in the east) and Mount Gordon Arch (in the west). Field data and airborne magnetic patterns indicate that basalts are preserved in a trough 65 km wide and at least 350 km, and possibly 550 km, long. The trough has subsequently been deformed into the Leichhardt River Fault Zone. To the west of the Mount Gordon Fault Zone (Arch) equivalent sequences on the Lawn Hill Platform are less intensely deformed and metamorphosed than those of
FAULT TROUGH
MARY KATHLEEN SHELF
BLOCK
EAST c1oncurry
____~___ CARPENTARIAN SEDIMENTARY COVERS croion
0-
(SUBJECT TO LATER OROGENY)
I”
2 M o r g no1 s u b s t d e n c e
u n d l.ansgresslon, Shallow s h e l f - c o o s t a l - OIIUYIOIp l a l n d e p o s i l i o n , l e r m i n o t e d by u p l l f t a n d d c o n sm , o n s e t of r e g t o n o l m e t a m o r p h i s m 7 1750 - 1680 m y MI
-
3
C
MK
P o r o o e o s y n c l i n a l stage regresslve sandstone t r o n s i t ~ o n a l lo s ~ l l s t o n e a n d s h a l e , f u r t h e r rnefomorphlsm ~n e a s t
1680- 71600 MI
MX
-~ 4 Parageosyncl no1 stage ( c o n t ~ n u e Id , c r a l o g e n c o r l h o q u o r t z ~ l etrons I on01 to f n e - gro n e d f l y s c h o n d p l o l f o r m Corbonotes ~
1680-PI600 m y
. __ _ _
Fig. 4 . 1 4 . Volcanic, sedimentary, and plutonic evolution of the Mount Isa Orogen-Lawn Hill Platform. (Adapted from Plumb et al., 1980). LOWER PROTEROZOIC BASEMENT Kalkadoon Gr. ( 2 ) with remains of high-grade ( l a ) and lower grade ( I b ) acid volcanics. = Yaringa Mets and Big Toby Gr.; 3, 4, 5 = Tewinga Group (Leichhardt Mets = 3 ; Magna Lynn Metabasalt = 4; Argylla Fm. = 5 ) ; 4a = dolerite dykes. lc
265 the Trough, and they thin t o the northwest, towards the Murphy Inlier. To the east of the basement blocks, thinner sequences of mainly shallowwater sediments ac imulated on the Mary Kathleen Shelf, but the shelf has undergone more complex folding, regional metamorphism, and granite intrusion than other areas, t o form the Mary Kathleen Fold Belt. Sedimentation, vulcanism and basin development in the region have been largply controlled by major long-acting north-trending lineaments or zones, some of which have then been a focus for major post-depositional faulting, metamorphism, and granite intrusion.
Pulaeogeogruphic evolution The Mount Isa Orogen evolved from a basement of dontinental crust, through a rift basin and shallow-shelf continental-margin stage with craton interior t o the west, t o parageosynclines whose depocentres migrated cratonwards as the continental margin foreland underwent uplift and tectonism. The palaeogeographic and tectonic evolution of the Orogen (Plumb et al., 1980) is summarized in Fig. 4.14. Early Proterozoic basement The earliest rocks in the Orogen appear t o be metadacite and gneiss, now preserved as roof pendants in the Kalkadoon Granite. They may be part of an older continental crust, or the basal part of the Tewinga Group. The 5.7 km-thick Tewinga Group is a rhyodacite-basalt-rhyolite suite, whose geochemistry is typical of continental-margin Andean-type calc-alkaline volcanism (Wilson, 1978), except that andesite is lacking. In the southwest, a major unconformity separates ?Tewinga Group acid volcanics and metabasalt (Argylla Formation/Magna Lynn Metabasalt) from older gneiss and migmatite (possibly Leichhardt Metamorphics) (Bultitude et
CARPENTARIAN SEDIMENTARY COVERS
Stage 1: 6a-1Oa = Haslingden Gp.; 6b-lob = Malbon Gp.; 6c-9c = Soldiers Cap Gp. 6 = basal clastics; 8 = basalt suites; 9-10 = upper clastics to chemical sediments; 7a = Ewen Gr.; 7b = dolerite feeders t o 8. Stage 2: 1 1 = arbitrary boundary between Quilalar Fm. and Mary Kathleen Gp.; 12a, b = basal arkosic conglomerate, sandstone; 13a = carbonates and fine clastics; 13b = basalt and basic tuff; 14 = Zn-anomalous black shale basins; 15, 16 = basal sedimentary facies and upper volcanic facies of Fiery Creek Volc; 17 = rising geotherms and regional metamorphism; intrusion of basic t o acid sills and plutons ( 1 7a = Weberra Gr.; 17b = Sybella Gr.; 17c = Burstall, Wonga Gr.; 17d = Williams, Naraku Gr.). Stage 3: Surprise Creek Fm.; Mount Albert Gp.; 18-20 = fining-up sequences from conglomerate to siltstone. Stage 4 : Mount Isa, McNamara Gps. 21 = basal quartzite blanket; 22, 26 = shale and siltstone facies; 23a, 236, 23c = dolomite, dolomitic siltstone, pyritic zones with Pb-Zn-Ag, respectively; 24-25 = sandstone, quartz wacke.
266
al., 1977). In the east, an unconformity within the Tewinga Group is manifested by conjugate dolerite dyke swarms which intrude the Leichhardt Metamorphics, and terminate at the base of, and are possible feeders to, the Magna Lynn Metabasalt. Phyllite, schist, gneiss, and rhyolite of the Yaringa Metamorphics, in the far west of the Orogen, may be equivalent t o basal Tewinga Group, or they may be representative of older elements of the North Australian Orogenic Province. All these metasedimentary and volcanic-plutonic complexes formed a shelf of continental crust, on which the succeeding supracrustal sequences were deposited.
Late Early Proterozoic-Carpen tarian covers Stage 1 : R i f t s and marginal basins with basic volcanic suites. Large rifts and basins developed in the basement. Apart from the cessation of acid volcanicity, deposition appears t o have been locally continuous from the Argylla Formation into the overlying Haslingden and Malbon Groups; relations between the Soldiers Cap Group and Tewinga Group are not definitely known. The Haslingden Group was deposited in the Leichhardt River Fault Trough and on the adjoining Lawn Hill Platform. Basal fluviatile sheets and wedges of arkose and volcanoclastic conglomerate formed along and westwards of the Quilalar Arch, with some interbeds of basalt and tuff. These grade upwards into well-sorted shallow-marine (?) shelf arenites, which pass abruptly into a 6 km-thick sequence of mainly continental tholeiitic basalt. A lack of pillow lavas, an abundance of amygdaloidal basaltsandstone breccias and flow breccias, and intercalations of fluviatile arenite sheets and lenses are indicative of mainly subaerial basalt extrusion, in the steadily subsiding rift zone. The basalt sequence thins rapidly eastwards towards the Quilalar Arch, and contains abundant interbeds of conglomerate. Thinning also occurs westwards across the Mount Gordon Arch, but conglomerate interbeds are lacking. Following the volcanics, further shallow-marine feldspathic sandstone, derived from both easterly and westerly source areas, grade upwards to a regressive ferruginous dolomitic sandstone and siltstone facies (Myally Subgroup). The Myally Subgroup thins eastwards, where wedges of conglomeratic arkose lie along the Quilalar Arch, and also displays marked thinning and facies changes westwards across the Mount Gordon Arch. Up to 5 km of shallow-shelf arenites, deposited in the Leichhardt River Fault Trough, are transitional t o 1.5 km of a hematitic, dolomitic, and stromatolitic intertidal red-bed sequence on the Lawn Hill Platform. Thin regoliths are present at the base of the Subgroup on the Mount Gordon Arch. East of the Kalkadoon-Leichhardt Block, the Malbon Group thickens towards the east. Basal sandstone and mainly subaerial basalt, deposited in a
267
'broad basin or shelf, grade upwards to a shoreline sandstone-carbonateiron formation association, deposited in very shallow, probably evaporative, shelf and supratidal zones, adjacent t o a mature low-relief land mass to the west and southwest. Average thicknesses of about 15 km and 4 km in the Haslingden and Malbon Groups, respectively, reflect greater sedimentation and subsidence in the Leichhardt River Fault Trough. Basal graded pelitic sediments (?turbidites) and an abundance of pillow basalt suggest that the Soldiers Cap Group, to the east of the Malbon Group, was deposited on a continental shelf or slope, in deeper water than the probable time-equivalent sequences to the west. Chert and manganese-rich iron-formation intercalated with arenite and basalt are hosts t o stratiform base-metal accumulations. Stage 2: Marginal subsidence and transgression, terminated b y uplift and volcanism. The arrangement of large linear separate eastern and western basins was terminated by regional tilting followed by general subsidence, resulting in region-wide transgression and deposition of the Mary Kathleen Group and Quilalar Formation (Figs. 4.11 and 4.14). Although these units are now located east and west of the Kalkadoon-Leichhardt Block, respectively, they are believed t o have been continuous across the Block during sedimentation. Most contacts with the Malbon and Haslingden Groups are apparently conformable; contacts with the Soldiers Cap Group may be both conformable and unconformable. Both units are representative of a quartzite-carbonate association. Basal fluviatile, mainly arkosic, wedges and sheets pass upwards t o a more mature linear sandstone sheet deposited along a migratory shoreline; thickness variations are most evident across highs such as the Quilalar and Mount Gordon Arches. A thin-bedded carbonate-rich facies overlies or is partly time-equivalent to the linear sandstone sheet, and formed in shallow shelves and lagoons; a westwards increase in stromatolitic dolomite, shale-clast breccia, and fine red-bed clastics suggests that more continental conditions prevailed towards the Lawn Hill Platform. Halite in the west, and scapolite in the metamorphosed Corella Formation near Mary Kathleen t o the east, are indicative of widespread evaporitic conditions (Ramsay and Davidson, 1970). Offshore lenses and bars of orthoquartzite accompanied dolomite deposition throughout the shelf areas. The regional transgression was terminated by crustal uplift in the west, possibly accompanied by granite intrusion (Sybella, Weberra Granites). A regional unconformity was developed. Sheets of fluviatile pebbly sandstone and conglomerate (basal Fiery Creek Volcanics) prograded unconformably akross shallow-shelf sediments to the west. Between the' Mount Gordon and Quilalar Arches red-bed and dolomite deposition succeeded the pebbly sandstones. Extrusion of acid and intermediate lava, agglomerate, and basalt followed in the west and northwest (upper Fiery Creek Volcanics and equiva-
268
lents). There is no conclusive evidence of an equivalent angular unconformity to the east: the uplift and volcanism evident in the west may be manifested by tuff, agglomerate, basalt, and conglomerate lenses (middle and upper Corella Formation); small zinc-rich sub-basins of black euxinic shale and limestone also formed at this time, in the east. Following deposition of Fiery Creek Volcanics and Mary Kathleen Group, regional uplift ensued, possibly accompanied by regional metamorphism and granite intrusion in the Mary Kathleen Shelf. The succeeding sediments of Stage 3 are separated from older rocks by a region-wide unconformity.
Stage 3 : Parageosynclinal stage: regressive sandstone transitional to siltstone and shale. Rapid erosion following uplift resulted in widespread conglomerate-sandstone deposition across the Orogen, in linear basins, possibly parageosynclines, apparently developed cratonwards of a slowly rising foreland to the east. Basal fluvial, delta-fan, or shallow-shelf sandstones, which display pronounced thickness variations across the Quilalar and Mount Gordon Arches, fine upwards into a fine sandstone and grey-green and carbonaceous siltstone sequence which thins eastwards, possibly a distal flysch facies (Mount Albert Group, Surprise Creek Formation). Stage 4: Cratogenic orthoquartzite transitional to fine-grained flysch and platform carbonates. This stage encompasses the Mount Isa and McNamara Groups, hosts for major Pb-Zn (Mount Isa, Hilton, Lady Loretta) deposits in the region. The depocentre of sedimentation appears to have migrated westwards or northwestwards with time, in response to the postulated slow emergence of a metamorphosed foreland t o the east (Fig. 4.14).Mount Isa Group rocks extend for 10 km east of the Quilalar Arch, but have not been recorded from the Mary Kathleen Shelf. McNamara Group rocks extend across the Lawn Hill Platform, westwards from the Mount Gordon Arch. A basal supermature conglomerate-orthoquartzite marine-shoreline blanket, subjected to periodic exposure, overlies older rocks both conformably and unconformably. It grades upwards into fine distal turbidities -- carbonaceous and dolomitic siltstones deposited in extensive backshore, lagoonal, and lacustrine environments. In the Mount Isa Group, overlying sequences of mainly flat-laminated dolomitic and pyritic siltstones with chert and tuff markers are host to the Mount Isa Cu-Pb-Zn deposits; time equivalents in the McNamara Group comprise a subtidal to supratidal, highly stromatolitic carbonate facies, which thins markedly onto the Mount Gordon Arch. Younger sequences are preserved only on the Lawn Hill Platform (Fig. 4.11 and 4.14),suggesting a further northwestward migration of the depocentre. Shallow-shelf carbonates and quartzites were followed by further turbidites and black shale, grading up into shallow-water micaceous sandstone.
269 S o u t h Nicholson Basin Following and/or during the final regional deformation of the Mount Isa Orogen and Lawn Hill Platform, alternating sandstones and micaceous siltstones (South Nicholson Group, Pilpah Sandstone), which resemble the Roper Group of the McArthur Basin, and are also interpreted as a sub-stable shelf association (or molasse facies?), were deposited on top of the Lawn Hill Platform. Palaeocurrents and facies clearly indicate derivation from the uplifted Mount Isa Orogen t o the east. Banded pisolitic iron ores (Constance Range Fe; Harms, 1965) were again deposited in marginal-marine environments. Igneous, metamorphic and structural style Igneous rocks. The Kalkadoon Granite is granitic to granodioritic in composition. Possibly comagmatic volcanics of the Tewinga Group (Leichhardt Metamorphics) are rhyolitic t o dacitic. Younger volcanics of the Tewinga Group (Argylla Formation) are generally more alkali-rich, and show enrichment in total Fe, TiOz, K20, Nb, Th, Y, and %, and depletion in CaO, A1203 and Sr compared with the Leichhardt Metamorphics (Wilson, 1978). Younger granites in the Mary Kathleen Fold Belt and Leichhardt River Fault Trough are silica and alkali-rich types, commonly highly potassic (more than 6% K,O), but with associated albitite phases. Basalts of the deep-rifting stage are continental tholeiites in the west (Eastern Creek Volcanics), but show increased oceanic tholeiite affinity in the east (Soldiers Cap Group), which suggests that continental crust may thin significantly from west t o east across the Mount Isa Orogen (Glikson et al., 1976). Most dolerite dykes are tholeiitic. Gabbroic intrusions from near Mary Kathleen are mainly K-rich subalkaline tholeiites, and most resemble the low-Mg Karroo Basalt magma type. All basic rocks show a trend of iron enrichment with increasing differentiation. The distinctly bimodal geochemistry of the plutonic and volcanic rocks may indicate a mainly extensional crustal regime (Martin and Piwinskii, 1972). Metamorphism. At least three and possibly five temporally distinct metamorphic episodes are evident. Episode 1 (about 1870 Ma) is related t o the batholithic Kalkadoon Granite; enclaves of grey dacitic volcanics are progressively metamorphosed t o grey gneiss and migmatite. Episode 2 is spatially related t o the Ewen Granite, and predates the tholeiitic volcanism of the Eastern Creek Volcanics. Basal Haslingden Group and/or Tewinga Group rocks were converted t o garnet and cordierite schist. Episode 3 is related t o the Sybella Granite, which metamorphoses the Eastern Creek Volcanics t o amphibolite grade, but which predates the Mount Isa Group. It may be of similar age to the Episode 4 metamorphism t o the
270
east, which was the main metamorphism t o affect the Mary Kathleen and Soldiers Cap Groups in the Mary Kathleen Fold Belt, and which was spatially related to high-level granites in the Fold Belt. Some of the latter, however, may post-date the main metamorphism. These metamorphic episodes were mainly a low pressure-high temperature type, attaining amphibolite facies grade; andalusite and cordierite are widespread, and staurolite, garnet, diopside, wollastonite, forsterite, and vesuvianite are present in rocks of appropriate composition. Polymetamorphism is evident, especially in the east, where sillimanite, and locally kyanite, replace andalusite and cordierite. Episode 5 metamorphism appears t o be related to the major conjugate faulting in the Wonga belt, near Mary Kathleen, where kyanite-corundumcordieriteanthophyllite, kyanite-chlorite, and cordierite-biotite assemblages have been recorded from the fault zones. This deformation may be related to the last regional deformation which affected the Mount Isa Group, to the west.
Structure. The growth faults which influenced sedimentation have probably been rejuvenated during later deformation. The oldest post-depositional structures are north-trending, apparently high-angle reverse faults, which show east-block-up and west-block-up movements, east and west of the Kalkadoon-Leichhardt Block, respectively. In the west, this system is followed by curved east-trending slices (“spoonfaults”), which are low-angle normal faults with north-block-up sense of movement (Dunnet, 1976b). All these structures were later displaced by a system of conjugate wrench faults; some of the latter, mainly northeast-trending, show up t o 25 km of rightlateral movement. Folding has accompanied all phases of faulting. The Soldiers Cap Group and adjacent rocks, in the far east of the Mary Kathleen Fold Belt, display multiple superposed folding and overthrusting towards the west, possibly reflecting the postulated foreland uplift (Fig. 4.14). Westwards of the Soldiers Cap Group, a complex pattern of basins and domes is preserved in the Mary Kathleen Fold Belt and on the Lawn Hill Platform and has probably resulted from inhomogeneous flattening during compression. These folds predate both the “spoon” and the wrench faulting. All the post-depositional fault and fold features are consistent with eastwest or craton-directed compression, in contrast t o the crustal extension postulated for most of the period of sedimentation and igneous activity. M inera1 deposits Mineralization in the Mount Isa Orogen and Lawn Hill Platform is dominated by the large shale-hosted Pb-Zn-Ag deposits at Mount Isa, Hilton, and Lady Loretta, and the brecciated sediment-hosted Cu deposits at Mount Isa and Mammoth. In addition, there are a number of stratiform Cu-Zn and Pb-
271 Zn deposits of the Broken Hill type; pyrometasomatic, pegmatite-related, and sedimentary U deposits, and innumerable small fissure Cu deposits.
Lead-zinc deposits. Massive Pb-Zn-Ag deposits formed during deposition of the Mount Isa Group, in the Leichhardt River Fault Trough. The Mount Isa deposit (reserves 56 m.t. at grades of 7% Pb, 6% Zn, 149 g/t Ag) and the nearby Hilton deposit (reserves 36 m.t. at 7.7% Pb, 10% Zn, 180 g / t Ag) consists of well-bedded sulphide ore layers hosted by pyritic and dolomitic siltstone and shale (Mathias and Clark, 1975), lacking obvious shallow-water features, and deposited adjacent to possible growth faults or fractures. Discovery of pseudomorphed sulphate evaporites within the deposit indicates possible ore formation from circulating metal-rich brines in an evaporite basin (McClay and Carlile, 1978). The McArthur and Mount Isa deposits show many similarities: they are about the same age, they occur in very similar host rocks with distinctive tuff layers, and they are located within or at the margins of major intracontinental or continental margin rift zones. In contrast to the unmetamorphosed McArthur deposit, greenschist facies metamorphism at Mount Isa has resulted in coarsening of galena and sphalerite, remobilization of sulphides, and the formation of pyrrhotite (Williams, 1980). A t Lady Loretta, on the Lawn Hill Platform, bedded Pb-Zn-Ag sulphides (reserves 8.7 m.t. at 6.7% Pb, 18.1% Zn, and 109 g / t Ag) are contained within a massive pyrite body, which passes laterally and vertically into carbonaceous shale and dolomitic siltstone of the McNamara Group (Loudon et al., 1975). Baryte lenses and potassium-rich siltstones are present near the ore zone. Isotopically, the deposit resembles the McArthur deposit (Carr and Smith, 1977). Stratiform Zn-Cu and Pb-Zn deposits, although presently subeconomic, are an important class of deposits recently discovered in the Soldiers Cap Group, in the far east of the Mary Kathleen Shelf. The Pegmont--Squirrel Hills type (Locsei, 1977) contains Pb-Zn in a magnetite-garnet-apatiteolivine-pyroxene-gahnite-hornblende-epidote-rich iron formation, within a sequence of amphibolite, feldspathic quartzite, schist, and pegmatite. The Monakoff Zn-Cu type contains sphalerite and chalcopyrite in a bedded baryte-fluorite-carbonate-garnet gangue, within a sequence of qu&zmagnetite schist, pillowed metabasalt, and garnet-amphibolite. The Pegmont-Squirrel Hills type closely resembles the Pb-Zn-Fe-iron formation association at Broken Hill (Richards, 1965). The Dugald River Zn-Pb deposit (Whitcher, 1975) appears to be the most extensively mineralized example of a series of Zn-rich euxinic sub-basins formed within the Mary Kathleen Shelf (Derrick et al., 1977). The deposit contains an ore resource potential of 40 m.t. at 10--12% equivalent (C.R.A., 1977). It is contained in a lode of graphitic slate, interlayered with impure dolomite and limestone containing anhydrite pseudomorphs, from near the top of the Mary Kathleen Group and metamorphosed t o amphibolite facies.
272
Copper deposits. Cu mineralization within brecciated sedimentary rocks is widespread in the Leichhardt River Fault Trough and Lawn Hill Platform. The enormous copper deposits in “silica-dolomite” breccia of the Urquhart Shale, in the Mount Isa Group at Mount Isa (reserves 126 m.t. at 3.2% Cu) are spatially related to the Ag-Zn-Pb deposits, and have been considered t o be syngenetic (Bennett, 1970; Mathias and Clark, 1975; Dunnet, 1976b) or epigenetic (Murray, 1961; Smith and Walker, 1971). Recent work tends to support Smith and Walker who suggested that Cu was mainly derived by leaching and transportation from older altered basalts or greenstone (Eastern Creek Volcanics), faulted against the Mount Isa Group. Circulating metasomatic fluids, adjacent t o the greenstoneshale contact, have silicified the dolomitic shale and siltstone and deposited massive Cu sulphides. Cu grades and degree of silicification decrease away from the faulted greenstone contact (Mathias and Clark, 1975). The recent discovery of anhydrite and gypsum pseudomorphs (McClay and Carlile, 1978) may indicate that the “silica--dolomite” was a more evaporative facies of the Urquhart Shale. Cu mineralization at Mammoth (reserves 6.1 m.t. at 3.3% Cu) is present as chalcocite-bornite-chalcopyrite--pyrite infilling of brecciated red sandstone and siltstone of the upper Myally Subgroup. Pyrite is syngenetic, but copper mineralization is late diagenetic or related t o late tectonic activity (Mitchell and Moore, 1975; Holyland et al., 1976). The deposit is overlain by a regolith at the base of the Surprise Creek Formation, and is located along the flanks of the Mount Gordon Arch. Most small copper deposits in the region result from hydrothermal activity related to dolerite and granite intrusion, metamorphism, and deformation (Wilson et al.,1972; Krosch and Sawers, 1974; McKay and Finlow-Bates, 1977; Williams, 1980). They are common in both basement areas and in the Mary Kathleen Fold Belt, and characteristically are small but of high grade (570 to 20% Cu). Many small deposits are stratabound within pelitic and arenaceous metasediments of the middle Corella Formation. Cu mineralization in trachytic to rhyolitic breccia pipes and dykes occurs in the Lawn Hill Platform, associated with a period of granite intrusion or Fiery Creek volcanism. It resembles the Redbank-type mineralization in the McArthur Basin. Uranium deposits. Some U occurs in tuffaceous siltstones and quartzite interbeds in the Eastern Creek Volcanics, usually adjacent to shears or faults. The uraniferous phases are zircon and brannerite, and deposits such as Skal and Valhalla each contain about 3500 tonnes of U308, at grades of 0.13 to 0.2% (Brooks, 1975). Mary Kathleen, the only U-producer (reserves 7000 tonnes U308 at 0.12%) is located in the Mary Kathleen Fold Belt: the ore is a network of uraninite-allanite veins in a massive skarn, which has replaced arkosic conglomerate and calc-silicate granofels of the Corella Formation; the source of metasomatic fluids and uranium is thought to be the
273 nearby Burstall Granite and associated alkali rhyolite dykes (Hughes and Munro, 1965; Hawkins, 1975; Derrick, 1977). Iron deposits. The subeconomic Constance Range Fe deposit, similar t o the Roper River Fe deposits of the McArthur Basin, consists of pisolitic hematitesiderite-chamosite deposited in a shallow-water marginal-marine environment (Harms, 1965). Northeast Queensland Precambrian Province (Principal references: Black et al., 1979; Bain and Withnall, 1 9 8 0 ; Withnall e t al., 1980).
The interrelated Georgetown and Coen Inliers (Fig. 4.1), and the smaller Dargalong and Yambo Inliers, are the exposed parts of the Northeast Queensland Precambrian Province (Withnall et al., 1980; Fig, 4.4).Early or Mid-Proterozoic sedimentary and volcanic rocks underwent repeated cycles of deformation, metamorphism, and granite emplacement during the Middle and Late Proterozoic, and then again during the Palaeozoic under the influence of the adjacent Tasman Orogenic Province. The province is separated from the Tasman Orogenic Province by the fundamental, transcurrent Palmerville Fault (De Keyser, 1963). It is unconformably overlain by Mesozoic rocks of the Trans-Australian Platform Cover, which obscure its relationships with Proterozoic domains farther t o the west (Fig. 4.1), but similar gravity trends t o those of the outcropping inliers may be traced beneath the cover, until truncated by the Mount Isa Orogen (Fig. 4.4; Wellman, 1976; Plumb, 197913). A thick pile of shallow-water arenite, lutite, and locally calcareous lutite, with intercalated basic volcanics and sills (Etheridge Group, Georgetown Inlier), accumulated during the Early or Mid-Proterozoic, in what may have been either a rifted marine basin or a marginal sea. 87Sr/86Srinitial ratios of the Etheridge Group metasediments suggest that the sediments in the Georgetown Inlier are not very much older than 1600 Ma. A second sequence (Yarman, Malacura Formations) was deposited unconformably on the Etheridge Group between the first and second deformations (D1,D2). Five separate episodes of deformation and metamorphism have been identified and precisely dated in the Georgetown Inlier. Emplacement of extensive granites in composite batholiths accompanied most of these events. Folding is superposed about alternate easterly and northerly trending axes. The first two deformations produced tight to isoclinal folds about largely east- (D1)and north-trending (D2)axes. They were accompanied by prograde metamorphism, ranging from greenschist to high-amphibolite and marginal granulite facies, of the low pressure-intermediate type. Grades increase, overall, from west t o east. Rb-Sr total-rock isochrons indicate ages of 1570 f 20 Ma for D1, and 1469 k 20 Ma for Dz (Black et al., 1979). After D2, mildly deformed rhyolitic ignimbrites and granites, overlain by
274 quartzose sandstone, were emplaced into the western Georgetown Inlier; these reflect an early stage of progressive cratonization (transitional tectonism) from west to east (Plumb, 1979b). The Croydon Volcanics yield a Rb-Sr total-rock isochron of 1399 t 75 Ma, and the comagmatic Esmeralda Granite is 1381 + 99 Ma (Black, 1973). Subsequent deformations have only been identified farther t o the east. D3 produced tight folds about east-trending axes and retrogressive greenschist facies metamorphism; Rb-Sr total-rock isochrons of 967 t 28 Ma have been obtained (Black et al., 1979). D4 is thought t o be related to SiluroDevonian deformation in the adjacent Tasman Orogenic Province, and is reflected in open folds coaxial with D3, retrogressive (greenschist) metamorphism, possible granite emplacement, and widespread mineral overprints around 400 M a (Richards et al., 1966; Cooper et al., 1975; Black et al., 1979). D5 is represented by broad warping of Proterozoic strata about northtrending axes, which accompanied the emplacement of Permo-Carboniferous granites and acid volcanics into cauldron subsidence structures, largely controlled by pre-existing Proterozoic features (Branch, 1966) - part of the transitional tectonism in the adjacent Hodgkinson domain (GSA, 1971). The Georgetown Inlier contains a wide variety of mineral deposits. The majority (U-F-Mo, porphyry Cu-Mo, Sn, W, Cu-U, base-metal veins) are clearly related t o the late Palaeozoic magmatism. Small stratabound Cu and Pb-Zn deposits occur in Proterozoic strata in many places and may be essentially syngenetic, although they may have been remobilized by late metamorphic and magmatic events. Close spatial relationships exist between numerous small vein Au deposits and interacting Proterozoic and Late Palaeozoic components of the regional geology : a multistage process is envisaged, in which originally syngenetic Au has been upgraded and relocated during subsequent metamorphic and magmatic processes (Bain and Withnall, 1980). Arunta Inlier The Arunta Inlier covers about 200,000 km’, along the southern edge of the North Australian Craton. It is a typical example of the intracontinental polymetamorphic mobile belts which are so characteristic of Proterozoic terrains, with their problems of evolution, origin, and interpretation. The stratigraphy of the Arunta Inlier is summarized in Table 4.VIII and the tectonics, structure, and metamorphism in Fig. 4.15. Tectonic setting The Arunta Inlier is the exposed part of a polymetamorphic belt adjacent t o the southern margin of the North Australian Craton (Figs. 4.1 and 4.2). Early Proterozoic rocks, apparently continuous with those of the North Australian Orogenic Province, were repeatedly metamorphosed, deformed,
275 TABLE 4.VIII Summary of stratigraphy and tectonic history, Arunta Inlier Event and approximate age
Northern Zone
Central Zone
Southern Zone
Alice Springs Orogeny 360-320 Ma
Overthrusting, retrogression
Faulting; local metamorphism u p t o amphibolite. Possibly Yambah Event. 500-350 Ma
Overthrusting; local greenschist metamorphism
______
PLATFORM COVER DEPOSITION - 900-300 Ma - AMADEUS-NGALIAGEORGINA BASINS - - - - - - .- _ _ - --- - - U n c o n f o r m i t y - ---- - - - - - - - -- - Post-tectonic intrusives 7 3 0 Ma Mud Tank Carbonatite 900 Ma Rare dolerite dykes; Rare dolerite dykes Dolerite dykes, metatroctolite plug gabbro, dacite
O r m i s f o n Event 1100-1 000 Ma
Local granite
Local granite
M o r d o r Igneous Complex 1185 Ma
Metamorphism t o amphibolite; migmatite, granite Potassic-ultramafic intrusives; unmetamorphosed
Metamorphism and igneous intrusions 1500-1400 Ma
Possible greenschist metamorphism. Extensive granite - “S-type”.
Local granulite metamorphism possibly migmatite - Woolanga Event
Not evident
Chewings E v e n t 1700-1550 Ma
Granite, pegmatite; metamorphism n o t evident
Local metamorphism t o amphibolite a t southern margin
Metamorphism t o amphibolite; “I-type’’ granite; local diorite, dolerite, and ultramafic
Strangways Event 1800-1750 Ma
Metamorphism t o granulite; dolerite; intrusive charnockite; granite
Metamorphism t o granulite ; migmatite ; anatectic granite
N o t evident
Division 3
Quartzite and metasediments - e.g., Reynolds Ra. Gp.
Quartzite; pelitic gneiss and schist; calc-silicates - e.g., Mendip Mets.
Quartzite; calcareous and pelitic rocks; quartzo-feldspathic gneiss - e.g., Iwupataka Met. Cplx. -Probable ---unconformity
__-_-_--__
U n c o n f o r m i t y - - - - - Inferred - - unconformity
276 TABLE 4 .VIII (continued) Event and approximate age
Division 2 2A
Northern Zone
Central Zone
Southern Zone
Slate, schist, metasandstone, amphibolite - e.g., Lander Rock Beds Not evident
Pelitic and calcareous Metasediments, calcrocks; amphibolite - silicate rocks; o r t h o e.g., upper Harts gneiss. e.g., Hayes Range Gp. Met. Cplx. 2B Layered granitic Layered granitic gneiss - e.g., Entia granodioritic gneiss; Gneiss; lower Harts amphibolite Ra. Gp. -~ _ _ _ _ .- _ . - _ - _--.---_ . - _ _-_ - -_ - - -_ - ----Contact relationFaulted; disrupted b y 2B locally unconform- l A / l B ouerthrust 2B ship granite able o n 1 A ; otherwise at northern margin faulted _.______-_._. - - - - - - - - - - -- - - - Division 1 1.4 N o t evident Calcareous and pelitic N o t evident rocks; quartzofeldspathic gneiss - e.g., Cadney Mets.
IB
Felsic and mafic granulites in Anmatjira Range
Felsic and mafic granulites; subordinate metasediments e.g., Strangways Met. Cplx.
and intruded by granite throughout the Mid-Proterozoic. The Inlier is overlain by Adelaidean-Palaeozoic platform covers of the Amadeus-NgaliaGeorgina Basins, and was reactivated by mid-Palaeozoic overthrusting and retrogressive metamorphism (Table 4.VIII). The Musgrave Block (Fig. 4.1) (Rutland, this volume, Chapter l ) , has many similarities t o the Arunta Inlier, and it is quite likely that they link up beneath the Amadeus Basin t o form an intracontinental mobile belt between the North Australian and Gawler Cratons (Figs. 4.2 and 4.4).
St ra t igrap hy The numerous rock units mapped within the Arunta Inlier have been grouped into Divisions 1, 2, and 3 (Shaw and Stewart, 1975b; Stewart and Warren, 1977). Unconformities are thought t o separate the Divisions, but many contacts are now faulted (Table 4.VIII; Fig. 4.15b). The divisions are based on rock types; they may represent chronological or stratigraphic equivalents, but this cannot be proved. The original sedimentary succession shows an overall decrease in volcanism and an increase in sedimentary maturity with time.
L
I
A
I
YORTHERN
ZCNE
I
CENTRAL ZONE
I
I
I
Metamorphic
lSOUTHERN
of
map
LONE1
I
~
Arunta
I
lniier
-r,
\ L SOUiHERN
6 Relationship of the informal Stratigraphic divisions of the Arunta Block
ZONE
C Distribution of Tectonic Zones
Fig. 4.15. Metamorphic rocks, stratigraphic relationships, and tectonic zones, Arunta Inlier.
27 8 The lower part of Division I (1B) comprises granulite facies rocks of felsic, mafic, and pelitic composition, the first two types including chemical equivalents of volcanic rocks; these rocks are interpreted as forming basement to all the younger successions. They are locally succeeded (Division l A ) , by a sequence of metapelitic and cordierite-rich gneisses and calc-silicate rocks of the amphibolite facies. Metagabbro sills and dykes intrude the Division. Division 2 comprises metapelite and arenite, and minor intermediate and basic flows or sills, now metamorphosed from greenschist t o granulite facies. Division 3 consists of metamorphosed shelf sediments - quartzite, shale, and carbonate rocks - and some acid porphyry sills, mostly in the greenschist facies but locally reaching granulite facies. Division 1 is tentatively correlated with the orogenic domains of the McClintock Subprovince, and Division 2 with the orogenic domains and Division 3 with the transitional domains of the Palmerston Subprovince (Plumb, 1979b).
Tectonic zones Three tectonic zones are recognized in the Arunta Inlier, each with a distinctive assemblage of rocks, structures, and tectonic history (Table 4.VIII; Fig. 4.15). The Northern Zone contains mainly low-grade metasediments of Divisions 2 and 3, which appear t o merge northwards into those of the North Australian Orogenic Province. A widespread weak metamorphism and granite emplacement event about 1500 Ma old characterizes the zone. Granites are mostly “S-type”. The Central Zone contains the main exposures of Division 1, and only minor exposure of Divisions 2 and 3. Almost all of the rocks lie in the granulite or amphibolite facies. It is the main area of development of the oldest (1800-1 750 Ma) readily recognized metamorphism (Strangways Event) in the Inlier, and evidence of older events (>1800 Ma) is also present. Granites are mostly anatectic. The boundary between the Central and Northern Zones is marked by a belt of high-level granites (Fig. 4.15A). The Central Zone represents a deeper crustal level than the Northern Zone. The Southern Zone is characterized by intense young (1050 Ma) migmatization and reactivation (Ormiston Event), superimposed on 1700-1550 Ma amphibolite metamorphism (Chewings Event). As such, it is closely related to the northern Musgrave Block, t o the south of the Amadeus Basin (Rutland, this volume, Chapter l), and similar gravity trends link both areas beneath the Amadeus Basin (Fig. 4.4). Most granites analyzed so far are “I-type”. The gneisses and migmatites of the Southern Zone are separated from the upthrust granulites of the Northern Zone by a narrow east-trending, north-dipping zone of complex intense mechanical deformation and flattening (Redbank Z o n e ) ; the Redbank Zone formed in the amphibolite facies (Chewings Event?), before the Ormiston Event (Marjoribanks and Black, 1974).
279
Description of zones Northern Zone. At the poorly exposed northern margin of the Zone, the metasediments appear t o merge with those of the Tennant Creek and The Granites-Tanami Inliers: Division 2 rocks have been correlated with the Warramunga Group and Tanami Complex, and Division 3 rocks with the Hatches Creek Group and the Pargee Sandstone/Mount Winnecke Formation, of the Tennant Creek and The Granites-Tanami Inliers (Shaw and Stewart, 1975a). A narrow northwesterly trending upfaulted block of high-grade metamorphic rocks and charnockites, in the Anmatjira Range (A in Fig. 4.1 5A), apparently metamorphosed during the Strangways Event, consists of mainly Division 1 rocks; these rocks may correlate with the Halls Creek Group (Plumb, 1979b). A widespread weak metamorphism overprinted all the Early Proterozoic rocks during a significantly younger event, provisionally dated at about 1500 Ma (L.P. Black, pers. commun., 1978); the full extent of this event in the poorly exposed northern area is unknown, The Zone was extensively intruded by granites throughout the interval 1700-1400 Ma; these are concentrated in the higher crustal levels adjacent to the Ngalia and Georgina Basins. Post-tectonic granites of similar age also occur in the adjacent Tennant Creek and The Granites-Tanami Inliers. Major folds and faults in the Northern Zone form a westerly-trending arc, convex t o the north (Fig. 4.15A). The rocks of the Northern Zone form basement t o the Ngalia, Georgina, and Wiso Basins, and the northern margin of the Ngalia Basin became the site of major southward-directed overthrust faulting, and the southern margin of the Georgina Basin the site of northeastward-directed overthrust faulting, during the Carboniferous Mount Eclipse (or Alice Springs) Orogeny (Walter, 1980; Wells and Moss, in press). Central Zone. The Central Zone comprises mainly rocks of Division 1 felsic and mafic granulites, intercalated with and overlain by pelitic, quartzofeldspathic, and calcareous metasediments, and includes meta-anorthosite in its lower part (Table 4.VIII). Anatectic granite occurs as a few small bodies. Rb-Sr total-rock ages of granites and high-grade metamorphic rocks are mostly about 1800--1750 Ma (Black, 1975; Iyer et al., 1976) - the Strangways Event. Iyer and others have identified a second, locally developed, high-grade metamorphism (Woolanga Event), which they have dated by RbSr total-rock isochrons at about 1440-1400 Ma. The Chewings Event (1680--1580 Ma) is only registered locally, adjacent t o the boundary with the Southern Zone (Redbank Zone). The Carboniferous Alice Springs Orogeny is presented by broad belts of retrogression, shear zones, and faults. Older K-Ar mineral ages (500-350 Ma) may possibly represent an older phase of the Alice Springs Orogeny - - the Yambah Event of amphibolite
280 facies (Shaw and Langworthy, in prep.) (see Southern Zone discussion). Three distinct phases of folding are evident in the Central Zone: the oldest has northerly striking, axial-plane traces; the second and main phase has eaststriking axial-plane traces; the youngest phase is only locally developed, it commonly has northerly axial-plane traces, and it is thought to correspond to the peak of the Strangways Event. The post-tectonic Mud Tank Carbonatite was emplaced adjacent t o a major cross-cutting lineament of the Central Zone, during development of the epicratonic Amadeus Basin (crystallization age 732 t- 5 Ma, U-Pb, zircon, Black and Gulson, 1978).
Southern Zone. The Southern Zone consists of extensively migmatized gneisses of granitic t o granodiorite composition (Division 2), overlain by metasediments and quartzite (Division 3); metamorphic grade decreases upwards from amphibolite t o greenschist facies (Table 4.VIII). Two major Proterozoic deformations and metamorphisms, and one during the Palaeozoic, may be identified. The Chewings Event produced isoclinal similar folds with a strong axialplane foliation and lineation, and was accompanied by amphibolite facies metamorphism, dated by a Rb-Sr total-rock isochron at 1585 70 Ma (Marjoribanks and Black, 1974; Marjoribanks, 1975). Metamorphism and syntectonic granite in the eastern part of the zone, dated by a Rb-Sr totalrock isochron at 1682 ? 24 Ma (Armstrong and Stewart, 1975), may belong to the same or a slightly earlier event. The Ormiston Event characterizes the Southern Zone. It produced intense migmatization and granite intrusion in the amphibolite facies, accompanied by the development of large east-plunging folds. A combined mineral-total rock Rb-Sr isochron indicates an age of 1053 ? 50 Ma for the migmatites (Marjoribanks and Black, 1974), while granites were emplaced about 980 Ma ago (Allen and Black, 1979). The unmetamorphosed “post-tectonic” Mordor Igneous Complex, a highly differentiated potassic-ultramafic intrusive with kimberlitic affinities, was emplaced 1185 t- 90 Ma ago (Langworthy and Black, 1978), apparently during a tectonic lull between the Chewings and Ormiston Events. Post-tectonic dolerite dyke swarms in the Southern Zone, and small gabbro bodies in the Redbank Zone, were emplaced about 900 Ma ago, foliowing which the platform cover of the Amadeus Basin was deposited above the Southern Zone (Black et al., 1980). The area was reactivated again during the Carboniferous Alice Springs Orogeny, when three thrust-nappe complexes, involving both basement and cover rocks, were thrust several kilometres over the northern margin of the Amadeus Basin (Stewart, 1967; Forman, 1971; Shaw et al., 1971; Marjoribanks, 1975, 1976). The cover rocks give K-Ar mineral ages about 358--322 Ma, and zones of retrogression in the basement give Rb-Sr total-rock mineral isochrons about the same age
*
281 (335--310 Ma). However, K-Ar mineral ages in retrogressed basement rocks are older, due to excess argon (Armstrong and Stewart, 1975), and are similar to the age of the supposed Yambah Event of the Central Zone. Me tam o rp h ism Most regional metamorphism in the Arunta Inlier appears t o belong t o the low pressure-intermediate facies series. Cordierite-garnet-sillimanite is widespread in high-grade terrains, and sapphirine is locally important. Staurolite is common in amphibolite facies terrains. Many of the metamorphic events involved more than one phase. Some metamorphism is syntectonic (e.g., Chewings Event). Others (e.g., Strangways and Ormiston Events) involve considerable late or post-tectonic migmatite emplacement. Kyanite is developed in the higher-stress environments of the retrogressed shear zones of the Alice Springs or supposed Yambah Events. Mega-struc tu re The structure of the Arunta Inlier is dominated by large west t o northwest-trending belts of overthrusting, shearing, mylonitization, and retrogression, which commonly separate the principal tectonic zones or zones of markedly different metamorphic grade (Fig. 4.15). These deformed belts originate from the earliest history of the inlier (Shaw and Stewart, 1975a), but they were also the loci for most of the overthrust faulting and retrogression which uplifted the Arunta Inlier and its overlying cover to its present attitude, during the mid-Palaeozoic Alice Springs-Mount Eclipse Orogeny. Forman and Shaw (1973) postulated that these overthrusts pass through the crust into the mantle, but Anfiloff and Shaw (1973) have shown that the gravity features of Central Australia may also be quantitatively explained by heterogeneities in the upper crust. Central Australia contains the largest gravity anomalies in Australia: for example, the Central Zone of the Arunta Inlier corresponds t o a latitudinal Bouguer Anomaly ridge up t o +45 mgal, flanked by parallel troughs down to -100 mgal, above the Northern and Southern Zones; another ridge and trough of smaller magnitude overlies the northern edge of the Northern Zone. These features continue beneath the Amadeus Basin, and are mirrored by similar large gravity troughs and ridges above the Musgrave Block (Langron, 1962; Lonsdale and Flavelle, 1963; BMR, 1976). All these anomalies terminate abruptly against two northerly trending lineaments in the east and west, interpreted as the sites of dislocation between which northksouth compression of the crust was largely confined (Anfiloff and Shaw, 1973). These linear gravity features may be interpreted in terms of major density contrasts, between 0.07 and 0.20 g/cm3, across the boundaries between the tectonic zones (Anfiloff and Shaw, 1973; Mathur, 1976; Wellman, 1978). The various interpretations differ in the amount of mass imbalance ascribed to distortions in the mantle-crust boundary.
282 It is apparent, from both the surface geology and the magnitude of the Bouguer anomalies, that major lateral density discontinuities which involve the upper crust, and possibly also the lower crust, separate zones of different overall composition, structure, and possibly age in central Australia. Gibb and Thomas (1976) suggest that similar gravity features in the Canadian Shield, which correspond t o major province boundaries, formed in response to convergence, collision, and suturing of cratonic blocks. Support for this concept is provided by preliminary data on granites in the Arunta Inlier (A.J. Stewart, pers. commun., 1979): 1700 Ma-old granite and diorite in a klippe near the boundary between the Central and Southern Zones resemble “I-type” granites of igneous source; many of the granites in the northern zone appear similar to “S-type” granites, considered to be derived from anatexis of metasediment at much higher crustal levels. However, the general pattern of tectonic evolution and palaeomagnetism in Australia suggest that the major Precambrian cratons in Australia have remained in essentially their present relative positions since at least the Early Proterozoic, and that most of the mobile belts are ensialic (McElhinny and Embleton, 1976; Plumb, 1979a).
Tectonic models Several models may explain the differences in crustal composition and history of the three zones in the Arunta Inlier: (1)They may be the result of interaction between continental and oceanic crust about 1800-1700 Ma ago. However, the high proportion of granites in the southern zone, and the lack of ophiolites, make this model unlikely. (2) They may be the result of intracontinental deformation, involving the incipient underthrusting of the crust into the mantle along a zone of high geothermal gradient (cf. Forman and Shaw, 1973). (3) Dense basic rocks, now seen as basic granulite, may have been emplaced into the Central Zone by a process of incipient rifting, followed by uplift and overthrusting, similar t o the ensialic rift model of Windley (1973). (4)The Southern Zone may correspond to a younger geosyncline, developed against an older basement corresponding mainly to the Central Zone (Division 2 rocks appear to overlie Division 1 in the overthrust boundary zone). Deep subsidence of this geosyncline resulted in the accumulation of sufficient radiogenic heat t o melt parts of both the lower and upper crust (Schruiling, 1972a,b). In this model the Northern Zone may also be interpreted as another geosyncline, originally overlying and now faulted against basement of the Central Zone but, in contrast to the Southern Zone, its position closer t o the North Australian Craton resulted in only moderate subsidence, less volcanic rocks, and generally low-grade regional metamorphism. Possibly each of the last three models apply in part. Virtually all the gravity interpretations require that the lateral density contrasts persist down t o the lower crust or mantle. The magnitude of the density contrasts requires
283 more basic--ultrabasic rock, at all depths, in the Central Zone than in the adjacent zones. The gravity anomalies have clearly been accentuated by thrusting and uplift during the younger Ormiston and Alice Springs events, as well as the older events. The Division 3 rocks, at least, seem to correlate well across all zones. Therefore all events, from at least the Strangways Event on, represent ensialic reactivation of an intracontinental belt. The Southern Zone “geosyncline” may well continue southwards beneath the Amadeus Basin, to the northern Musgrave Block, and thus represents the main reactivation zone between the North Australian and Gawler Cratons (Fig. 4.2), bounded by large zones of thrusting and basic rock emplacement, which extend to the crust or mantle - the Redbank Zone (Arunta Inlier) and the Woodroffe Thrust (Musgrave Block). The post-Amadeus Basin Alice Springs Orogeny represents a higher-level and more brittle deformation than the earlier events, involving intracontinental overthrusting, metamorphism, and reactivation of older deformed zones, up t o 700 Ma after “cratonization”. LATE: PROTEROZOIC PLATFORM COVERS
Amadeus, Ngalia and Georgina Basins (Principal references: Wells et al., 1970; Preiss e t al., 1978; Walter, 1980; Wells and Moss, in prep.)
Following cratonization of the Arunta Inlier and the Musgrave Block by the Ormiston Event (Arunta Inlier) and by the Musgravian Orogenic Cycle (Musgrave Block) up t o 10 km of Adelaidean and Palaeozoic sediments were deposited, with several breaks, in the Amadeus, Ngalia, and Georgina Basins (Figs. 4.1 and 4.3). At the close of the Proterozoic, the Petermann Ranges Orogeny (Forman, 1966) uplifted the Musgrave Block: a large thrust-nappe complex, incorporating both basement and cover rocks, extends several kilometers into the southwestern corner of the Amadeus Basin; the effects of this deformation decrease northwards. During the Carboniferous, the Arunta and Musgrave Blocks were uplifted by the Alice Springs--Mount Eclipse Orogeny, when thrust-nappe complexes, thrust sheets, and reverse faults developed at the margins of the present basins (Wells et al., 1970; Milton and Parker, 1973; BMR, 1978; Wells and Moss, in prep.). This event brought the development of the Central Australian Platform Cover, in central Australia, to a close. The Amadeus Basin is well known for its assemblages of stromatolites (Walter, 1972), and for its exceptionally well-preserved microfossils (Schopf, 1968; Schopf and Black, 1971; D.Z. Oehler, 1976, 1977). The Adelaidean stratigraphy and correlations of the basins are summarized in Table 4.IX and Fig. 4.16.
284 TABLE 4.IX Summary of Adelaidean (Late Proterozoic) stratigraphy, Amadeus-Ngalia-Georgina Basins Unit
Rock types and thickness
Remarks
Basal diamictite, arkose, and “cap” dolomite (Olympic Fm.); siltstone, shale, red sandstone (stromatolitic) dolomite. 500-2750 m
Unconformable between L. Cambrian and Aralka Fm. or equiv. Thickens t o SW. “Ediacara” faunaArumbera Sst. Pink-cream laminated “cap” dolomite
Basal diamictite and “cap” dolomite (Areyonga Fm.); black shale, siltstone, feldspathic and lithic sandstone, stromatolitic dolomite. Up t o 1800 m
Unconformable on Bitter Spr. Fm. or equiv. Laterally discontinuous. Thick in SW. Lenticular grey silty “cap” dolomite
Quartz sandstone, siltstone, conglomerate; dolomite and limestone (stromatolitic), shale, sandstone, gypsum and halite. 750-1500 m
Unconformable on basement - Arunta Inlier, Musgrave Block. Thickest in west. Abundant stromatolites, microfossils
Basal diamictite (Mt. Davenport Dia. Mbr.), “cap” dolomite, red shale and sandstone. 0-800 m
Unconformable between L. Cambrian and Rinkabeena Sh. Trace fossils Yuendumu Sst. Pink laminated “cap” dolomite
Naburula FormationRinkabeena Shale
Basal diamictite, “cap” dolomite (Naburula Fm.); grey siltstone, green shale. Up t o 100 m
Unconformable on Albinia Fm. - Vaughan Spr. Qtzte. Dark grey “cap” dolomite
Vaughan Springs QuartziteAlbinia Formation
Quartz sandstone, siltstone, basal conglomerate; siltstone, shale, stromatolitic dolomite. 0-2500 m
Unconformable on basement-Arunta Inlier. Albinia Fm. disconformable on Vaughar. Spr. Qzte.
Arkose, siltstone, shale, (stromatolitic) dolomite, glauconitic sandstone. 300-1800 m
Unconformable between L. Cambrian and Keepera Gp. “Mt. Skinner” faunaCentral Mt. Stuart Beds. More extensive than units below
Red-brown arkose, siltstone, shale, dolomite. 0-1100 m
Unconformable on Yardida Till. Deposited only in fault troughs
AM ADEUS BASIN Olympic-Pertatataka-Julie Formations; Arumbera Sandstone I ; and equivalents
Areyonga-Aralka and equivalents
Formation,
Heavitree Quartzite-Bitter Springs Formation, and equivalents
NGALIA BASIN Mount Doreen FormationYuendumu Sandstone I
G EO KGINA B A S I N MOPUNGA GROUP
KEEPERA GROUP
285 TABLE 4. IX (continued) Unit
Rock types and thickness
Remarks
GEORGINA BASIN (continued) Yardida Tillite
Diamictite, arkose, siltRelationship t o Yackah stone, grey “cap” dolomite, Beds concealed. Deposited shale. 0-3000 m only in fault troughs
Yackah Beds
Feldspathic sandstone, shale, stromatolitic dolomite. 0 -250 m
Unconformable on basement - Arunta Inlier. Preserved only in fault troughs
NGALIA BASIN EAST Walbiri Dolomite
Todd River Dolomite 4rumbera Sandstone
GEORGINA BASIN SOUTH Red Heart Dolomite
?-
IU
Yuendumu Sandstonc
I r
irumbera Sandstone
Adam Shale
I
I I Arumbera Sandstone
1
Julie Formation
I
I - --- -- - - -
Cyclops Member Waldo Pedlar Member UpLer-mKker c3io%@e_ - - - Olympic Formation M t Doreen Formatioi
I
-_-_-_
____--
-----
-__-.
Keepero Group
Aralka Formation Yardida Tillite
Bitter Springs Formation Heavitree Quartzite
L+
Albinia Formation Quartzite
Vaughan Springs Quartzite
Yackah Beds
F52lA12154
+
1 Granitic and metamorphic basement
Vnconformrtv
Fig. 4.16. Correlation chart, Amadeus-Ngalia-Georgina Basins. (Adapted from Burek e t al., 1979, with revised nomenclature from Walter, 1980).
286
Tectonic setting The basin sequences were deposited on top of (Amadeus, Ngalia), or adjacent to (Georgina), the intracontinental Central Australian Mobile Belts (Figs. 4.1 and 4.2). This belt was the site of much thicker accumulations of Adelaidean and Palaeozoic sequences than on the North Australian and Gawler Cratons to the north and south (Plumb, 1979a,b). The Amadeus, Ngalia, and southern Georgina Basins are only the structurally preserved remnants of more extensive probably interconnected sedimentary covers, which existed prior t o the Petermann Ranges and Alice Springs-Mount Eclipse Orogenies. Thin remnants of Adelaidean rocks still extend around the western end of the Arunta Inlier, from the Amadeus Basin to The Granites--Tanami Inlier (Redcliff Pound Group; Blake et al., 1979) and, during the Adelaidean, the Amadeus Basin was probably connected with the Adelaide Geosyncline and Officer Basin to the south (Preiss et, al., 1978). Stratigraphy The Adelaidean successions of each of the basins may be divided into three tectosomes, each separated by unconformities (Fig. 4.16); other breaks are only identified locally. The tectosomes may be characterized from the Amadeus Basin. The Heavitree Quartzite-Bitter Springs Formation and equivalents comprises a widespread basal arenite passing up into a carbonate-rich succession, and is relatively uniform in thickness. This is overlain with regional unconformity by the Areyonga-Aralka Formations and equivalents: a sequence of basal diamictite overlain by arenites, lutites, and carbonates. The “cap dolomite” immediately overlying the basal diamictite is characteristically grey, silty, laminated, and lenticular. The units are laterally discontinuous. FormationsThe unconformably overlying Olympic-Pertatataka-Julie Arumbera Sandstone and equivalents again comprise a basal diamictite passing up into lutites, arenites, and carbonates but, in this case, the sucession is more laterally continuous. The “cap dolomite” is everywhere present on top of the younger diamictite, and is characteristically pink or cream and laminated. An erosion break exists between the Arumbera Sandstones I and I1 (Fig. 4.16), and their equivalents, passing southwestwards into an angular unconformity between the Adelaidean and Lower Cambrian - the Petermann Ranges Orogeny. The Arumbera Sandstone I and its equivalents contain trace fossils of the “Ediacara” and “Mount Skinner Faunas”, similar to those of South Australia (Wade, 1969; Glaessner, 1969; 1971). Correlations By the use of stromatolite zonation (Preiss, 1977), the diamictites and their characteristic “cap dolomites” (Preiss et al., 1978), and magneto-
287 stratigraphy (Burek et al.. 1979), correlations have been refined within and between the central Australian basins (Fig. 4.16), and between central Australia and the type-Adelaidean of South Australia: The Heavitree Quartzite-Bitter Springs Formation equals the Burra Group; the Areyonga-Aralka Formation equals the Umberatana Group (excluding Yerelina Subgroup) ; The Olympic Formation-Arumbera Sandstone equals the Y erelina Subgroup and Wilpena Group (see Rutland et al., this volume, Chapter 5). The terminology and correlations have been significantly revised within the Amadeus Basin by Preiss et al. (1978) and Wells (in press) (cf. Wells et al., 1970). The revised correlations are used here.
Age The maximum age for the base of the Amadeus Basin succession is provided by dyke swarms about 900 Ma old, in the underlying Arunta Inlier (Black et al., 1980). This age for the Amadeus Basin (<900 Ma) is consistent with the Upper Riphaean age determined by stromatolite zonation (Walter, 1972; Preiss, 1977), and with isotopic data from the Adelaide Geosyncline (Rutland et al., this volume, Chapter 5). Shales and coarser clastic sediments from almost the whole thickness of the Pertatataka Formation, and its equivalents, yielded a Rb-Sr isochron, of uncertain significance, of 725 f 45 Ma (Compston and Taylor, 1969); the ages of about 725 m.y. and 680 or 660 Ma, for the equivalent upper and lower tillites of the Kimberley region (this paper), are the best estimates for the age of the Areyonga and Olympic Formations (Walter, 1972). Palaeogeography The Adelaidean succession, as a whole, accumulated as a southwardthickening wedge, thickest at or just south of the present southern edge of the Amadeus Basin. A mirror-image, northwards-thickening wedge occurs in the Officer Basin, t o the south of the Musgrave Block (M.J. Jackson, pers. commun., 1979). A thinner shelf zone, with local depressions and ridges, extended across the Amnta Inlier from the northern Amadeus Basin to the Ngalia Basin. Much of the Adelaidean in the southern Georgina Basin accumulated in a series of narrow, deep, northwesterly trending grabens or aulacogenes, impressed into the southern edge of the North Australian Craton. Stage 1 . Following uplift and erosion of the basement rocks of the underlying Arunta Inlier and Musgrave Block, Adelaidean deposition began with a largely westward-thickening tabular body of fluvial and shallow-marine sand (Heavitree Quartzite and equivalents), passing up into stromatolitic carbonates and evaporites, deposited in a marginal-marine and lagoon complex (Bitter Springs Formation and equivalents).
288
Stage 2. The tectonism which followed radically changed the depositional framework. A west-trending hinge zone developed through the centre of the Amadeus Basin, and sedimentation of the Areyonga--Aralka Formations and equivalents to the north of this line was laterally discontinuous, controlled in the Ngalia and Georgina Basins, at least, by west to northwest-trending horsts and grabens. Basal diamictites of glacial origin were derived from the Arunta Inlier and from the underlying Bitter Springs Formation and equivalents. These pass up into sequences of paralic and shallow-marine lutites, arenites, and carbonates. To the south of the hinge line, the equivalent diamictites and paralic sediments were deposited as a southward-thickening wedge, immediately adjacent t o a provenance area in the Musgrave Block perhaps an uplifted block t o the south of the Woodroffe Thrust. Stage 3. Following further mild tectonism, relatively continuous deposition of the Olympic-Pertatataka-Julie Formations, Arumbera Sandstone, and equivalents, occurred across a shelf, t o the north of the hinge line of the central Amadeus Basin. At the base, diamictites of glacial origin in the northeastern Amadeus and Ngalia Basins (Olympic Formation, Mount Davenport Diamictite Member) pass elsewhere into arkosic sandstones and conglomerates; clasts are from underlying Adelaidean strata and from the Arunta Inlier. These units pass up into shallow-marine lutites, carbonates, and arenites. In the southern Amadeus Basin, another southward-thickening wedge of paralic sandstone, siltstone, and conglomerate, passing northwards into more shallow-marine deposits, was again derived from a geanticlinal welt on the Musgrave Block.
Structure The structure of the Amadeus Basin is dominated by large thrust-nappe complexes, incorporating both basement and cover rocks, at its margins. The Late Proterozoic or Early Cambrian Petermann Ranges Orogeny produced a large nappe complex, hundreds of kilometers long, at the southwestern margin which transported rock masses tens of kilometers to the northeast (Forman, 1966). The Carboniferous Alice Springs Orogeny produced several nappes along the northern and northeastern margin of the basin, with movements of several kilometers t o the south (Stewart, 1967; Forman, 1971; Shaw et al., 1971; Marjoribanks, 1975, 1976). These nappe-thrusts pass, in the centre of the basin, into Jura-type folds with axes hundreds of kilometers long, by virtue of a dkcollement along two surfaces - one within the Bitter Springs Formation, and one within the Early Cambrian Chandler Limestone (Wells et al., 1970). Intense retrogressive metamorphism occurred in the basement during these events. The Ngalia Basin is now preserved as a faulted asymmetrical syncline deformed along its northern margin, mostly during the Carboniferous Mount Eclipse Orogeny by a series of southward-directed thrust sheets and reverse
289 faults (Wells and Moss, in prep.). The southern margin of the Georgina Basin was deformed during the Carboniferous by northward displacement on shallowly southward-dipping planar thrusts (BMR, 1978). The reason for this intracontinental tectonism in central Australia, hundreds of millions of years after “cratonization”, remains elusive (Stewart et al., 1976). To explain the major gravity features of the region, Forman (in Wells et al., 1970) postulated deformation of both the crust and mantle. Forman and Shaw (1973) described a process of underthrusting of subAmadeus crust down deformed zones passing through the crust to the mantle, and regarded the area as an example of intraplate reaction. However, Anfiloff and Shaw (1973) showed that all the gravity features of central Australia may be quantitatively explained by heterogeneities in the upper crust. They also showed that the major gravity ridges terminate abruptly against major north-trending gravity lineaments and interpreted these as the sites of major crustal dislocations, between which n o r t h s o u t h compression was confined. Duff and Langworthy (1974) suggested that intraplate tectonics took place which included supracrustal folding and thrusting along ancient zones of weakness, in response t o the meridional compression postulated by Anfiloff and Shaw, and they attributed the compression t o rapid intermittent drift of the Gondwanide plate. K i m herley glacial successions (Principal references: Dow, 1965; Perry and Roberts, 1968; Dow and Gemuts, 1 9 6 9 ; Roberts et al., 1 9 7 2 ; Sweet et al., 1974a,c; Gellatly et al., 1975; Plumb and Gemuts, 1976; Sweet, 1977; Plumb, in press).
Mildly disturbed, late Adelaidean glacial successions are well exposed in three separate areas of the Kimberley and Victoria River regions (Fig. 4.17). It is not known whether the deposits were once continuous between these areas or not. The successions rest unconformably on a variety of rocks, and are unconformably overlain by the Early Cambrian Antrim Plateau Volcanics. The diamictites are noted for their well-preserved evidence of undoubted glacial features, and the abundance of glacial pavements. The stratigraphy is summarized in Table 4.X and Fig. 4.18. Stratigraphy Two successions of glacigene and interglacial rocks, referred to as the “Landrigan Glaciation” * and “Egan Glaciation” (Dow and Gemuts, 1969),
* Because of controversy about correlations within the Kimberley region, the “Moonlight Valley Glaciation” of Dow and Gemuts (1969) is here replaced by “Landrigan Glaciation”. Both the upper and lower glaciations are now named from the one section (Fig. 4 . 1 8 ) .
290
124°00 ~
-~
-- -
0
50
130°00'
I00
150
2 0 0 hm
Fig. 4.17. Distribution of "Glacial Successions", Kimberley region. (After Plumb and Gemuts, 1976; with permission of International Geological Congress).
291 MOUNT
RAMSAY A R E A
Metres
a 3 0 [L
a
3000
~
:
br
9" 686272
Do1 Err
.739?30 Do1 Do1
Fig. 4.18. Diagrammatic stratigraphic columns and correlations, "Glacial Successions", Kimberley region. (Adapted from Plumb and Gemuts, 1976). N.B.: Isotopic ages here are original data (1.39 X lo-" yr-' decay constant of 87Rb), and are c. 2%older than recalculated ages in text.
292 TABLE 4.X Summary of Late Adelaidean (latest Proterozoic) stratigraphy, Kimberley region Unit
M o u n t Harnsay area LOUISA DOWNS GROUP
KUNIANDI GROUP
East Kimberlr>yarea ALBERT EDWARD GROUP
DUERDIN GROUP
M o u n t House area MOUNT HOUSE GROUP
Rock types and thickness
Remarks
Basal diamictite, dolomite, arkose, “cap” dolomite (Egan F m ) ; dark shale, sandstone, subgrey Wac ke, dolomite, conglomerate. 3800 m +
Unconformable between L. Cambrian and Kuniandi Gp. Persistent yellow and pink laminated “cap” dolomite. Turbidites in upper part of Gp.
Basal diamictite and “cap” dolomite (Landrigan Till.); green-grey shale, siltstone, ferruginous sandstone. U p t o 1025 m
Unconformable o n L. Proterozoic. Impersistent purple and cream laminated silty “cap” dolomite. Limited inliers beneath Louisa Downs Gp.
Green-purple shale, siltstone; sandstone, dolomite. 1450-2000 m +
Uncomformable between
L. Cambrian and Duerdin
Basal diamictite, sandstone (Fargoo Till., Frank R . Sst.); diamictite and “cap” dolomite (Moonlight Valley Till.); green-red shale, siltstone, sandstone. 980-1450 m
GP. Moonlight V. Till. unconformable o n Frank R. Sst. discontinuous beneath, and unconformable o n basement. Persistent pink and cream laminated “cap” dolomite - Moonlight V. Till. Scattered erratics in and above “cap” dolomite
Basal diamictite, sandstone, shale, and “cap” dolomite (Walsh Till.); grey-green shale, siltstone, ferruginous sandstone, dolomite. 350-500 m +
Unconformable o n Kimberley Basin succession. Persistent pink to purple laminated “cap” dolomite. Scattered erratics above “cap” dolomite
are separated by an unconformity (Fig. 4.18), and may be correlated with reasonable confidence with the Sturtian and Marinoan sequences of South Australia (Dunn et al., 1971; Rutland et al., this volume, Chapter 5). The Mount House, Kuniandi, Louisa Downs, and Duerdin Groups each comprise a basal unit of diamictite overlain by thin persistent distinctive pink, cream, or purple laminated fine-grained dololutites - the “cap dolomites” - and then pass up into thick sequences of mostly green or grey siltstones, shales and sandstones (Table 4.X, Fig. 4.18).
293 The typical diamictites are unsorted, poorly stratified, and contain scattered polished and striated megaclasts, up to 5.5 m across, in a massive silt matrix which shows characteristic tillite fissility or structure. Locally, the matrix is partly sorted t o sand. The number of clasts decrease, and bedded shales appear, in the upper few metres of most diamictites. The diamictites grade laterally into sandy conglomerate and bouldery sandstone of fluvioglacial and periglacial or marine origin. The matrix of the Fargoo Tillite, and the associated Frank River Sandstone, is carbonate-rich, reflecting a carbonate-rich provenance seen in many of the megaclasts. The Egan Formation comprises diamictite, interbedded with a complex sequence of dolomite, limestone, arkose, and sandstone. Dropstone clasts are occasionally present in the “cap dolomites” or in the immediately overlying beds. The “Landrigan cap dolomite” is impersistent and silty, in contrast to the laterally persistent, pure dolomites above the other diamictites. Glaciated pavements and palaeotopography Twenty individual glacial pavements are now known from the Kimberley region, representing all the tillites except the Fargoo (Fig. 4.17); the largest single polished surface measures 55 m X 23 m. Polished surfaces, striae, grooves, crescentic gouges, crescentic fractures, and whaleback and stoss-andlee forms confirm a glacial origin. The predominant directions of ice movement, from the northeast or north, are consistent with the provenances of megaclasts in the tillites: widespread sources in the Kimberley Basin for the Walsh, Landrigan, and Egan Tillites, and widespread sources on the Sturt Block for the Moonlight Valley Tillite. The Fargoo Tillite has a distinctly different provenance; varied, locally derived megaclasts from the Halls Creek Mobile Zone, and unidentified carbonate rocks, indicate a more westerly source. Distribution of pavements and tillite can be related t o palaeotopography ; glacial valleys can be identified in places. The widespread provenance, and the thin, laterally persistent diamictites of the Walsh, Moonlight Valley, and Egan Tillites, suggest broad ice sheets. The Fargoo Tillite is impersistent, laterally variable, and clearly fills in an irregular palaeotopography.
Correlations The successions in the Kimberley region have been correlated, as indicated in Fig. 4.18, on the basis of similarities between the overall sequences of the Mount House, Duerdin, and Kuniandi Groups, the identical “cap dolomites”, and the unconformities below the Louisa Downs and Albert Edward Groups (Dow and Gemuts, 1969; Roberts et al., 1972; Plumb and Gemuts, 1976). Alternatively, Coats and Preiss (1979) consider that the “cap dolomite” of the Egan Formation is identical t o those of the Walsh and Moonlight Valley Tillites, which in turn resemble the distinctive Marinoan “cap dolomites”
294 throughout central and South Australia (Preiss et al., 1978), while the silty, impersistent Landrigan “cap dolomite’’ is more characteristic of the underlying Sturtian, and they correlate, in detail, the Duerdin plus Albert Edward Groups with the Wilpena Group (Marinoan) in South Australia (see Rutland et al., this volume, Chapter 5). This alternative implies correlation of the Landrigan Tillite with the Fargoo Tillite. Furthermore, J.C. Crowell and K.A. Plumb (unpublished data, 1976) have found that the Landrigan Tillite comprises only one thin diamictite; the “lenticular sand unit”, which many have correlated with the Frank River Sandstone (Fig. 4.18), is infolded Stein Formation in a structurally complex area. Plumb (in press) has also suggested alternative and simpler correlations of the various “members” within the Egan Formation, to those of Roberts et al. (1972). These alternative correlations cannot be finally resolved without further work, but Coats and Preiss’ alternative resolves many anomalies (isotopic ages; different source of Fargoo Tillite; etc.) and is perhaps t o be preferred (Plumb, in press). Age o f the sequences * Shale whole-rock Rb-Sr isochrons (Bofinger, 1967) provide internally consistent estimates of the age of the glacial successions (Fig. 4.18), regardless of the correlation chosen. A composite isochron from the Moonlight Valley/Landrigan Tillites is largely controlled by samples from the Landrigan Tillite, so the age of 723 30 Ma remains a reasonable estimate for the base of the “Landrigan Glaciation”. Identical isochrons of 670 85 Ma (Throssell Shale), 671 72 Ma (Johnny Cake Shale Member), and 671 f 44 Ma (composite; Throssell Shale/ Johnnv Cake Shale Member) confirm the correlation of these units. The Elvire Formation is 639 2 48 Ma old. A composite isochron from the McAlly/Timperley Shales is dominated by McAlly Shale samples, so 651 43 Ma becomes a reasonable estimate for this unit. These latter two ages provide age constraints on the “Egan Glaciation”, whilst that of the Throssell Shale/Johnny Cake Shale Member can be applied to either glaciation, depending on correlations : the “Egan Glaciation” commenced at about either 680 Ma or 660 Ma.
*
*
*
*
Palaeogeograp hy Perry and Roberts (1968), Roberts et al. (1972), and Sweet et al. ( 1 9 7 4 ~ ) interpret the diamictites as terrestrial tillites, while Dow and Gemuts (1969) consider that most of the East Kimberley diamictites were deposited in a marine environment beneath a floating ice sheet. Plumb and Gemuts (1976) and Plumb (in press) have concurred with both viewpoints, in different areas.
* These isotopic ages have recently been revised and reinterpreted by Plumb (in press), but the broad conclusions remain similar.
295 Most of the diamictites have been deposited directly onto basement by extensive grounded or floating ice sheets, at or near the edge of marine basins. Terrestrial influence predominates in the Walsh Tillite and in the northeastern areas of the Moonlight Valley Tillite. Marine deposition becomes increasingly important southwards in the Moonlight Valley Tillite, and in the Egan Formation. Marine transgression accompanied glacial retreat and produced periglacial deposits (bedded shales, carbonates) at the top of the tillite units, and scattered dropstones in the immediately overlying deposits. The interglacial deposits are shallow-marine. Terrestrial glaciers are suggested for the Landrigan Tillite. The thick massive diamictite of the Fargoo Tillite and the overlying fluvial Frank River Sandstone show marked lateral variations, were deposited onto a highly irregular palaeotopography, and contain abundant outwash material. Perhaps a mountain glaciation, from the uplifted Halls Creek Mobile Zone, is indicated. Conclusion Northern Australia is dominated by the North Australian Craton. The craton is bounded on the north and south by younger mobile belts - the Northeast Orogens and the Central Australian Mobile Belts. Most of the region is now overlain by mildly deformed Proterozoic and Phanerozoic platform covers. The tectonic evolution of northern Australia may be directly related to the nature of its basements, and t o the times of their cratonisation. A uniform pattern of major fractures developed throughout the region very early in its history, and this has controlled the subsequent tectonic evolution of the region by the repeated reactivation of these fractures. These ancient basement mobile belts and fractures have repeatedly controlled the distribution of the thickest sequences and the most intense deformation of the later platform covers. Early Proterozoic orogenic domains of the North Australian Orogenic Province were initially distributed right across the region. Most were cratonized by the end of the Early Proterozoic and were then overlain by Early t o MidProterozoic covers of the North Australian Platform Cover, t o form the North Australian Craton. Along the borders of the craton the Early Proterozoic rocks were incorporated into and reworked by Mid-Proterozoic mobile belts -the Central Australian Mobile Belts. These were, in turn, overlain by the Late Proterozoic-Palaeozoic Central Australian Platform Cover. Elements of the 1950-1900 Ma McClintock Subprovince underlie much of the region, but only t o the west of the Halls Creek Fault was the Halls Creek Inlier cratonized during the interval 1920-1815 Ma, and then overlain by 1815-1760 Ma platform cover - the Kimberley Basin. East of the Halls Creek Fault, the younger Palmerston Subprovince developed on and between McClintock Subprovince rocks, was cratonized during the interval 1840-
1680 Ma, and then overlain by various basins of the 1700-1300 Ma North Australian Platform Cover. Farther east, elements of the McClintock and Palmerston Subprovinces were incorporated into the Mount Isa Orogen, overlain by younger sequences, then all were cratonized about 14901460 Ma ago and overlain by platform cover of the South Nicholson Basin. This progressive cratonization and development of platform covers from west to east, in northern Australia, is identical t o the progressive west t o east cratonization of the Phanerozoic Tasman Orogenic Province and the development of the Trans-Australian Platform Cover above it (GSA, 1971). This may reflect the presence of a continental margin in the east, although all the domains of the North Australian Orogenic Province were ensialic. The Central Australian Mobile Belts are typical intracontinental ensialic polymetamorphic belts, in which Early Proterozoic rocks apparently continuous with those of the North Australian Orogenic Province t o the north, and the Gawler Craton t o the south, were repeatedly metamorphosed, deformed, and intruded by granite along a zone of high heat flow throughout the Mid-Proterozoic. A probable combination of incipient rifting, thrusting or underthrusting, and uplift along deep-seated structures, which probably extend down t o the mantle, have brought slabs from a wide range of crustal levels into contact at the surface. This belt remained a zone of relative instability throughout the development of the Late ProterozoicPalaeozoic Central Australian Platform Cover, culminating in the formation of unusual intracontinental supracrustal thrust-nappe complexes during the Late Proterozoic and again during the mid-Palaeozoic. The equivalent Central Australian Platform Cover which overlies the North Australian Craton is thin and mildly deformed. Although major vertical displacements are apparent, the structural evolution of much of the region can be related t o horizontal displacements along major faults. Much of the structure can be fitted into a simple model of uniform movements in which a central block, roughly between the Kimberleys and Mount Isa, was displaced northwards relative t o the blocks on either side and, locally, the horizontal displacements were absorbed along east-west-trending zones of thrusting and folding, where the cover was crumpled against rigid blocks during southward transport. In this model, the Central Australian Mobile Belts would represent a zone of interaction between the North Australian and Gawler Cratons, during northward transport of both of them. The Early Proterozoic rocks of northern Australia are a major source of U and 17-Au deposits, while the Mid-Proterozoic rocks of the McArthur Basin-Mount Isa Orogen comprise a major province of stratiform Pb-Zn deposits and, to a lesser extent, Cu.
.
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306 Sweet, I.P., 1977. The Precambrian geology of the Victoria River region, Northern Territory. Bur. Miner. Resour. Aust. Bull., 168. Sweet, I.P., Mendum, J.R., Bultitude, R.J. and Morgan, C.M., 1974a. The geology of the southern Victoria River region, Northern Territory. Bur. Miner. Resour. Aust. Rep., 167. Sweet, I.P., Mendum, J.R., Morgan, C.M. and Pontifex, I.R., 1974b. The geology of the northern Victoria River region, Northern Territory. Bur. Miner. Resour. Aust. Rep., 166. Sweet, I.P., Pontifex, I.R. and Morgan, C.M., 1974c. The geology of the Auvergne 1 : 250,000 Sheet area, Northern Territory. Bur. Miner. Resour. Aust. Rep., 161. Thomas, W.A., 1977. Evolution of Appalachian-Ouachita salients and recesses from reentrants and promontories in the continental margin. Am. J. Sci., 277: 1233-1278. Thomson, B.P., 1966. The lower boundary of the Adelaide System and older basement relationships in South Australia. J. Geol. SOC.Aust., 13: 203-228. Traves, D.M., 1955. The geology of the Ord-Victoria region, northern Australia. Bur. Miner. Resour. Aust. Bull., 27. Wade, M., 1969. Medusae from uppermost Precambrian or Cambrian sandstones, central Australia. Palaeontology , 1 2 : 3 51-365. Walker, R.N., Muir, M.D., Diver, W.L., Williams, N. and Wilkins, N., 1977. Evidence of major sulphate evaporite deposits in the Proterozoic McArthur Group, Northern Territory, Australia. Nature, 265: 526-529. Walker, R.N., Logan, R.G. and Binnekamp, J.G., 1978. Recent geological advances concerning the H.Y.C. and associated deposits, McArthur River, N.T. J. Geol. Soc. Aust., 24: 365-380. Walpole, B.P., Crohn, P.W., Dunn, P.R.and Randal, M.A., 1968. Geology of the KatherineDarwin region, Northern Territory. Bur. Miner. Resour. Aust. Bull. 82. Walter, M.R., 1972. Stromatolites and the biostratigraphy of the Australian Precambrian and Cambrian. Palaeontol. SOC.Lond., Spec. Pap. in Palaeontol. 11. Walter, M.R., 1980. Adelaidean and Early Cambrian stratigraphy of the southwestern Georgina Basin: correlation chart and accompanying notes. Bur. Miner. Resour. Aust. Rep., 214. Walter. M.R., Krylov, I. and Muir, M.D., in prep. Carpentarian (Middle Proterozoic) stromatolites from the McArthur Basin and Mount Isa region, Australia. Alcheringa. Webb, A.W., 1973. Geochronology of the McArthur Basin. AMDL Rep. AN 1814/73 (unpubl.). Webb, A.W. and Page, R.W., 1977. Geochronology of glauconitic sandstone and shale beds in the Victoria River region, Northern Territory. In: I.P. Sweet, The Precambrian Geology of the Victoria River Region, Northern Territory. Bur. Miner. Resour. Aust. Bull., 168: 68-72. Wellman, P., 1976. Gravity trends and the growth of Australia: A tentative correlation. J. Geol. SOC.Aust., 23: 11-14. Wellman, P., 1978. Gravity evidence for abrupt changes in mean crustal density at the junction of Australian crustal blocks. BMR J. Aust. Geol. Geophys., 3 : 153-162. Wells, A.T., in press. Late Proterozoic diamictites of the Amadeus and Ngalia Basins, central Australia. In: W.B. Harland and M.J. Hambrey (Editors), Pre-Pleistocene Tillites: A record of Earth's Glacial History. Cambridge University Press, Cambridge. Wells, A.T. and Moss, F.J., in prep. The Ngalia Basin, Northern Territory, stratigraphy and structure. Bur. Miner. Resour. Aust. Bull. Wells, A.T., Forman, D.J., Ranford, L.C. and Cook, P.J., 1970. Geology of the Amadeus Basin, central Australia. Bur. Miner. Resour. Aust. Bull. 100. Whitcher, I.G., 1975. Dugald River zinc-lead lode. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall., Monogr., 5: 372-376.
307 Williams, N., 1978a. Studies of the base metal sulphide deposits at McArthur River, Northern Territory, Australia I. The Cooley and Ridge Deposits. Econ. Geol., 73: 1005-103 5. Williams, N., 1978b. Studies of the base metal sulphide deposits at McArthur River, Northern Territory, Australia 11. The sulphide-S and organic-C relationships of the concordant deposits and their significance. Econ. Geol., 73: 1036-1056. Williams, N., 1979. The timing and mechanisms of formation of the Proterozoic stratiform Pb-Zn and related Mississippi Valley-type deposits at McArthur River, N.T., Austxalia. Soc. Econ. Geo1.-Am. Inst. Min. Eng. Joint Meeting, New Orleans, Preprint NO. 79-51. Williams, N., 1980. Precambrian mineralisation in the McArthur-Cloncurry region, with special reference t o stratiform lead-zinc deposits. In: R.A. Henderson and P.J. Stephenson (Editors), The Geology and Geophysics of Northeastern Australia. Geol. SOC. Aust., Qld Div., Brisbane, pp. 89-107. Wilson, I.H., 1978. Volcanism on a Proterozoic continental margin in northwestern Queensland. Precambrian Res., 7: 205-235. Wilson, I.H. and Derrick, G.M., 1976. Precambrian geology of the Mount Isa region, northwest Queensland. 25th Int. Geol. Congr., Sydney, Excursion Guide 5A, 5C. Wilson, I.H., Derrick, G.M. and Hill, R.M., 1972. Copper mineralization, (including Mount Isa) in the Precambrian Cloncurry Complex of northwest Queensland, Australia. 24th Int. Geol. Cong., Montreal, Sect. 4,234-240. Windley, B.F., 1973. Crustal development in the Precambrian: Philos. Trans. R. SOC. London, Ser. A, 273: 321-341. Withnall, W., Bain, J.H.C. and Rubenach, M.J., 1980. The Precambrian geology of northeast Queensland. In: R.A. Henderson and P.J. Stephenson (Editors), The Geology and Geophysics of Northeastern Australia. Geol. SOC. Aust., Qld Div., Brisbane, pp. 109127.
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309
Chapter 5
THE PRECAMBRIAN OF SOUTH AUSTRALIA R.W.R. RUTLAND, A.J. PARKER, G.M. PITT, W.V. PREISS and B. MURRELL
INTRODUCTION
In South Australia the Gawler Province displays a history of Early Proterozoic orogenic evolution, and transitional tectonism leading to deposition of the Gawler Range Volcanics (c. 1500 Ma) on a relatively stable platform. A fragmentary record of platform deposition (the Stuart Shelf) on the resulting Gawler Craton provides the link with the Adelaide “Geosyccline” (mainly
*
In trod uct ion The principal outcrop of the Gawler Province is in the Gawler orogenic domain (Fig. 5.1). However, although the area of basement rock as shown in
____
* By R.W.R.
Rutland, A.J. Parker and G.M. Pitt.
310
___
Local Archaean relics - - - - _ _ 79 44
0
KILOMETRES SA
Department of Mines and Energ
Fig. 5.1. Main subdivisions of t h e Precambrian of S o u t h Australia. WT = Woodroffe Thrust ; THZ = Torrens Hinge Zone; T = Tarcoola region.
Fig. 5.1 is extensive, it is generally of very low relief and commonly blanketed by a thin superficial cover. Thus the geology is relatively well known only in the area of higher relief between the Gawler Range Volcanics and the NW shore of Spencer Gulf (Fig. 5.2). There, a series of Early Proterozoic metasedimentary rocks (the Hutchison Group) and granitic gneisses (the Lincoln Complex) have been highly metamorphosed and tightly interfolded t o form a major north- to northeast-trending orogenic belt. Recent work has shown that Archaean gneisses, forming the basement t o the Early Proterozoic metasedimentary sequences, are widespread in the less well-exposed areas in the W and NW of the domain. Some of the granitic gneisses of the
311
100
0
I
K l LOM ETR ES
Syntectonic granites - - - - - -
Adelaidean and Palaeozoic sediments-UNCONFORMITY
CB
1100 my
Post-tectonic granites - _ - _ _ _ - _
--
I+
Lincoln Complex---
4 1
--_ _ - --_-
_______________
Sleaford Complex-
*.
1
-
BIF Hutchison Group
UNCONFORMITY co 2100 my
to._ '7"
_ _ -_
-v-v-v-
v
___ ____---__ S
A
Department of Mines and Energy
Fig. 5.2. General geology of the Gawler orogenic domain in the Eyre Peninsula. Locatior shown o n Fig. 5.1.
4 8 Low to medium grade metamorphics.
2 Calc-silicate
and albitic formations-
Mutooroo area Medium to high grade metamorphics__.
0M
type iron formations
________
fg.{.,
Q. Arbitrary datum based on iron formation horizon _ _ _ .
Fault inferred
2 d
3
,A.
Granitic gneiss ________._..__._ 70
71
Aeromagnetic lineament__.,;’
/
,*,
S A D e ~ a r t m e n ot f M nes and Enerqy
Fig. 5 . 3 . General geology of t h e Olary sub-domain in the Willyama orogenic domain. Location shown o n 5 . 1 .
Fig. 5.4. General geology of t h e Broken Hill sub-domain in t h e Willyama orogenic domain, showing inferred axial traces of superposed folds of three generations. Redrawn from map prepared b y N. Archibald as part of t h e Broken Hill Lode Horizon research programme at t h e Department of Earth Sciences, Monash University. Unpublished geological information from t h e N.S.W. Geological Survey is included. Location shown o n Fig. 5.1.
314 Lincoln Complex may represent reworked Archaean gneisses. In the southern part of the Musgrave domain, overprinting by later Musgravian events is weak and therefore this part is also included within t h e Gawler Province. Again, however, exposure is very poor and little attention has yet been paid t o the Early Proterozoic history. Early Proterozoic rocks of t h e Gawler Province also reappear in a major inlier, t h e Willyama orogenic domain, east of the main part of the Delamerian fold belt. This domain is divided by an area of relatively poor outcrop into t h e Olary subdomain in South Australia (Fig. 5 . 3 ) and the Broken Hill subdomain in New South Wales (Fig. 5.4). However, t h e tectonic development o f both these subdomains is so similar t o that of the Gawler domain that it is convenient t o consider them together. Smaller inliers of t h e Gawler Province are also found within t h e Delamerian fold belt both near Adelaide (Talbot, 1963; Mills, 1973) and, more extensively, in the M t . Painter domain (Coats and Blissett, 1971) and in the Peake and Denison Ranges (Ambrose and Flint, 1979) (Fig. 5.1). il rchaean basement rocks Several occurrences of Archaean basement rocks have recently been described from the Gawler domain (Cooper et al., 1976; Webb and Thomson, 1977; Daly et al., 1978). The gneissic rocks of the Eyre Peninsula were originally allocated t o the Flinders “Series” by Tilley (1921) who regarded these rocks as intrusive into the metasedimentary Hutchison Group. Johns (1961 ), however, suggested that the Flinders “Group” of his revised terminology represented a higher grade of metamorphism, with local granitization, of the metasediments, whereas Whitten (1966) suggested that the Flinders “Gneiss” of his revised terminology represented basement t o t h e overlying metasediments. This latter view has been supported by Coin (1976) and by Glen et al. (1977) although these authors noted that, because of the strong tectonism, no major stratigraphic unconformity had been demonstrated between t h e t w o groups. Granitic gneisses in the Broken Hill area are interlayered with t h e metasediments and have taken part in all the deformation episodes that affect the metasediments. These gneisses have been dated a t about 1660 Ma (Pidgeon, 1967; Shaw, 1968) which is interpreted as the age of t h e Willyama metamorphism. The relatively low initial ratios, however, suggested a maximum primary age for the rocks of little more than 1800 Ma, although t h e uncertainties involved in making such inferences were emphasized by Pidgeon. Cooper et al. (1976) dated some of the granitic gneisses in the Eyre Peninsula a t 1816 ? 1 0 Ma with an initial ratio of 0.7043 t 0.0008. They regarded the rock as a granite intrusion and the date as the date of intrusion. They suggested that if t h e date represents a metamorphism, only about 1 0 Ma could have elapsed between intrusion and metamorphism. It
315 must be presumed, therefore, that in both areas the dated events provide a minimum age for the associated metasediments. General similarity with the succession in the Hamersley Basin (Whitten, 1966) has led t o the view that the main metasedimentary sequences are of Early Proterozoic age. However, Cooper et al. (1976) also showed that some granulite facies metasediments and associated augen gneisses at Cape Carnot on the southc’rn end o € Eyre Peninsula are of Archaean age. They regarded a date of 2587 ? 130 Ma with an initial ratio of 0.7015 t 0.0043 as the most meaningful interpretation of their data for the augen gneisses (regarded as anatectic) but they were unable t o evaluate any effect of the subsequent Proterozoic metamorphism. Webb and Thomson (1977) obtained further evidence from granites in the south and similar rocks 400 km to the north in the Lake Harris-Tarcoola region. Their isochrons suggest dates between 2400 and 2300 Ma and provide a minimum age for the paragneisses which the granites rntrutle. Cooper et al. (1976) note that the Archaean rocks at Cape Carnot include sillimanite-xorderite gneiss with up t o 20% sillimanite. Basic granulites are interpreted as original dolerite intrusions. Daly et al. (1978) showed that a large area of paragneisses in the Tarcoola region are also sillimanite and cordierite bearing. They contain an iron formation which is a well layered quartz-magnetite-diopside-hypesthene-amphibole gneiss. Thin calcite layers containing minor mafic minerals are also present. The date of 2417 k 59 Ma from the sequence is inferred t o be the metamorphic age and the low initial ratio of 0.7036 t 0.0015 is interpreted t o indicate that the material had not resided in the crust for more than 100-150 Ma. Thus these late Archaean metasediments appear t o represent a relatively mature sedimentary sequence, unlike the main greenstone belts of the Eastern Goldfields sub-province of the Yilgarn orogenic domain, in either sediimentary character or their high grade of metamorphism (see Chapter 2). They are more similar in these characteristics to the older gneisses of the Wheat Belt sub-province in Western Australia. However, southeast of Tarcoola, Forbes et al. (1977) record an occurrence of relict pillow structures in basic hornfels intruded by the Glenloth Granite (2318 k 55 Ma). The iron formations are similar to, and were previously correlated with, those of the Middleback Ranges in the Hutchison Group (Whitten, 1966). However, Daly et al. (1978) report that initial isotopic ratios for units of the Hutchison Group suggest that their deposition could not have occurred prior t o c. 2000 Ma. Also the older iron formations are feldspathic whereas Hutchison Group iron formations are not and trace element compositions are distinct. It therefore seems clear that part of Tilley’s Flinders “Series” is indeed Archaean, at least in the sense that it forms an older basement, metamorphosed and intruded by granite, before the deposition of the Hutchison Group. That part is now assigned to the Sleaford Complex Thomson, 1980)
316 while the remaining c. 1800 Ma granite gneisses - those derived from pretectonic granite intrusions or from reworked Archaean basement or from migmatizetl Early Proterozoic metasediments (such that their original identity is no longer distinguishable) - are assigned to the Lincoln Complex. I t appears that the Hutchison Group has its principal development in the eastern part of the Eyre Peninsula but that westwards it has been largely removed by subsequent erosion to expose its reworked Archaean basement. It can be reasonably inferred that the quartzites at the base of the Hutchison Group were deposited on a major unconformity (Parker, 1978) although evidence of a structural contrast across the unconformity has not yet been presented. It is possible that much of the structure in the Archaean rocks is the consequence of the Proterozoic deformation episodes which also affected the overlying Hutchison Group. In any case, it is clear that the Archaean rocks do not represent a stable Archaean craton like the Pilbara and Yilgarn orogenic domains in Western Australia. Instead they represent part of the reworked basement t o the Early Proterozoic sequences. The absence of Archaean basement rocks in the Broken Hill sub-domain has already been noted. In the Olary sub-domain, granitic rocks consist of granitic. t o adamellitic intrusives of 1500-1550 Ma age, and semiconcordant granodioritic gneisses which, like the granite gneisses of Broken Hill, have the same metamorphic history as their enclosing metasediments. Stratigraphically interlayered with, and underlying the granodioritic gneiss bodies is a quartzofeldspathic gneiss sequence probably comparable t o the lower parts of the Hutchison Group of Eyre Peninsula. These appear to be part of the Willyama sequence and there is no evidence t o suggest they are Archaean. In fact their correlation with the lower gneisses of Broken Hill, and their low initial strontium ratios support their interpretation as part of the Willyama Complex. Thus the Archaean rocks that must have been basement t o the Willyama sedimentation are not present or not recognized as such, due t o Early to Middle Proterozoic metamorphism and anatexis (cf. Lincoln Complex). The Early Proterozoic sequences Early Proterozoic sedimentation throughout the Gawler Province is notable for its relatively shallow water, mature and uniform c h a i & m -over wide areas (Glen et al., 1977). Details of the sedimentation and stratigraphy are most readily established for the higher parts of the sequence where the metamorphic grade is generally low and the deformation less severe (e.g., Tuckwell, 1978). In the lower part of the sequence highly complex superposed deformations and high grade of metamorphism make it difficult to correlate locally established successions over wide areas. In the Gawler domain, quartzites (the Warrow Quartzite) dominate the base of the succession, immediately above the Archaean basement gneisses. Sedimentary facing criteria have not been firmly established but a base has been inferred from structural evidence (Parker, 1978$, from the pr.Psence
317 of local quartz “pebble” bands near the supposed base of the quartzite (M. Fanning, pers. commun., 1978), and from a general similarity of the total succession with other Early Proterozoic sequences throughout the world. The quartzites are locally calc-silicate-rich at the base whereas higher up they are intercalated with pelitic schists. Immediately overlying the quartzites is a sequence of mixed chemical and semipelitic metasediments ranging from massive dolomite and associated banded iron formation and chert t o quartzo-feldspathic gneisses and schists. There are numerous conformable amphibolite sills throughout this part of the sequence but their origin is obscure. Their massive, uniform lithology and their chemical affinity to quartz tholeiites suggest an igneous, possibly intrusive, origin although amphibolitic dykes have only been recorded in the Middleback Ranges, west of Whyalla (Furber and Cook, 1976), where they may be related t o younger tectonic features. The chemical/semipelitic part of the sequence is well developed in the Middleback Ranges where it is locally known as the “Middleback Group” and where iron ore has been mined for over half a century. The underlying leucogneiss unit in that area is locally known as the Gneiss Complex, but is considered t o be a granitized equivalent of the basal (Warrow) quartzites that are so well developed t o the west and southwest. The composition of the iron formations is dominantly quartz hematite at the surface but below the weathered zone ranges from quartz-magnetite t o quartz-magnetite-amphibole (grunerite-actinolite) t talc, calcite, sulphide and diopside. Feldspar is absent. The weathering of the ore body was, a t least, partly Precambrian since clasts of the ore are found in the Corunna Conglomerate (see below). The rnetasedimentary sequence above the iron formation horizons has mostly been eroded away, but where part of it is locally preserved, it is of pelitic t o semipelitic character and generally of slightly lower metamorphic grade. Deformation and igneous intrusion have greatly modified the apparent thickness of the succession in the Gawler domain, but it is estimated t o be of the order of 1500 to 2000 m thick. In the Broken Hill mine area it has been shown that the lower part of the succession was inverted during the first deformation so that the major (F,) fold structures are downward facing (Laing et al., 1978). The “Mine Sequence” which was regarded as a stratigraphic succession facing northwest is now interpreted as containing both limbs of a downward facing antiformal structure separated by a syn-metamorphic slide (Fig. 5.6). Within the Mine area therefore the principal stratigraphic units are as follows (Fig. 5.5): 5 (Top). Pelitic and psammopelitic sillimanite gneiss (without metavolcanics). 4. Psammopelitic sillimanite gneiss (c. 1000 m thick) containing units of Mine sequence - Banded Iron Formation, amphibolite, Potosi gneiss, Lode horizon (with lead lodes stratigraphically above zinc lodes).
31.8
SUITE
5
.
.
.
. .~. . . .
-.
.
.
.
.
.
-
Psammopelite 8 pelite Potosi type gneiss Lode horlzon 8 ore lenses Potosi
type gneiss
Non-bedded
Sc J I 1-E
pelite, K . ( F e ) rich
Pscirnrnopslite 8 psammite
4
AmDhibol I te
Pelite
B
psarnmopelite
. . . . . . . . .
SUITE 3
. . . .
. . . .
. . . .
. . . .
. . . .
. . . .
. . . .
. . . .
. . . .
Quartrofeldspathic gnelss 8 arnphibolite Composite gneiss B arnphibolite
I ~ ~ q5 . 5. Stratigraphic sequence in the Broken Hill mines area. The bar scale is very :ipproximately 100 m . From W.P.Laing in unpublished report GS 1979/062, Geological Survey 6)f New South Wales.
3 (Rase). Granite gneiss (pre-F, granite or metavolcanic). T h e “lode horizon” and its associated “Potosi” and “Parnell” gneisses can hc. traced throughout the Broken Hill sub-domain. The “Potosi” and “Parnell” gneisses apparently represent metavolcanic rocks (especially metadacites) closely associated with Banded Iron Formation and with the Broken Hill mineralization, while the stratigraphically overlying sillimanite gneisses are generally regarded as mature non-volcanogenic metasediments. In contrast, Stanton (1976) interprets the whole environment as of volcanic island arc character (see also Rutland et al.,1978). Some feldspathic metasediments and associated gneisses east of Broken Hill (the Thorndale and underlying gneisses, possibly several thousand metres thick), appear t o lie stratigraphically below the mine sequence, while t o the northwest, major synclinal structures carry lower-grade metasediments
319 /-,
/
\
\
m
I
SE
w -
1
NW
Fig. 5.6. Schematic geological cross-section of the Broken Hill mines area. Numbers refer to metamorphic rock suites of Fig. 5.4. Ornamentation as o n Fig. 5.5. HWS = Hanging Wall Synform; B H A = Broken Hill Antiform; B H S = Broken Hill Synform. The bar scale is very approximately 1 km. From W.P. Laing in unpublished report GS 1979/062, Geological Survey of New South Wales.
stratigraphically above the mine sequence. Tuckwell (1978) has divided these younger metasediments into the Wookookaroo Beds (up t o approximately 4000 m thick) below, and the Bijerkerno Beds (up to 1500 m thick) above, separated by a prominent calc-silicate unit. The Wookookaroo beds commonly show graded bedding in peliti-psammite beds and were probably deposited by traction and turbidity currents. Psammitic beds are commonly feldspathic or calcareous. By contrast the Bijerkerno beds are well-laminated and consist of suspension deposited fine-grained muds and quartz-rich sediments. The sediments are often carbonaceous and the psammites are relatively pure quartzites.
320 No major stratigraphic breaks have been observed anywhere in the Broken Hill succession. Unit 5 above the mine sequence probably corresponds t o the lower part of the Wookookaroo Beds. It is notable, however, that the granitic gneisses and amphibolites are confined t o the high-grade part of the sequence although structural studies suggest that the whole succession has been involved in the same deformation episodes. Tuckwell (1978) has suggested that the mafic intrusive activity may have preceded deposition of the higher part of the succession and that a major tectonic or stratigraphic break may be present. Three distinct areas within the Olary sub-domain (Fig. 5.3) are recognized on metamorphic, lithological and deformational grounds (Forbes and Pitt, in prep.). Regional mapping has shown that specific stratigraphic units can be traced through both the Old Boolcoomata and Kalabity areas. The Old Boolcoomata arcba is, as a rule, more mobilized and gneissic than the less penetratively deformed Kalabity area. The recognition of a sequence, in the former area, depends o n the validity of interpreted isoclinal F2 folds which often lack exposed closures. Grades of deformation and metamorphism can vary from limb t.o limb resulting in schistose, gneissic or migmatoid variants of any particular unit or sequence of units. The consistency of local and regional sequences, erected on this basis, however, supports the assumed structure. The interpreted sequence of the Old Boolcoomata area is therefore broadly as follows: 4 (Top). Interlayered schist and quartzite. 3 . Knotted aluminosilicate and/or muscovitic schist, with ubiquitous tourmaline and garnet. Fibrolite often appears t o be later than both kyanite and andalusite. The andalusite is often retrogressed t o corundum plus pyrite. Chiastolite is developed locally in an upper, carbonaceous member. Calcsilicate and calc-albitic marker beds are present at the t o p and base - the “Ethiudna C!alc-silicate Group” of Campana and King (1958). 2. Layered, often magnetitic and/or feldspathic, gneisses, containing one to three thin, but laterally persistent, iron formations which display a 2 sulphide + magnetite hematite k baryte k quartz 2 albite mineralogy. 1 (Base). A thick, lower sequence of quartzo-feldspathic gneisses, which may be migmatoid or anatectic, containing lenticular semiconcordant granodiorite gneiss bodies near the top. There are no indications of an Archaean basement t o t h e succession and preliminary estimates give a thickness of some 4000-6000 m. Nevertheless this succession appears t o correspond very closely t o that described from Broken Hill. The Kalabity area is stratigraphically equivalent, but less penetratively deformed and kyanite has not been observed. However, the local development of fibrolitic sillimanite indicates that the metamorphic grade is not necessarily very much lower. _+
321 The structure of the Kalabity area is that of a broad north-closing anticlinal arch. Relative t o the Old Boolcoomata succession, Unit 4 is similar, while Unit 3 is a very thick, more homogeneous, carbonaceous metapelite up to 1.5 km thick with ubiquitous (now retrogressed) chiastolite and andalusite porphyroblasts. Iron formations equivalent t o those of Unit 2 at Old Boolcoomata are thicker and more massive and are directly associated with calc-silicate, calc-albitic and albitic horizons at the base of Unit 3 . Thus the correlation of Units 2 and 3 is clear but there appear t o be significant differences in thickness and facies between the two areas. Possibly the Kalahity area represents the more distal facies. Unit 1 is represented by psammopelitic schists and gneisses which occupy the core of the anticline and are intruded by late-tectonic granitoids, while granodiorite gneiss bodies like those of the Boolcoomata area are absent. The origin of albite-rich units in both areas is controversial. Acid volcanics are believed to have made some contribution, in the form of tuff or of clastics reworked from the original extrusives (cf. Vernon, 1961Stanton, 1976). Evaporites and related sediments are possibly also important t o the origin of both calc-silicate and albitic rock-types, but this is not proven. The Mutooroo area is of notably higher grade (upper amphibolite) than the other two areas described above, consisting of undifferentiated coarsergrained sillimanite-kyanite-garnet schists and gneisses. Local occurrences of “Aplite”, “augen gneiss” and “Q.M.” (quartz-magnetite iron formation) lithologies like those of Broken Hill have been recorded, but have not allowed a stratigraphic succession t o be established. However, it seems likely that, these rocks are equivalent t o the lower-grade successions to the north. A notable feature, not present in the other areas, is the presence of conformable amphibolites, regarded as of igneous origin. The summaries of recently acquired stratigraphic evidence give support to the general conclusions of Glen et al. (1977) regarding the stratotectonic unity of the Gawler Province. It appears that the whole province was characterized by relatively shallow water sedimentation on an older continental crust. An increase in thickness appears t o accompany a change from more proximal to more distal environments from west t o east. The lower part of the sequence with its banded iron formations and amphiholite sills is similar throughout the province and invites analogy with the Hamersley Group in Western Australia (Whitten, 1966). The available isotopic data place the sequence between 2400 Ma and 1800 Ma approximately, so that it could indeed be approximately the same age as the Hamersley Group. Unlike the Hamersley Group, however, it has suffered strong deformation and metamorphism. The probable metavolcanics (e.g., Potosi Gneiss) in the lower part of the sequence are well developed only in the Broken Hill sub-domain and appear to be genetically associated with the major stratiform base-metal sulphide
322 mineralization of that region (see e.g., Both and Rutland, 1976; Stanton, 1976; Laing et al., 1978).
Deformation and metamorphism Deformation and metamorphism is also similar in character throughout the Gawler Province (Glen et al., 1977). The earliest prominent schistosity in the high-grade rocks is a layer-parallel schistosity (S,) which in the Broken Hill region is known to be associated with stratigraphic inversion of the high grade rocks in the Broken Hill Mine area. In the low-grade rocks S1 is an axial plane schistosity and there is evidence of metamorphic crystallization prior to its development (Glen, 1978a). There is less strong evidence for porphyroblast development in the high-grade rocks prior t o S1 (Parker, 1978) and correlation of S, between the high- and low-grade rocks is not firmly established (e.g., Glen, 1978a). The second deformation was also accompanied by high-grade metamorphism and produced tight t o isoclinal folds with local transposition in the highgrade rocks. F2 folds are the major structures of the Broken Hill Mine area for example and they played an important part in localizing the Pb-Zn orebody (Laing et al., 1978). The ages of the F, and F, deformations in the Gawler Province are not easily separated. The isotopic dates reflect an age of metamorphism associated with one or both of these deformations. It can be concluded only that F, is not younger, and F, is not older, than these dates. The two deformations are bracketed in the Gawler domain between about 1816(?10) Ma given by Cooper et al. (1976) for the Kirton Point gneiss (which probably antedates both F, and F,) and about 1650 f 35 Ma for the syn F, Middle Camp Granite on east-central Eyre Peninsula (N.B. the post-F, Carpa Granite in the same area gives an age of 1677 k 125 Ma). The third deformation appears to be significantly younger throughout the province and is associated with the intrusion of late tectonic granites. These are dated at c. 1520 in the Broken Hill domain. In the Gawler domain regional F3 folds deform Carpa Granite equivalents but are unconformably overlain by the basal Gawler Range Volcanics dated at 1525 f 1 5 Ma (Blissett, 1977; Branch, 1978). The F3 deformation has produced major north-northwest- t o northeasttrending upright folds throughout the province and is generally accompanied by retrogressive metamorphism and the development of ductile shear zones. For example, in the Gawler domain a major mylonite zone was developed (Fig. 5.2) and in the Broken Hill sub-domain retrograde schists zones were initiated. The axial trends and plunge directions of F3 folds are somewhat variable throughout the province reflecting in part the effects of the weak overprinting F4 deformation and in part the variable nature of F3 strain (as is also evidenced by the production of local F3 shear zones). F3 folds generally do not have strong axial planar fabrics by contrast with F2 folds (Fig. 5.7).
323
Fig. 5.7. Tight recumbent F, folds refolded by upright F3, central Eyre Peninsula (Fig. 5.2). Photograph, A.J. Parker.
Fig. 5.8. Amphibolite pod in strongly mylonitized gneisses o n coast of Eyre Peninsula (Fig. 5.2). Photograph, A.J. Parker.
324 They also have influenced the localization of orebodies and in the Gawler domain the iron ores of the Middleback Ranges lie in an F3 synform. Glen et al. (1977) note that the province was characterized by a high geothermal gradient and that there is general correlation between metamorphic grade and height in the structural succession. In the NW part of the Broken Hill subdomain, Glen (1978a) observed that metamorphic boundaries were subparallel t o bedding and folded by F1 folds. He considered that there was some modification of the metamorphic boundaries by this and subsequent deformations, but that the major metamorphism in that area antedated the F1 deformation. Consequently, he concluded that the basal rocks of the local succession were higher grade and more penetratively deformed than the upper succession. Similar relations may well hold elsewhere in the province for the M, M , metamorphism but it is clear that metamorphic boundaries are not always bedding parallel and that subsequent metamorphism can be locally more important. In the higher-grade parts of the Broken Hill region sillimanite growth locally continued into the F3 deformation and isograds clearly transgress the early formed structures. In the Olary sub-d.omain the iron formations (Unit 2) are present in both higher- and lower-grade regions. Likewise in the Gawler domain amphibolite facies iron formations of the Middleback Ranges can be traced south into granulite facies terrain (Fig. 5 -8). Furthermore, metamorphic boundaries in that domain locally transgress major F, fold structures thus necessitating post M I modification of the original, perhaps bedding parallel, metamorphic boundaries. Superimposed metamorphism is also important in the Olary sub-domain where fibrolite overprinting andalusite has already been noted.
Late tectonic history of the Gawler Province Following the F3 deformation there was a major disruption t o the parallelism of tectonic events in the Gawler, Willyama and Musgrave domains. In the Gawler domain the eruption of the Gawler Range Volcanics, the synchronous deposition of the predominantly clastic Moonabie Formation, Tarcoola Beds and Corunna Conglomerate and the subsequent co-magmatic intrusion of the post-tectonic granites define a major tectonic episode which represents the transition from orogenic t o cratonic character (cf. Tectonic Map of Australia and New Guinea, G.S.A., 1971). In the Kokatha and Tarcoola areas, the Gawler Range volcanic complex rests directly on Archaean rocks, demonstrating that uplift and erosion of the entire Hutchison Group took place prior t o their deposition. This uplift had already begun by F3 times since the metamorphic conditions during F, indicate uplift of at least 4-6 km after F, and Fz (Parker, 1978). The volcanics are relatively flat-lying and undeformed and form a large basin-like feature surrounded by a ring of post-tectonic granites approximately in the centre of the Gawler domain (Fig. 5.1). They were erupted
325 subaerially in a continental environment and consist essentially of acid ignimbrites with minor lavas and agglomerates. In the Kokatha area, Branch (1978) recognizes several phases of volcanic activity (collectively called the Chitanilga Volcanic Complex) initiated by the development of large stratovolcanoes at 1525 5 15 Ma on a folded Archaean basement. This initial phase was dominated by basaltic and andesitic flows with some acid ignimbrites but the subsequent premonitory caldera, caldera formation and post-caldera phases are dominated by the eruption of thick rhyodacite t o dacite ignimbrites. The recognition of the calderas themselves is tenuous because of the outflow of a thick sheet of dacite during the so-called caldera phase. However, there are several circular features tens of kilometres across with volcanic units gently dipping towards their centres suggestive of a caldera origin (Crawford, 1963a). The age of the thick sheets of dacite (Chandabooka Dacite and Yardea Dacite) is 1511 ? 36 Ma. The caldera phase is succeeded by thin rhyodacites and breccias. and the whole sequence is intruded by the Hiltaba Granite suite at 1478 5 38 Ma and by rhyodacite dykes at 1457 22 Ma (Blissett, 1977; Branch, 1978). The relationship between the volcanics of the Gawler Ranges and clastic sediments is not clearly established in that area. However, in the Moonabie Range 50 km southwest of Whyalla (Fig. 5.2) volcanics which have been likened to those of the Gawler Ranges (C. Giles, pers. commun., 1978) are interlayered with volcanoclastic sediments of the Moonabie Formation (Fig. 5.9). Overlying the Moonabie Formation is a sequence of fluvio-marine conglomerates and sandstones, the Corunna Conglomerate sequence, and both sequences have been intruded by the Charleston Granite dated by Rb-Sr methods at 1445 5 39 Ma (N.B. K-Ar gives an anomalous age c. 1560 Ma). Deposition of the Moonabie Formation and Corunna Conglomerate is largely controlled by fault graben which are believed to have been developed by reactivation of F3 structures since they parallel F3 fold trends. F, deformation features are present in both the basement and Middle Proterozoic sediments and local K-Ar dates of c. 1470 are common in the basement (Parker, 1978). This date is believed t o represent the cessation of tectonic activity in the Gawler domain. Subsequent volcanism, the Roopena Volcanics and Beda Volcanics, and Adelaidean sedimentation took place on a stable cratonic platform, the Stuart Shelf (Fig. 5.1). It is significant to this point that Rb/Sr and K/Ar Delamerian ages (c. 490 Ma) have not been recorded in either Olarian basement or Adelaidean cover rocks of the Gawler Platform west of the Torrens Hinge Zone. In contrast, tectonic events in the Willyama domain continued through t o Delamerian time and Delamerian K/Ar ages are obtained from basement rocks. Adelaidean rocks resting on the Willyama domain are strongly deformed and they show large thickness changes across faults believed to be extensions of retrograde schist zones. In fact the retrograde schist zones are believed to have been intermittently reactivated from F3 time through to
*
326 East
West IRON DUKE
MOONABIE RANGE
CORUNNA CONGLOMERATE
r\
MOONABIE FORMATION
pebbly sandstone
g r i t t y sandstone
breccia
Fl
VOLCANICS
c o n g lo rn e ra t e Z,ooo
basalt
r
1
dacite and andesite
V
r h y o l i t e and r h y o d a c i t e
0 0
1
0
1
/ , ) , , I
I
nCRS
KILOMETRES 2 3
4
5
I
I
I
I
79-189
8A
Deportment of M
ond Ene 9)
Fig. 5 . 9 . Schematic section west of Whyalla (Fig. 5 . 2 ) showing rock relationships of units representing the late-tectonic history of the Gawler domain. After N. Lemon, reproduced by permission of Broken Hill Proprietary C o . Ltd.
Delamerian time. Furthermore, Berry et al. (1978) have recorded five deformations in basement rocks north of Olary in the Willyama domain, and they identify the F4 and F, deformations with deformation of the Adelaidean sediments. Consequently the final development of a stable craton in the Willyama area did not take place until c. 490 Ma. This also applies to the other ,smaller, basement inliers east of the Torrens Hinge Zone.
Post-1400 Ma evolution In the Musgrave orogenic domain, the Early t o Middle Proterozoic orogenic complexes were subjected t o strong orogenic reworking between 1400 and 1000 Ma. This domain was part of the Musgrave-Fraser orogenic province. During the same period, the newly established Gawler platform remained relatively stable but became subject to basic volcanism and limited, mainly continental, sedimentation. At about 1000 Ma both regions were the sites of deposition of blanket quartzites which were followed by the development of the Amadeus Basin and the Adelaide “geosyncline”. In the following account the Adelaidean
327 deposition in South Australia will be treated, together with the earlier platform development. For a brief account of the Musgrave orogenic domain the reader is referred to Chapter 1. THE STUART SHELF AND ADELAIDE GEOSYNCLINE
*
Tectonic framework Sprigg (1952) employed the term Adelaide Geosyncline and regarded its western boundary with the “Stuart Stable Shelf” as a major N-S lineament, the Torrens Lineament. Thomson (1969, 1970, 1976) recognized a transitional depositional zone between the thin near-horizontal Stuart Shelf sequences and the thick, more complete Late Proterozoic and Cambrian sequences to the east which have been relatively strongly deformed by the Cambro-Ordovician Delamerian Orogeny to produce a complex, sinuous and branching system of folds. This transitional zone he called the Torrens Hinge Zone, bounded on the west by the Torrens Lineament and on the east by the fault system along the western margin of the Mt. Lofty and Flinders Ranges. The rocks of this hinge zone are less strongly folded than those t o the east and thus there is a broad correspondence between three principal structural and tectono-stratigraphic zones, the Stuart Shelf, the Torrens Hinge Zone, and the Adelaide Geosyncline (Fig. 5.10). At their northern extension, the Torrens Hinge Zone and adjacent foldbelt curve t o the northwest, and the Adelaidean sediments here were probably partly continuous with those of the Officer and Amadeus Basins of central Australia. Between the central Flinders Ranges and the Muloorinna Ridge (Fig. 5.11) of interpreted shallow basement (Thomson, 1970) a fold belt extends from the Mount Painter basement inlier in the east through the north Flinders Ranges to the Peake and Denison Ranges in the northwest. In the south, the fold belt swings westerly through Kangaroo Island, to the west of which it is truncated by the present continental margin. The present morphological expression of the Delamerian fold belt is largely due to Cenozoic, and especially Late Cenozoic, tectonics. Erosional etching in a semi-arid climate, following Cenozoic block uplift and tilting, allows the structure to be particularly well seen in aerial and satellite photographs and, because of this, the mapping of the belt at 1 : 250,000 scale or better has been rapidly achieved. Stratigraphic successions are best studied in the banks of the principal drainage channels. In contrast, the flat-lying rocks of the Stuart Shelf are largely obscured by thin younger cover. The Torrens Hinge Zone was the site of Cenozoic graben structures, complimentary t o the uplift of the ranges to the east. Consequently, the Pre-
* by W.V. Preiss, R.W.R. Rutland and B. Murrell.
328
100
200
I 1 --
km I
Fig. 5.10. General geology of t h e Adelaide Geosyncline, Torrens Hinge Zone, and Stuart Shelf, showing t h e principal tectonic subdivisions of t h e Delamerian fold belt. Note t h a t t h e zones shouth of the Central Flinders zone, together form the Mt. Lofty-Olary fold belt, which includes t h e Nackara and Fleurieu arcs.
329
Fig. 5.11. Distribution of sedimentary facies during various stages of Adelaidean deposition.
330 Cambrian geology is largely concealed here also. Recent drilling programmes by mining companies and by the S.A. Geological Survey have, however, provided much information on stratigraphic relationships between the three principal zones. The Delamerian fold belt can be conveniently divided into a number of subsidiary structural zones, which again have considerable tectono-stratigraphic significance (Fig. 5.10). An area with only broad open folding (the Central Flinders Zone) separates the North Flinders fold belt, in the north, from the main Mt. Lofty-Olary fold belt in the south. This Central Flinders zone has fault boundaries on both E and W but provides a relatively stable link between the Stuart Shelf to the west and the Curnamona Cratonic Nucleus (Thomson et al., 1976) t o the east. The North Flinders belt is bounded by Precambrian fault structures, rejuvenated in the Cenozoic, on the SE and SW. In the north, however, the boundary is provided by the feather edges of Cretaceous and Cenozoic transgressive sequences. This younger cover obscures the Adelaidean rocks, and the Muloorinna basement ridge, which extends from the Mt. Painter inlier WNW to the Denison inlier, as indicated by gravity and magnetic evidence (Thomson, 1970). The Mt. Lofty-Olary fold belt is divided by an anticlinorial zone, which joins the Olary basement zone to the basement inliers of the Adelaide region, one of which, the Houghton Inlier, gives its name to the zone. NW of this zone the South Flinders zone displays labile shelf sedimentation and is little metamorphosed. SE of the zone, a more basinal sedimentary facies is developed (Preiss, 1973 and see below) and the rocks become significantly metamorphosed (McKirdy et al., 1975) in the inner part of the Nackara arc. The Cambrian rocks in the outer part of the Fleurieu arc are even more strongly metamorphosed. In general Adelaidean rocks east of the Flinders zones are concealed by younger cover around Lake Frome. This area is interpreted as the stable Curnamona Cratonic Nucleus. Adelaidean rocks east of the Mt. Lofty-Olary fold belt are concealed by younger cover of the Murray Basin and perhaps also by metamorphosed Cambrian rocks like those of the Fleurieu arc. Northeast of the Olary basement zone, however, the Broken Hill zone displays the most easterly outcrops of Adelaidean rocks and their basement. Principal problems The main Mt. Lofty-Olary fold belt is very similar in shape and size t o the Jura fald belt and, like the latter, it is a belt of foreland folding adjacent to a major orthotectonic belt. However the Adelaide Geosyncline displays thicker and more varied sedimentary facies which may partly account for significant differences in the deformation. The diapiric structures of the Adelaide Geosyncline in particular appear to be unique.
331 The main folding in the belt occurred at about the beginning of Ordovician time. It is earlier than the principal deformations of the Lower Palaeozoic rocks of the Lachlan belt to the east, but later than the late Precambrian orogenies recorded in Tasmania (Williams, 1978). The existence of late Precambrian orogeny east of the Delamerian orogeny is an apparent exception to the general rule of eastward migrating orogeny and progressive cratonization in eastern Australia since the Ordovician (e.g., Rutland, 1976). The Tasmanian evidence indicates that an active continental margin already existed in late Precambrian time although that margin lay east of Tasmania, and the orogeny may have been of either collision or Cordilleran type. The main problem of the rocks of the Delamerian fold belt therefore, concerns their relationship t o continental or plate margins. Sprigg (1952) suggested that the Adelaidean sediments are essentially miogeosynclinal and were developed as a continental terrace, implying the existence of a continental margin immediately to the southeast. In current terms, Sprigg interpreted the Adelaidean as a miogeoclinal wedge developed across a passive plate margin. He in fact suggested that “many Beltian type miogeosynclines originated as continental terraces” and he thus anticipated more recent interpretations of the Cordilleran region (Dietz, 1963; Stewart, 1972). More recently there has been a tendency to place more emphasis on the evidence from the Flinders zones and to interpret the Adelaide “Geosyncline” as an aulacogene (e.g., Rutland, 1973a; Scheibner, 1973) rather than a miogeoclinal wedge. It has also been argued that a continental margin was produced by rifting and drifting in the Cambrian (Veevers and McElhinny, 1976) but this denies any relationship of the Adelaidean rocks to a passive continental margin, unless an earlier Wilson cycle, involving a passive continental margin and late Precambrian collision is also envisaged. These hypothesis will be discussed after the stratigraphic and sedimentological evidence has been summarized below. A related problem, which affects the stratotectonic interpretation, is whether parts of the fold belt have suffered rotations, or translations, relative to each other, or to other parts of the continent. Crawford and Campbell (1973) invoked both oroclinal bending and large translations on major strike-slip faults, while Coward (1976) suggested a quite different interpretation involving 150 km of sinistral simple shear across the fold belt. Arguments against these interpretations have been provided by Daily et al. (1973) and Glen et al. (1977). This problem will be discussed after the main characteristics of the folding and faulting have been summarized in a further section below.
Stratigraphic systematics The Adelaide Geosyncline is the stratotype for the Adelaidean Period, the time during which the Precambrian sediments of the Adelaide Geosyn-
332 cline were deposited (Thomson, 1969). It contains in general the thickest and most complete record of Adelaidean sedimentation in Australia. In its type area near Adelaide, the Adelaide(an) System was originally divided into the Torrensian (oldest), Sturtian (containing glaciogene sediments) and Marinoan (youngest) Series (Mawson and Sprigg, 1950), while Sprigg (1952) added the pre-Torrensian Willouran Series in the northern regions. Today these terms are used in a chronostratigraphic sense as subdivisions of the Adelaidean (Thomson, 1969), providing a useful time framework for the discussion of geological history, although their actual boundaries can be located precisely only in the stratotype area. As presently defined, the Torrensian-Sturtian and Sturtian-Marinoan boundaries fall within continuous stratigraphic sequences, thus conforming to Hedberg’s (1976) recommendations. The Willouran-Torrensian boundary is, however, not yet well defined, since in the stratotype area the base of the Torrensian of Mawson and Sprigg (1950) coincides with a transgressive basal unconformity, and Willouran sediments are absent. Thomson et al. (1964) proposed a lithostratigraphic subdivision of Adelaidean rocks that has been used by most subsequent authors. The oldest sequence, the Callanna Beds of Willouran age, is incompletely preserved and partially disrupted in isolated anticlinal cores, fault blocks and diapirs. The better preserved Burra Group includes all type Torrensian rocks, the Torrensian-Sturtian boundary, and earliest Sturtian sediments. The Umberatana Group is everywhere unconformable on older units, mostly Burra Group but also locally Callanna Beds or crystalline basement. It is of Sturtian to early Marinoan age, and contains widespread glacials at its base and more restricted glacials at the top (Sturtian and Marinoan glaciations respectively), with a major interglacial episode between them. The postglacial Wilpena Group of late Marinoan age rests on the Umberatana Group mostly conformably, and represents a major transgression on to the Stuart Shelf. Early Cambrian deposits everywhere disconformably overlie older units, generally Wilpena Group. The stratigraphic and geographic relationships of the groups and their constituent formations t o each other and to the chronostratigraphic divisions are represented in Table 5 .I. Without conventional biostratigraphic control and without unique isochronous marker beds, precise time correlation is not possible, and Daily (1963) has argued against the use of chronostratigraphic subdivisions. Other authors have retained them while acknowledging their limitations e.g., Dalgarno and Johnson (1968). Correlation within the Adelaide Geosyncline is achieved by mapping and lithostratigraphic comparisons, tracing distinctive marker beds and sequences as far as they persist. Certain lithologies characterize particular stratigraphic levels, though not always uniquely, e.g., Willouran volcanics and replaced evaporites, Torrensian sedimentary magnesite, Sturtian and early Marinoan tillites. Tillites and associated rock types are the best markers for lithostratigraphic correlation with other Adelaidean
333
334 sequences in Australia. Major uncertainties remain, however, in the subdivision and correlation of lower Adelaidean rock units, especially the Callanna Beds. Preiss (1977) summarized the application of stromatolite biostratigraphy to the problems of correlation of late Precambrian sequences. It was concluded that certain Willouran stromatolites suggest correlation with the Late 50 Ma and 680 50 Ma) Riphean of the USSR (dated at between 950 whereas those from the Torrensian Skillogalee Dolomite (Table 5.1) most resembled late Middle Riphean forms. In view of geochronology then available, and of the uncertainty in the total time range of the Middle Riphean forms in the USSR, correlation of the early Adelaidean with part of the Late Riphean was preferred, representing an extension in the range of the Middle Riphean forms. Sturtian-Marinoan forms from the Umberatana Group closely resemble forms known from Late Riphean and Vendian deposits of the USSR, suggesting that they fall close t o the Riphean-Vendian boundary (680 50 Ma). The uppermost Adelaidean Pound Subgroup contains the well-known Ediacara assemblage of soft-bodied metazoan fossils. Glaessner (1971) reviewed world-wide occurrences of elements of this assemblage and demonstrated their restriction to latest Precambrian sequences younger than about 680 Ma. Geochronological control on the age of the Adelaidean is still very limited, because few of the rocks are suitable for dating and because of widespread Delamerian overprinting. Attempts to date pre-Adelaidean basement and Willouran volcanics have produced equivocal results. Cooper and Compston (1971) interpreted a Rb/Sr age of 849 31 * Ma as representing the age of high-grade metamorphism in pre-Torrensian basement near Adelaide. Compston et al. (1966) and Cooper (1975) suggested an age of approximately 800 Ma for the Willouran Wooltana Volcanics of the Adelaide Geosyncline, though Thomson (1966a, b) had correlated them with the approximately 1300 Ma-old Roopena Volcanics on the Stuart Shelf. Mason et al. (1978) prefer to correlate the Willouran Volcanics with the newly defined Beda Volcanics, also on the Stuart Shelf. Webb and Horr (1978) published a preliminary age estimate for the Beda Volcanics, which has now been revised to 1076 f 34 Ma (Thomson, 1980). The conflict between these data cannot at present be resolved, but it is likely that Adelaidean sedimentation commenced between 1100 and 800 Ma ago, although the platform history of the Stuart Shelf had begun considerably earlier with the extrusion of the Roopena Volcanics and deposition of sediments associated with them.
*
*
*
* Recalculated
using XRb87 = 1.42 X
lo-"
*
335 History of sedimentation Willouran The Adelaide Geosyncline was initiated at the beginning of Willouran time with widespread deposition of basal quartzites and carbonates, followed by basic volcanism representing a phase of tensional rifting. Fine to coarse clastics, largely immature, and thin bedded carbonates were then deposited under shallow water and evaporitic conditions, possibly in restricted basins. The Adelaide and northern Olary regions were probably emergent basement, as they contain no preserved Willouran sedimentary record (Fig. 5.11a). Although the Willouran Callanna Beds were defined in the Willouran Ranges, the basal relationships are not clear in that area. Consequently, theMount Painter region was chosen as stratotype for the lower Callanna Beds (Coats, 1971). There the basal Paralana Quartzite (maximum thickness 1100 m) is a clean cross-bedded, partly heavy-mineral laminated and pebbly quartzite with minor siltstone and local talus breccia formed in fault-angle depressions. It was deposited under fluviatile t o shallow marine conditions (Coats, in prep.). Locally, stromatolites occur that were formed by the trapping of quartz grains by algal mats. The Wywyana Formation (100 m ) of calcitic and dolomitic marble, calc-silicates and minor siltstone follows conformably, and these two formations partly envelop the Mount Painter basement inlier. Early Willouran volcanism is then represented by altered amygdaloidal basalts, trachytes, andesites and minor rhyolites of the Wooltana Volcanics, attaining a maximum thickness of 2500 m (Crawford, 1963b; Coats, in prep.). Red quartzites, siltstones and tuffs are associated locally. These three Willouran formations of the Mount Painter region have close analogues in Western New South Wales, (Lady Don Quartzite, stromatolitic Boco Formation dolomites, and Wilangee Volcanics: Cooper and Tuckwell, 1971) and in the Peake and Denison Ranges of far northern South Australia (basal conglomerate, stromatolitic dolomite and basic volcanics: Ambrose and Flint, 1979). In all three regions, the basal formations rest with pronounced unconformity on pre-Adelaidean metamorphic basement. Volcanics petrologically similar to the Wooltana Volcanics occur unconformably beneath the Burra Group at Depot Creek (southwest Flinders Ranges), and as blocks out of stratigraphic context in disrupted and partly diapiric sequences throughout the Flinders Ranges, suggesting originally widespread extrusion of flood basalts. The Beda Volcanics on the eastern Stuart Shelf are also possible equivalents (Mason et al., 1978). On the Stuart Shelf, the widespread red, cross-bedded, lithic, feldspathic and kaolinitic sandstones of the Pandurra Formation were deposited, perhaps at approximately the beginning of Willouran time, in fluviatile environments on down-faulted basement blocks (Mason et al., 1978), attaining a maximum thickness of over 600 m and lapping out on to crystalline basement to the west (Thomson, 1976). Fossil liesegang weathering rings are characteristic. The Pandurra Formation rests unconformably on the approximately 1300 Ma-old Roopena Volcanics, and is itself disconformably overlain by the Beda Volcanics and associated coarse clastics of the Backy Point Beds (Mason et al., 1978). During the later Willouran, sedimentation was restricted to the Adelaide Geosyncline, excluding the emergent basement of the Adelaide and Olary regions. A common association is carbonaceous laminated pyritic siltstones and shales, micaceous and silty thinly bedded immature sandstones (often with mud-cracked, ripple-marked shaly partings and halite casts), cross-laminated sandstones (in places with heavy mineral bands), grey limestone and thinly bedded buff dolomites with mainly stratiform stromatolites, tepee structures, gypsum pseudomorphs and local fenestral fabrics. Low-energy, evaporitic, tidal-flat to subtidal basin conditions are indicated (Preiss, 1978). Broadly similar lithological asso-
336 ciations have been recognized in very thick sequences in the Peake and Denison Ranges (Thomson and Coats, 1964; Ambrose and Flint, 1979), through the Willouran Ranges (Thomson and Coats, 1964; Murrell, 1977) to the Spalding Inlier (Preiss, 1974), and in disrupted zones and “diapirs” throughout the Flinders Ranges (Coats, 1964a, 1965; Dalgarno and Johnson, 1968; Mount, 1975; Preiss, 1978). Detailed correlations are not yet possible between these sequences, partly because of post-depositional disruption, and partly because of primary facies changes that may reflect deposition in only partially connected basins. Relationships in the Peake and Denison and Willouran Ranges indicate that these sequences post-date the main Willouran Volcanics. Thin basalts in the River Broughton Beds of the Spalding Inlier (Preiss, 1974) are possibly younger than the Wooltana Volcanics, while Kitch ( 1975) records volcaniclastics in possible Willouran sediments in the southern Flinders Ranges. Local volcanic highs probably existed in marginal areas, e.g. Depot Creek and Mount Painter, where the Burra Group directly but unconformably overlies the Willouran Volcanics. Evidence for a more widespread unconformity in this stratigraphic position is as yet inconclusive, though pre-Torrensian deformation of the Callanna Beds has been suggested by Kitch (1975) and Murrell (1975). In the Willouran Ranges the contact between the basal parts of the Burra Group and the partly disrupted Willouran rocks is more extensive (up to 100 km of strike length) than elsewhere. The evidence adduced by Murrell(l977) in support of a widespread basal Burra Group overlap and unconformity over a previously disrupted Willouran sequence is circumstantial, though consistent with continued halfgraben development during sedimentation of the Burra Group. The boundary zone commonly shows evidence of faulting, consistent with continuing block movements. Stratigraphic thinning and persistent development of near shore facies in the Burra Group against boundaries with Willouran rocks are difficult to explain without invoking an unconformity. The disposition of Burra Group sediments supports this interpretation (Fig. 5.13). However, the lack of unequivocal sedimentary contacts and the chaotic and disharmonic deformation of the Callanna Beds (probably originally incompetent due to contained evaporites) make the hypothesis of an unconformity between Callanna Beds and Burra Group difficult to prove. The structural style of the Callanna Beds has been ascribed to later pulses (Sturtian to Cambrian) of diapiric intrusion into overlying sequences, culminating in the Delamerian Orogeny (Webb, 1960; Coats, 1965; Dalgarno and Johnson, 1968).
Torrensian During Torrensian time, the greater part of the Burra Group (Mirams and Forbes, 1964) accumulated in more continuous basins, by transgressions and regressions across a possible evolving continental shelf. The Gawler Craton, including the Stuart Shelf, contains no preserved Torrensian sediments but was a source of clastic sediment that entered the Adelaide Geosyncline, probably as a series of eastward-prograding deltaic complexes. Another source area probably existed to the north around the Muloorina Ridge. No Torrensian sediments are preserved in the Broken Hill region. Clastic sedimentation was episodically superimposed by progradation on widespread carbonate shelf sedimentation characterized by “primary” dolomite without limestone, and by the presence of sedimentary magnesite and diagenetic black chert (Fig. 5.11b). The base of the Torrensian was placed by Mawson and Sprigg (1950) at the unconformable base of the Aldgate Sandstone on metamorphic basement
337 near Adelaide. Equivalents of the Aldgate Sandstone are widespread, but relationships with underlying Willouran rocks are preserved in only a few areas. At some of these, the Aldgate Sandstone equivalents rest unconformably on basic volcanics (e.g., Depot Creek, Mount Painter) while in others they appear to pass gradationally down into Willouran sequences younger than the main volcanics (e.g., Spalding). Despite these uncertainties, it is known that Torrensian sedimentation commenced with widespread transgressive clastic deposition. Near Adelaide, the Aldgate Sandstone (Mawson and Sprigg, 1950) is up to 1200 m thick and consists of a lower heavy-mineral laminated, cross-bedded pebbly arkose, deposited under fluviatile to tidal flat conditions and an upper, fine- to coarse-grained feldspathic sandstone, wavy- and cross-bedded, possibly reflecting tidal conditions (Forbes, in prep.). In the southwest Flinders Ranges, the equivalent Emeroo Quartzite (Mawson, 1947) is lithologically similar and up to 1400 m thick : a lower heavy mineral banded conglomeratic, feldspathic sandstone is separated from an upper more mature feldspathic quartzite by silty and micaceous sandstones with salt casts (the Wirrabara Formation of Mirams and Forbes, 1964). Towards the east and southeast, the Emeroo Quartzite passes laterally into two major arenites separated by the silty and dolomitic River Wakefield Subgroup (Table 5.1). The lower arenite, the Rhynie Sandstone (Wilson, 1952) corresponds in lithology and stratigraphic position to the lower Algate Sandstone and is probably largely of fluvial origin, although local herringbone cross-beds and salt casts suggest intermittent tidal flat conditions. The River Wakefield Subgroup (Wilson, 1952; Forbes, 1977) is about 700 m thick and is interpreted as an offshore equivalent of the Wirrabara Formation or middle Emeroo Quartzite. It includes carbonate platform sediments that are largely indistinguishable from the younger Skillogalee Dolomite (locally with similar sedimentary magnesite), and wavy and lenticular-bedded, fine sandy siltstones that were probably deposited at the fronts of prograding deltas. The upper arenite, the Bungaree Quartzite (Forbes, 1977) is a 500 m-thick feldspathic, cross-bedded quartzite lacking heavy-mineral bands, with minor dolomite interbeds. It is a coarser, more proximal facies that prograded eastward over River Wakefield Subgroup sediments. Transitions from Rhynie Sandstone to River Wakefield Subgroup to Bungaree Quartzite thus record first transgressive, then regressive sedimentation. Similar events are recorded in the Mount Painter region (sequence Humanity Seat Formation to Wortupa Quartzite: Table 5.1) and in the Willouran Ranges (Murrell, 1975). Following early Torrensian clastic sedimentation, a wide carbonate platform developed over most of the area of the Flinders Ranges. There was still a persistent sand influx from the western and northern source areas, and in the most marginal western zone, only sandstones are represented at this interval. Near Adelaide, platform carbonate deposition was more local, and gave way to basinal black shale deposition to the southeast. The term Skillogalee Dolomite (Wilson, 1952) has been used by most authors for these carbonates throughout the Flinders Ranges, though Uppill (1979) proposes new nomenclature. The Skillogalee Dolomite attains a thickness of up to 4000 m in the Willouran Ranges, where it passes eastwards into a basinal black shale facies. It is characterized in its lower member by pale coloured dolomites, commonly stromatolitic, with variable admixtures of arkose and siltstone, and in its upper member by dark grey, carbonaceous, laminated dolomites, occasionally stromatolitic, carbonaceous shales, bedded and reworked conglomeratic magnesite, sandstone, and early diagenetic black cherts. Silicified stromatolites are commonly microfossiliferous (Schopf and Barghoorn, 1969; Schopf and Fairchild, 1973). The Skillogalee Dolomite represents deposition in lagoonal and shallow subtidal to intertidal settings. Mudcracks record periodic subaerial exposure. The lower part shows local evidence of tidal influence (Uppill, 1979), with reducing conditions during deposition of fine sediment but early diagenetic oxidation of most dolomite. The upper
338 part records almost entirely reducing conditions, the dolomites forming in mainly low energy subtidal areas and magnesite in marginal zones (Forbes, 1961; Uppill, 1979). A single type of stromatolite grew under both high- and low-energy conditions (Preiss, 1973). In western marginal areas, sand deposition persisted after Skillogalee Dolomite deposition. In the southeast, the silty Woolshed Flat Shale (300 m thick; Wilson, 1952) and overlying 200-600 m and Stonyfell and Undalya Quartzites (Mawson and Sprigg, 1950; Wilson, 1952) are lithologically very similar to the silts of the River Wakefield Subgroup and sands of the Bungaree Quartzite, respectively. The sequence that follows the Skillogalee Dolomite is thus analogous to the earlier Torrensian prograding deltaic cycle. In the southern regions, this style of sedimentation persisted through the remainder of the Torrensian and earliest Sturtian. Thus the 600 m carbonaceous dolomites and dolomitic siltstones (both dominantly subtidal) of the lower Auburn Dolomite (Wilson, 1952) represent the next transgressive phase, and the eastward-lensing Watervale Sandstone Member (Wilson, 1952; Mirams and Forbes, 1964) the next progradation. The upper Auburn Dolomite (600 m ) and overlying siltstone of the Saddleworth Formation (Mirams and Forbes, 1964), locally up to 4000 m thick, represent a major transgression in the area between Adelaide and the southern Flinders Ranges, at the end of Torrensian time. In the northern areas, the deltaic cycles are not so obviously repetitive. Deposition of the Burra Group in the Willouran Ranges (Murrell, 1977) was apparently restricted to a northwest- to southeast-trending trough less than 5 0 km wide. The early depositional history shows a developing half-graben between a hinge a few kilometres to the southwest of the Norwest Fault and a faulted margin along the East Willouran Fault. Marginal elastic wedges are associated with both sides of the trough in the early part of the sequence and it is considered that these were the approximate limits of sedimentation. Later Torrensian sediments overstepped the earlier deposits to a small extent. As the basin subsided it was broken into a complex of second order half-grabens by movements on intrabasinal faults. Rotations of individual blocks (all northeastern edges downthrown) resulted in bald-cap highs, facies and thickness changes and the formation of growth folds in the accumulating sediments. Deposition of the upper Burra Group was dominated by silts and fine sands, with thin quartzite and dolomitic intercalations (including the 1100 m thick Myrtle Springs Formation of Coats, 1973). Some of the dolomites resemble Skillogalee facies, and are partly stromatolitic, especially in the western Willouran Ranges and Peake and Denison Ranges.
Sturtian The Sturtian was conceptually linked by Mawson and Sprigg (1950) to glaciation, but it is now considered that pre-glacial sedimentation (final phases of the Burra Group), two distinct glacial phases, and post-glacial transgression and regression took place during Sturtian time. Mawson and Sprigg (1950) placed the Torrensian-Sturtian boundary at the base of the Belair Subgroup, in the belief that its fresh feldspar content and varve-like lamination are related t o glaciation. It is now considered that the Belair Subgroup is genetically related to the underlying Burra Group (Coats, 1967), within which it is now included (Binks, 1971). However, its conformable base remains a convenient position for the chronostratigraphic boundary. The Belair Subgroup, up t o 3600 m thick and consisting of alternating feldspathic coarse quartzites and dark coloured siltstones and slates, may
339 represent the final deltaic progradation of the Burra Group. Coats (1967) recorded abundant mudcracks, suggesting very shallow water environments. Slumping (Forbes, in prep.) and sand dykes (Sprigg, 1946) indicate syndepositional instability. Finely laminated siltstones represent lowenergy subtidal conditions. The Sturtian to early Marinoan Umberatana Group (Coats, 1964b) contains all known glaciogene sediments. The earliest of these OCCUT only in the northern and eastern regions (Fig. 5.11c), and are not represented in the stratotype near Adelaide, where Coats (1967) mapped an unconformity at the base of the Sturt Tillite (Howchin, 1920; Mawson and Sprigg, 1950). All subsequent mapping has demonstrated the Umberatana Group t o be regionally unconformable on the Burra Group. It has been suggested that the first phases of diapirism, and local block faulting, were associated with this early Sturtian tectonism (Binks, 1971; Coats, 1973). In the Olary region, the Pualco Tillite (Forbes and Cooper, 1976) occurs at the base of the Umberatana Group. It is 3300 m thick in its type section, but contains numerous quartzite and laminated siltstone intercalations. Glaciomarine deposition from an extensive floating ice sheet is inferred (Coats, in prep.). The overlying Benda Siltstone (Forbes, 1970) contains limestone and quartzite interbeds and occasional pebbles, suggesting basinal marine sedimentation with limited glacial influence. The Braemar Iron Formation of Whitten, (1970) is now regarded as a ferruginous facies variant of both Pualco Tillite and Benda Siltstone (Forbes and Cooper, 1976) and represents chemical precipitation of iron oxides and minor dolomite during glacial and interglacial phases. These glaciogene sediments are unconformably overlain by the Wilyerpa Formation, containing local tillite and representing the younger Sturtian glacial phase (Coats and Forbes. 1977). The younger phase of Sturtian glaciation is recorded by the marine glacial Sturt Tillite (400 m) near Adelaide, Appila Tillite (1500-2100 m) of the SW Flinders Ranges, Merinjina Tillite (600 m) of the northern Flinders Ranges (Coats, in prep.) and probably the McLeay Beds of the Stuart Shelf, all of which were deposited on low-relief unconformity surfaces. The Appila Tillite passes eastwards into basinal marine sediments - the alternating sandstones and laminated siltstones of the Wilyerpa Formation (Dalgarno and Johnson, 1966) with a maximum thickness of 2800 m. Tillites and glacial erratics are restricted t o the basal parts of this sequence which itself rests unconformably on the older Sturtian glacials. The thickness variations and distributions of tillites and basinal clastics imply westerly ice derivation and deposition in a series of marine troughs, with periodic advance and retreat of a floating ice sheet (Coats, in prep.). The 6000 m-thick glacials of the Yudnamutana Subgroup of the northernmost Flinders Ranges, originally regarded as partly equivalent t o the Sturt Tillite (Coats, 1964b) have been reinterpreted by Coats (1973; in prep.) as marking the earlier Sturtian glacial phase. The younger phase would then be recorded by an unconformably overlying quartzose conglomerate in that area (Table 5.1). The Yudnamutana Subgroup resembles the Pualco TilIiteBenda Siltstone sequence in having initial tillites followed by basinal marine sediments with limited glacial influence, but differs in lacking the characteristic iron formations. Murrell et al. (1977) prefer the original correlation with the younger Sturtian glacial phase. Deposition was restricted t o a faultcontrolled e a s t w e s t trough, and transgressed on t o the basement of the Mount Painter inlier, which was a source of clasts only in the basal boulder beds but not in the overlying tillites (Coats, in prep.). The thickest developments of lower Sturtian glacials are in eastern zones, and Murrell (1977) has inferred an easterly provenance for them.
340 Sturtian glacials are also known from the Officer Basin (northwestern South Australia) and from north and south of Broken Hill (western New South Wales). Iron formations associated with tillites of the Officer Basin and south of Broken Hill suggest lithostratigraphic correlation with the earlier Sturtian phase (Coats, in prep.). Basalts conformably overlie Sturtian tillites in the Officer Basin (Krieg and Major, 1972). Following the second Sturtian glacial phase, there was widespread marine transgression, represented by the deposition of up to 2000 m of very finely laminated carbonaceous and pyritic siltstones of the Tapley Hill Formation (Howchin, 1904; Coats, 1964b) throughout the whole Adelaide Geosyncline, much of the Stuart Shelf, western New South Wales and the eastern Officer Basin. Characteristic laminated cabonaceous dolomite and shale occur in the basal few metres. Most of the remainder of the formation has very even, uniform, carbonaceous laminations up to 0.5 mm thick, implying relatively deep water deposition below wave base. Local starved ripples near Adelaide indicate southeast flowing downslope currents (Preiss, 1973). Coarser elastic interbeds (coarse siltstone, greywacke, conglomerate) in the central Flinders Ranges imply a local intrabasinal derivation, perhaps from submarine fault scarps or rising diapirs. They were possibly deposited partly by turbidity currents. On the Stuart Shelf there was generally black shale deposition, although locally the oolitic and stromatolitic Woocalla Dolomite (Johns, 1968) accumulated over erosional highs of Pandurra Formation during the Tapley Hill Formation transgression. The upper beds of the Tapley Hill Formation become increasingly calcareous and show evidence of higher energy conditions: cross-bedding, slumping and scouring are common near Adelaide. Stromatolites grew on higher erosional remnants, and intraclasts derived from the destruction of algal mats accumulated in erosional channels as edgewise breccias (Preiss, 1973). These changes reflect shallowing of water depth to above wave base and within the photic zone, and are most marked along the western margin of the Adelaide Geosyncline. In this zone, the Tapley Hill Formation is capped by the latest Sturtian regressive carbonates of the 5 0 m-thick Brighton Limestone (Howchin, 1904). The 140 m-thick Balcanoona Formation (Coats, 1971) is an analogous regressive carbonate in the NE Flinders Ranges. The lower member of the Brighton Limestone is typically crossbedded, massive, ooid and intraclast grainstone deposited under high-energy, intertidal to shallow subtidal conditions, with laterally intertonguing stromatolite bioherms in the Flinders Ranges. The upper member is buff, early diagenetic micritic dolomite with tepee structures and intraclast wackestones, formed on supratidal flats only in the marginal zones (Preiss, 1973; Preiss and Kinsman, 1978). The Brighton Limestone is absent from the eastern deeper water zones where the regression had less effect and deposition of laminated silts persisted into the early Marinoan. However, offshore ooid banks up to 3 0 0 m thick formed well into the basin in the latest Sturtian (Fig. 5 . 1 1 ~ )possibly delineating a shelf margin. Late Sturtian syndepositional tectonism is demonstrated by faults active during deposition (e.g., Plummer and Gostin, 1976) and by large-scale slump scours in the Tapley Hill Formation in the Willouran Ranges and north of Broken Hill. Lenses of coarse detritus adjacent to inferred diapiric structures have been interpreted as evidence of diapirism during sedimentation (Dalgarno and Johnson, 1968).
Marinoan Mawson and Sprigg (1950) defined the Marinoan as encompassing all deposition between the Sturtian and the base of the Cambrian. They placed the Sturtian-Marinoan boundary at the top of the Brighton Limestone in the type area near Adelaide. Subsequent mapping by the Geological Survey of South Australia showed that this lithologic boundary loses signifi-
341 cance in more basinal areas, and led to the definition of the Umberatana Group (late Sturtian-arly Marinoan) as a mapping unit. The gradational top of the Brighton Limestone is retained as the chronostratigraphic boundary. The marked polarity of sedimentary facies developed in the Adelaide Geosyncline during the late Sturtian persisted through most of the Marinoan, and is defined by more westerly shelf and more easterly basinal zones, as first demonstrated by Honvitz (1962). The shelf zone is characterized by redbeds and eastward-lensing sands, while the basinal areas accumulated drab silts, and during the Marinoan glaciation, tillites. Shelf carbonates are widespread in the early Marinoan of the central Flinders Ranges. Post-glacial transgression inundated the Stuart Shelf, but late Marinoan sedimentation may have been restricted to the area of the present Flinders Ranges (Fig. 5 -1Id). The late Sturtian regression left the western shelf areas under extremely shallow water. During the early Marinoan, an influx of clastics from the west was deposited in the shelf areas as the Willochra Subgroup (Thomson, 1969; Binks, 1971). The lower 300 m of this subgroup consists of ripple marked, cross-bedded siltstones and thin lenticular, feldspathic sandstones, with red, ripple-marked and mudcracked shale partings, indicating deposition under oxidizing conditions with frequent subaerial exposure on extensive tidal mudflats (Preiss, 1973; Plummer and Gostin, 1976). Recent drilling has established the westward extension of these redbeds on t o the Stuart Shelf (Yudnapinna Beds of Gersteling and Heape, 1975). In the northeast Flinders Ranges, the Angepena Formation (1300 m thick) represents analogous marginal sedimentation. The upper 400 m of the Willochra Subgroup in the western marginal zone is dominated by sand deposition but still under very shallow water conditions. In the more basinal areas of the northern and southeastern Flinders Ranges, up t o 900 m of laminated grey-green silts accumulated in subtidal environments (Amberoona Formation and Tarcowie Siltstone, respectively). Early Marinoan deposition in the Broken Hill area, NSW, was also of a basinal character. In the central Flinders Ranges, the shelf carbonates of the Etina Formation (Dalgarno and Johnson, 1966) were deposited during the latest Sturtian and early Marinoan. The base of this unit is diachronous, as it intertongues t o the southwest with the Willochra Subgroup and to the southeast with the Tarcowie Siltstone (Binks, 1971; Preiss, 1973). In the north, the lower Etina Formation partly merges with the Balcanoona Formation, and partly lenses out against the drab siltstones of the Amberoona Formation. The Etina Formation is characterized by cross-bedded sandy, gritty, oolitic and stromatolitic limestones interbedded with grey laminated siltstone, representing alternating high- and lowenergy conditions, in intertidal t o shallow subtidal zones, possible tidal channels and offshore carbonate banks. I t attains a maximum thickness of nearly 1300 m in the central Flinders Ranges, but here regional subsidence closely compensated for both sedimentation and the local rising of diapirs, some of which Dalgarno and Johnson (1968) demonstrated to have been exposed and shed detritus during deposition of the Etina Formation. The Enorama Shale (Dalgarno and Johnson, 1966), 500 m thick, is a greenish laminated siltstone and shale, lacking shallow-water features, that overlies Etina Formation, Amberoona Formation, Angepena Formation and Tarcowie Siltstone in the central Flinders Ranges and in basinal areas. It probably represents a transgressive phase, while the overlying 240 m of stromatolitic and intraclastic reddish limestone of the Trezona Formation (Dalgarno and Johnson, 1966) represent regression and a return to shelf carbonate deposition, chiefly in the central Flinders Ranges. The earlier differentiation of shelf and basinal environments persisted during the
342 Marinoan glaciation. In the western zones (including Stuart Shelf) as well as in the central Flinders Ranges, shallow-water sand deposition predominated, while glaciomarine tillites were deposited in the eastern and northern deeper basins. The 60 m-thick Elatina Formation (Mawson, 1949 ; Dalgarno and Johnson, 1966) is dominantly pinkish, massive, fine- to medium-grained, feldspathic sandstone, with pebbly lenses, deposited in the shelf environment: there is local evidence of an erosional base, and local developments of tillite with a red-brown matrix at the top. In the Adelaide region, the western shelf zone of the south Flinders Ranges and over part of the Stuart Shelf, the red, massive, granulebearing Reynella Siltstone Member, 120 m thick (Thomson, 1966b),resembles the matrix of the Elatina Formation tillites, but rare dropstones provide the only evidence of limited glacial influence (Coats, in prep.). The Yerelina Subgroup (Coats, 1971) records marine glacial deposition in the basinal areas. In the north Flinders Ranges, 1000 m of laminated siltstones with dropstones (Fortress Hill Formation of Coats, 1971), locally disconformable on the Amberoona Formation, record the first evidence of galciation, while in the southeast basin, the 75 m-thick Gumbowie Arkose (Mirams, 1964) also contains pebbles in interbedded shales (Coats, in prep.). The main phase of Marinoan glaciation is represented by 9 0 m of Mount Curtis Tillite of Coats, (1971) (northern basin) and up to 180 m of Pepuarta Tillite of Mirams, (1964) (southeast basin). Coats (in prep.) considers the overlying 2 0 0 m of feldspathic Balparana Sandstone of Coats (1971) (north) and Grampus Quartzite of Mirams (1964) (southeast) to be basinal equivalents of Elatina Formation sandstones, and the Elatina Formation tillites, and dropstones in the Reynella Siltstone Member, to record the last phases of Marinoan glaciation. Current direction measurements by Forbes (1975) on Yerelina Subgroup in the southeast basin imply a southeasterly palaeoslope, though not necessarily a northwesterly provenance for the glacial erratics. Indeed, the distribution of facies (Fig. 5.11d) has led to the suggestion of northerly and easterly source areas (Murrell, 1977). Glaciomarine tillites and associated quartzites, conglomerates and siltstones with dropstones were deposited in a similar basinal setting north of Broken Hill (Cooper and Tuckwell, 1971). Coats (in prep.) also recognizes possible Marinoan tillites in the eastern Officer Basin. Marinoan post-glacial sequences are referred to the Wilpena Group (Dalgarno and Johnson, 1964), and record at least two major transgressive and regressive cycles. The constituent formations of the Wilpena Group tend to be more uniform and laterally persistent than most earlier Adelaidean formations. The lower part of the Wilpena Group is widespread over the Stuart Shelf, but within the Adelaide Geosyncline it still displays the shelf and basin facies differentiation evident in the early Marinoan. The upper Wilpena Group transgressive-regressive cycle is preserved essentially only within the Flinders Ranges (e.g., distribution of Pound Subgroup in Fig. 5.11d). The base of the Wilpena Group is marked by a laterally persistent, though not always continuous unit, about 1 0 m thick, of pink flaggy, laminated dolomite and red shale (Nuccaleena Formation of Dalgarno and Johnson, 1964). Similar “cap” dolomites are useful stratigraphic markers overlying Marinoan glacials also in other Australian basins (Preiss et al., 1978; Coats, in prep.). In the Adelaide and Olary regions, the Nuccaleena Formation intertongues with elastics of the Seacliff Sandstone Member (Thomson, 1966b). The upper boundary of the Nuccaleena Formation is gradational into the silty shales of the lower Brachina Formation (Dalgarno and Johnson, 1964), recording a major shallow marine transgression over the whole Adelaide Geosyncline and Stuart Shelf. The Brachina Formation of the western and central Flinders Ranges shelf areas consists of about 1200 m of reddish-brown micaceous siltstones and fine sandstones, with greenish interbeds and locally with tuffaceous contribution (Plummer, 1978a), but passes laterally into green Ulupa Siltstone ( 1 5 0 0 m ) in the southeast and northern basinal areas. The upper Brachina Formation and its gradation into the overlying ABC Range Quartzite of Mawson (1939), with abundant mudcracks and clay pellets, represent regressive sedimen-
343 tation. Plummer (1978b) describes progressive shallowing from subtidal mud flats (widespread in lower Brachina Formation) to intertidal mud flats (middle and upper Brachina Formation). Geological Survey of South Australia mapping has shown the ABC Range Quartzite to thin and lens out to the east (Fig. 5.11d), suggesting a westerly derivation of clastics. Plummer (1978b) also presents evidence that it intertongues with the middle and upper Brachina Formation in the southwest Flinders Ranges where it attains a maximum thickness of 1800 m, and that it represents an intertidal deltaic plain environment. The sequence Woomera Shale-corraberra SandstoneArcoona Quartzite Members (Johns, 1968) of the Stuart Shelf is a Ethostratigraphic equivalent of the lower Wilpena Group, but the lithologic boundaries are likely to be diachronous in this eastward prograding sequence. The upper Wilpena Group is represented neither in the Adelaide region, nor on the Stuart Shelf (except for the Yarloo Shale) where Early Cambrian sediments transgressed disconformably over the lower Wilpena Group. In the Flinders Ranges, the upper Wilpena Group transgression is represented by 700 m of uniform and widespread red and greenish silty shales of the Bunyeroo Formation (Dalgarno and Johnson, 1964). Local erosional remnants of the equivalent Yarloo Shale (Johns, 1968) imply that the Stuart Shelf was also transgressed at this time, but the sediments were subsequently stripped before Early Cambrian deposition. The Bunyeroo Formation includes thin cupriferous dolomite bands and fine sandy dolomitic lenses with sole marks suggesting east-flowing currents (Forbes, in prep.). According to Coats (1973), several diapirs were active in the north Flinders Ranges. The overlying Wonoka Formation, 460 m thick (Dalgarno and Johnson, 1964) consists of greenish calcareous siltstones in the lower part and cross-bedded silty limestones with sole marks in the upper. Its base is usually gradational, though large-scale slump scours are locally present at this level throughout the Flinders Ranges (e.g., Coats, 1 9 6 4 ~ ) .Like the Bunyeroo Formation, the Wonoka Formation lacks shallow-water indicators, but the gradation into the overlying Pound Subgroup does record a major regressive phase. A thin oolitic and locally stromatolitic limestone commonly occurs at the top of the Wonoka Formation. In the extreme northeastern Flinders Ranges, Coats (1971) described green siltstones, sandstoiies and quartzite above the Wonoka Formation (Billy Springs Beds). Although these are 4800 m thick, they are restricted to that area. The Pound Quartzite (Mawson, 1938) was divided into a lower red, soft, silty, micaceous Bonney Sandstone Member and an upper more massive white Rawnsley Quartzite Member by Forbes (1971), but Jenkins (1975) proposed these as formations of the Pound Subgroup. Unlike preceding units, the Bonney Sandstone contains abundant shallow water features (mudcracks, clay pellets, lenticular bedding) indicating an oxidizing shallow marine to tidal flat environment. The Rawnsley Quartzite consists of clean washed, medium-grained sandstones deposited under higher energy conditions. Its lower part is probably transgressive (Jenkins, 1975), the flat- and cross-bedded sands being a shallow marine off-shore facies (Goldring and Curnow, 1967). The widespread Ediacara Member, occurring within the Rawnsley Quartzite, represents sheltered intertidal to lagoonal environments (Jenkins, 1975), and contains the well-known Ediacara assemblage of soft-bodied metazoan fossils (Sprigg, 1947; Glaessner and Wade, 1966). A seaward barrier for such a lagoonal system may have been provided by diapirs rising in late Marinoan time (Wade, 1968) o r by long arcuate off-shore sand bars (Jenkins, 1975). Diapiric activity during Pound Subgroup deposition is suggested by Leeson (1970). The Pound Subgroup is restricted to the Flinders Ranges, where i t is the main ridgeforming unit in the present landscape. I t attains a maximum thickness of almost 3000 m in the northeastern Flinders Ranges. The Punkerri Beds are a probable equivalent in the Officer Basin (Major, 1974). The end of the Marinoan is marked by very extensive regression of the sea and substantial erosion of Marinoan sediments, especially on the Stuart Shelf. Marine transgression
344 during the Early Cambrian reinitiated sedimentation within both the Adelaide Geosyncline and Stuart Shelf.
Summary of history of sedimentation The Adelaide Geosyncline is a complex system of basins that received very thick accumulations of sediment in the Adelaidean and Cambrian. The stratigraphic history can be divided into a number of phases. The nature and loci of deposition changed substantially from one phase t o another. The following major events and relationships are recognized in the Adelaide Geosyncline and adjacent Stuart Shelf. Pre-Adelaideanplatform history (c. 1400-1000 Ma): (1)Basic volcanism and associated sedimentation. (2) Continental Pandurra Formation deposition (possibly early Willouran). Will0uran : (1)Deposition of transgressive clasticcarbonate blanket (known only within Geosyncline). (2) Rifting and basic vulcanism (known in both Stuart Shelf and Geosyncline). (3)Deposition of clastics, carbonates and evaporites in restricted basins (known only in Geosyncline). Torrensian: Not known on Stuart Shelf, and perhaps never deposited there. (1)Transgressive fluvialshallow-marine clastics. (2) First prograding deltaic cycle. (3)Shelf carbonate deposition with marginal sands, basinal black shales. (4)Second prograding deltaic cycle. (5) Third and fourth prograding deltaic cycles. Sturtian: (1)Final prograding deltaic cycle. (2) Fault movements and differential uplift producing widespread unconformity. (3)First Sturtian glacial phase. Thick glacial sediments rest on basement in north and northeast. Areas of greatest pre-Sturtian uplift also suffered greatest Sturtian subsidence. (4)Continued fault movements and differential uplift. (5) Second Sturtian glacial phase. (6) Post-glacial marine transgression, extending on to Stuart Shelf. (7) Regression in labile shelf areas; off-shore ooid banks; syndepositional tectonism, slumping. Marinoan: Labile shelf and basin facies well differentiated in the Geosyncline. Only (1)t o (4)known on Stuart Shelf. (1)Tidal flat redbeds on shelf, drab silts in basin. Platform carbonates in intracratonic area. (2) Sand deposition with local glacial influence in shelf zone; glaciomarine deposition in basins.
345 (3) Post-glacial marine transgression.
(4)Regression and deltaic progradation. (5) Transgression; diapirism; syndepositional slumping. (6) Deposition of a regressive sand blanket. (7) Cessation of Adelaidean sedimentation; erosion prior t o Cambrian transgression.
Tectono-stratigraphic discussion The above summary indicates that tensional tectonism was characteristic of the evolution of the Stuart Shelf and Adelaide Geosyncline from preAdelaidean into Sturtian time. Basic volcanism occurred intermittently on the Shelf from about 1300 Ma and is also represented by Willouran volcanics within the geosyncline. This suggests a phase of extensional tectonism in which the Shelf formed part of an uplifted region and the Geosyncline was already developing adjacent subsidiary basins. Graben development, especially on the western margin of the Geosyncline, was also characteristic of the succeeding Burra Group, which is essentially restricted t o the Geosyncline. Vertical fault movements were of particularly large amplitude and widely distributed in the early Sturtian. The most notable activity was then in the north and east of the Geosyncline, as in the Mt. Painter and Broken Hill areas. Similar strong relief was never developed on the west of the Geosyncline in the area of the Stuart Shelf. In spite of some recurrences, there was a diminution of syndepositional faulting and of tectonic relief from the time of deposition of the Sturtian Tapley Hill Formation. Regional patterns of sedimentation became simpler and more uniform and a record of major transgressions is preserved on the Stuart Shelf. The Tapley Hill Formation itself indicates a deepening of the Geosyncline and a transgression onto the Shelf which was extended (with intermittent regressions) during the Marinoan and Cambrian. Thus it is possible t o distinguish between a protracted period of strong intermittent tensional tectonism up t o early Sturtian time and a relatively quiescent period during the remainder of Adelaidean time. In view of the presumed age of the Sturtian glacials (c. 725 f 30 Ma; Compston and Arriens, 1968) a lower limit of about 700 Ma can be set to the protracted extensional phase, which apparently began more than 600 Ma earlier. It may be significant that compressive tectonics both in the Amadeus Transverse Zone (see Chapter 1; the Petermann Ranges Orogeny) and in Tasmania (the Frenchman and Penguin Orogenies) began at about the same time as the extensional phase waned in the Adelaide Geosyncline. The stratigraphic record indicates a further change in the tectonic regime during the Cambrian (e.g., Daily, 1976). The regression indicated by the evaporites and red beds of both the late Lower Cambrian Billy Creek Formation (with its associated ash fall tuffs) and the Middle Cambrian Lake Frome
346 Group correlates with the development of the Early Cambrian Kanmantoo Group Trough with its rapid deposition of deeper water clastics. It also correlates with the development of volcanic arcs on the margins of back-arc basins further east (e.g., the Mt. Wright volcanic arc; Scheibner 1973) and ushers in the folding of the Delamerian Orogeny. At this time Wopfner (1972) postulates the development of a pericratonic circum-Denison Arc, “a major depositional locus which extended from the Bancannia Trough in northwestern New South Wales to the Officer Basin”. The unifying factor during this tectono-stratigraphic evolution is the Torrens Hinge Zone which was always a significant zone between Geosyncline and Shelf. The N-S trend of the hinge zone (and of Late Cenozoic structures) has led to an oversimplified view of the Geosyncline as a N-S trough. Figure 5.11 shows that the basin configuration has changed markedly during Adelaidean times. During the protracted extensional period up to early Sturtian time there is repeated evidence of uplifts by differential faulting, providing sediment source areas to the north and north east of the Geosynclinal area, and the inferred Curnamona Cratonic Nucleus may have been emergent for much of this time. The distribution of early Marinoan tillites also raises the possibility of northerly and south easterly source areas at that time (Murrell, 1977). However, these eastern source areas were of more mobile (taphrogenic) character than the Stuart Shelf and the Geosyncline was essentially asymmetric. Moreover, two main basinal areas evolved, corresponding t o the North Flinders and Mt. Lofty-Olary fold arcs. The intervening area of the South and Central Flinders zones had a more stable labile shelf character linking the Stuart Shelf t o the Curnamona Cratonic Nucleus. It is possible to envisage that in later Adelaidean times the two main basinal areas opened northwards and south-eastwards respectively towards deep ocean across a nearby continental margin. However, the evidence of the Muloorina Ridge and of source areas during earlier Adelaidean deposition suggests the alternative possibility that subsided continental basement areas may lie in these directions. In any case, it is clear that the Adelaide Geosyncline as a whole was never a simple N-S fault trough with stable craton on both sides. Rather the region east of the Torrens Hinge Zone has been disrupted by a mosaic of faults controlled by pre-existing basement structures. These structures have controlled the development of the two main basinal areas and the subsequent folding deformation. It certainly seems likely that the whole region east of the Torrens Hinge Zone formed part of the earlier Proterozoic craton. However, this region apparently lost its cratonic character early in the Adelaidean, except for the remnant Curnamona nucleus. Thus only the Central Flinders zone can reasonably be described as intracratonic during the depositional history. Since it links complex basinal areas both t o north and south it can scarcely be regarded as an aulacogen, in the currently popular restricted sense as the failed arm of a plume generated triple junction
347 (Burke and Dewey, 1973). It evidently has not been developed as a fault trough propagated from a triple junction either t o north or south. The aulacogen hypothesis is perhaps more relevant on a much larger scale t o the Amadeus Basin. Its development as a broad E-W aulacogen could have occurred concomitantly with the inferred fragmentation of the continental margin during Adelaidean deposition. There is, however, little evidence that deposition in the Amadeus Basin was controlled by major normal faults and the initial blanket quartzites apparently extended across the basement complexes both t o north and south. The present margins of the basin are controlled by late Precambrian and Palaeozoic reactivation of the major thrust faults which developed during the earlier Musgravian orogeny, and which cut across thick Adelaidean and Lower Palaeozoic accumulations (Wells et al., 1970). The temporal and spatial relationships in evolution of the Adelaidean Geosyncline suggest that analogy is better made with the evolution of multiple rifted arch systems in the development of Phanerozoic passive continental margins (Veevers and Cotterill, 1978). The protracted phase of tensional tectonics up t o early Sturtian times could correspond t o the slow development of a multiple rifted arch system. Late Sturtian and Marinoan deposition could then correspond t o the development of a broad inter-arch and subsequent rim basin. Finally the Lower Cambrian carbonates could correspond to the development of a prograding carbonate wedge. In this model, the Torrens Hinge Zone would form the western margin of the rifted arch system which might be expected to extend eastwards for a distance of the order of 1000 km. The Curnamona Nucleus could be regarded as a stable part of the inter-arch basin and the Broken Hill region would form part of a more easterly rift system than the troughs adjacent t o the Torrens Hinge Zone. Recent correlations between Australia and Antarctica give some support t o this model. Laird et al. (1977) recognize the possibility of Cambrian sedimentation occurring in two distinct belts. A westerly belt, represented by the Adelaide Geosyncline and the Byrd Group in Antarctica, displays Early t o Middle Cambrian archaeocyathid limestone sequences resting with varying degrees of unconformity on late Precambrian sediments. Early Ordovician granitic rocks occur in both regions and suggest the correlation of the Delamerian and Ross orogenies. An easterly belt is represented by the Bowes Group in Antarctica and the Dundas Group in Tasmania. Unfossiliferous clastic sequences with basic volcanics are overlain by clastic Middle t o Late Cambrian sequences with shelly fossils. The Dundas trough also has ultramafic rocks and has been commonly interpreted as a narrow marginal basin floored by oceanic crust. “The youngest fossiliferous marine deposits are in both cases overlain unconformably by dominantly continental coarse quartzose sandstones and conglomerates of probably latest Cambrian or earliest Ordovician age” (Laird et al., 1977). No important folding precedes this unconformity and early Ordovician granites are
348 absent so that the Ross or Delamerian orogenies are therefore of much lesser significance in the eastern belt. In contrast late Precambrian orogenies are observed (Beardmore, Frenchman, Penguin orogenies). The Cambro-Ordovician sequences have also been folded in a later, possibly Devonian, orogeny and both areas are intruded by Devonian-Carboniferous granites. Thus the two belts could be parts of a multiple rifted arch system, the easterly belt being closer to the developing continental margin. The concept of two belts of differing character is somewhat reinforced by the similarity of the area east of Broken Hill t o the eastern belt. The Mt. Wright Volcanics in the lower part of the sequence on the eastern side of the Bancannia Trough, recall the Mt. Read Volcanics on the eastern side of the Dundas Trough (Scheibner, 1973). Fossiliferous Middle Cambrian sediments with tuffaceous components are also present, while a pebbly quartzite again begins the Ordovician sequence which is unconformable on probable Lower Cambrian rocks (Wopfner, 1966; Webby, 1978). Again the Cambrian and Ordovician rocks have been folded together, in an orogeny for which a Late Ordovician to Silurian age is suggested (Wopfner, 1966). Possibly the eastern belt can be recognized again in Queensland, where the Windsor Volcanics form part of the Cambro-Ordovician sequence of the Lolworth-Ravenswood Block (Murray and Kirkegaard, 1978). Thus it seems possible to envisage the Adelaide Geosyncline as part of a major western belt of rifting, broadly parallel to, and several hundred kilometres from, an eastern rift belt. The latter may itself have been well t o the west of the main continental margin which developed as a result of the rifting. If the Adelaide Geosyncline is thus related to the development of a continental margin as part of a multiple rifted arch system it can be postulated that the continental margin developed prior to 700 Ma ago. Direct evidence of the position of the continental margin at that time is lacking since the oldest rocks in the Lachlan Geosyncline are of Cambrian age. It is possible that this Geosyncline formed part of a marginal basin floored by oceanic crust from the Cambrian (e.g., Packham and Falvey, 1971). Alternatively it is possible that thin older Proterozoic crust extends under much of the Lachlan belt (e.g., Rutland, 1973a, b). Evidence of an active continental margin, probably of Andean type is found on the eastern edge of the Lachlan belt. A volcanic arc is first recognized in the brdcvician and it occupied much the same position throughout the Palaeozoic (e.g., Rutland, 1976). Some evidence for subduction-related volcanism has been presented for the Stavely Greenstone Belt on the west of the Lachlan Fold belt (Crawford and Keays, 1978). Nevertheless compressional tectonics represented by the late Precambrian Petermann Ranges, Frenchman and Penguin Orogenies, as well as by the Cambro-Ordovician Delamerian Orogeny, suggest that an active continental margin was already developed in the Late Proterozoic. The different timing of deformation in the eastern and western rift belts presents an intriguing
349 problem in a plate tectonics context. The orogenies may represent phases of compression alternating with the later phases of tension in the rifts of the original multiple rifted arch system. Some of these rifts may have evolved further as back-arc basins after the development of an active continental margin, and may be floored by oceanic crust. The Delamerian Orogeny Timing of deformation The principal fold structures of the fold belt are attributed t o the Delamerian Orogeny, when major faults were also reactivated, especially along the western margin. Since Cambrian rocks are involved in the fold structures and since Early Ordovician granites occur sporadically the orogeny is usually placed at about the Cambro-Ordovician boundary. However, as already noted, the main folding on the east of the Bancannia Trough was post-Lower Ordovician and there is some stratigraphic evidence for varying age of folding elsewhere. According to Wopfner (1972) deposition in the extensions of the Adelaide Geosyncline adjacent to the Musgrave and Denison blocks was interrupted by fold movements towards the end of Precambrian time. In the Fleurieu arc itself there is also some evidence of folding (the Duttonian Folding of Thomson, 1969) prior t o the deposition of the Early Cambrian sediments and basic volcanics which rest unconformably on the Umberatana Group. The relations of the Cambrian Kanmantoo Group t o the Houghton anticlinal zone may have been in existence and subject to erosion before the Kanmantoo sediments were deposited across it. Nevertheless, the main folding and metamorphism in this region is clearly post-Cambrian, and indeed part of the Cambrian Kanmantoo Group shows the strongest deformation and the highest grade of metamorphism (e.g., Offler and Fleming, 1968). As indicated in the preceding section the tectonics prior t o 700 Ma was principally extensional. Nevertheless there is a widespread early-Sturtian unconformity and there is some evidence of folding in the Willouran Ranges prior to unconformable deposition of the Burra Group. Murrell (1977) suggests that the latter may be essentially due t o soft-sediment sliding, diapirism and growth structures formed during syndepositional basin development in Willouran times, mobility being increased by the evaporites intercalated in the sequence. He attributes the brecciation of these beds to this phase of mobility rather than to the Delamerian Orogeny. Mount (1975), on the other hand, (who studied the Thomson’s Gap and Arkaba “diapirs” in particular) , while recognizing the extensional nature of the breccias, attributes both breccias and “diapirs” t o permitted intrusion during the Delamerian orogeny rather than t o forceful diapiric intrusion as envisaged by most previous authors. In any event there is abundant evidence that the “diapirs” are related t o the fault pattern (e.g., Thomson, 1970) and that they were source
350 areas of debris shed into the geosyncline at various times in the stratigraphic history. Whether the “diapirs” were then masses of ductile flowing breccia, or upfaulted blocks of already consolidated breccia, or upfaulted blocks of as yet unbrecciated material, is still open t o dispute (Coats, 1964a; Dalgarno and Johnson, 1968; Haslett, 1975). Crucial questions are whether the breccias contain “xenoliths” of rocks derived from higher, stratigraphic levels, and whether the surrounding sediments, derived from the “diapirs”, contain clasts of breccia, but both these points are arguable at present. Whatever the interpretation of the “diapirs” it is likely that faulting and vertical movements, rather than compressional deformation, were responsible for the unconformities in the Adelaidean sequence. Compressional deformation varies in age from latest Precambrian to post-Lower Ordovician. Style of deformation The general similarity in the size and shape of the Mt. Lofty-Olary fold belt to the Jura, and the possibility of decollement in the Callanna Beds, have both been noted above. However, the Mt. Lofty-Olary fold belt displays thicker, more variable and more basinal sedimentary facies than those of the Jura. Moreover, a slaty cleavage is commonly developed (see e.g., Bell, 1978), and organic, mineralogical and magnetic indicators reveal significant metamorphism in the eastern part of the belt (McKirdy et al., 1975). There is also abundant evidence for basement control of the folding in the Delamerian fold belt. The basement rocks of the fold belt show rather variable trends of their macroscopic fold structures. These fold structures have been subjected to concordant pre-Adelaidean retrogression and in addition they are cut by preAdelaidean retrograde schist zones of varying trends. Both the syndepositional faulting and the subsequent folding of the Adelaidean rocks can be closely correlated with these basement trends. The clearest evidence is in the Broken Hill region (e.g., Glen et al., 1977). North of Broken Hill, Adelaidean rocks are faulted against basement rocks on NW-trending faults which follow shear zones in the basement. The shear zones were initiated prior to Adelaidean deposition and were also active during deposition especially in the early Sturtian (Cooper and Tuckwell, 1971). They are oblique t o the N-S trend of the basement fold structures in the same vicinity. The Adelaidean rocks show a fold interference pattern between NW- and N-trending folds which is clearly due t o the interaction of the two basement trends. South and west of Broken Hill the basement fold trends are WSW and the macroscopic folds in Adelaidean rocks have the same trend, again with local interference by NW-trending basement faults. In the Adelaide region, both basement and Adelaidean cover folds have northerly trends. Basement fold trends in the Gawler Platform, which was undeformed during the Delamerian Orogeny, show similar variation to those within the Geosyncline and suggest that the basement trends were already established
351 prior to Adelaidean deposition (Glen et al., 1977). Consequently, it can be inferred that the arcuate fold trends in the Delamerian fold belt are essentially inherited from the earlier basement trends, and are not due t o oroclinal bending of originally linear fold systems. This conclusion is supported by the stratigraphic evidence which shows that the two principal fold arcs, in the North Flinders zone, and in the Mt. Lofty-Olary fold belt, are both broadly parallel to the boundaries between labile shelf and more basinal facies. The shapes of the depositional basins were themselves controlled by basement fault movements and the basement trends have been utilised during the subsequent compression of the basins. Coward (1976) supposed for the Mt. Lofty-Olary system that an original linear fold trend was ENEWSW and that the present near N-S segment has been related by a process of sinistral simple shear on a broad deep seated shear zone trending NNW. This is indeed a prominent trend of basement shear zones but it can hardly have been active as postulated by Coward. The proposed shear zone has no effect on the approximately meridional Torrens Hinge Zone; and the Adelaidean sediments on the Stuart Shelf which lie on the NNW continuation of such a shear, show only the gentlest of folding. Coward (1976) supported his argument with an isogon map purporting t o show that the pattern of axial traces has the form of a ductile shear. This conclusion, however, appears t o be derived from an oversimplification of the macroscopic axial traces especially in the apical zone of the Nackara arc. In fact the enBchelon style there, with loss of continuity of fold axial tracw in the apical area, is further evidence that the curvature of the Nackara arc is a primary feature. The folds in the N-S limb of the Nackara arc are tighter than those in the apical zone or in the ENE trending limb but the strain does not increase into the apical zone of the Fleurieu arc which is concave towards the craton. On the contrary, the strain reaches a maximum there and is characterized by overturned folds and thrusts with NW vergence (e.g., Offler and Fleming, 1968). Penetrative structures plunge t o the south or southwest and do not change as the regional strike swings into the Fleurieu arc. In the E-W limb of the Fleurieu arc the stratigraphic history and style of deformation is quite different from elsewhere in the fold belt. Thus it can be concluded that neither the Nackara nor Fleurieu arcs is due to oroclinal bending but that both are essentially primary. The folding can be attributed to shortening in both basement and cover along variable trends controlled by basement structures and the subsequent Adelaidean deposition. Paleomagnetic evidence also opposes the orocline hypothesis (Giddings and Embleton, 1974). Compression in a NW-SE direction most readily accounts for the structures in the Mt. Lofty-Olary fold belt. It is evidently quite inappropriate, however, t o the deformation in the Central and Northern Flinders zones which, along with the South Flinders Zone, display the principal occurrences
352
Fig. 5.12. General geology of part of the Flinders Ranges, showing relations between the North, Central and South Flinders zones. Based o n maps of the Geological Survey of South Australia.
I-.
EAST
Fig. 5.13. Geological section across the Torrens Hinge Zone and North Flinders zone, northeastwards from the north end of Lake Torrens. Modified after Murrell, 1977.
353
354 of diapiric structures (Fig. 5.12). Moreover, the Central Flinders Zone is much less strongly folded than the arcs to north and south. The principal fold structures of the Central Flinders zone are a broad anticlinesyncline pair trending NNW approximately parallel t o the Norwest Fault. Analogy with the Houghton zone, in the Mt. Lofty-Olary arc, suggests that the broad anticline is underlain by a broad basement culmination following the dominant basement trend. Alternatively as implied by crosssections accompanying Geological Survey Maps, the broad anticline may be underlain by thick diapiric material and the basement may remain at great depth. A set of NE- to ENE-trending faults cuts across these broad folds, and localises “diapiric” bodies principally in anticlinal culminations but also locally in the synclinal zone. This pattern is consistent with compression in an ENE direction t o produce the broad folds, and with tensional faults in the same direction. This contrasts with the pattern in the Mt. Lofty-Olary belt already noted and with the pattern in the North Flinders zone. The northeasterly fault trend continues locally into the east of the North Flinders zone where it corresponds with basement trends in the Mt. Painter basement complex. The broad fold trend of the Central zone on the other hand continues into the west of the North Flinders zone where it forms a broad belt parallel to the Norwest Fault and, implies a different basement trend in that region (which also controlled the syndepositional faulting). In addition the North Flinders zone displays a series of unusually tight synclines which trend approximately E-W. In the Mt. Painter region there is also a pronounced cleavage in the same direction, oblique t o the main basement trend. It appears, therefore, that the North Flinders arc is a composite structure. It is most simply explained by a N-S compression t o produce the tight synclines, combined with control of broad folds by both NE- and NWtrending basement structures. The different directions of compression indicated for the different zones in the fold belt are remarkable. Cambrian rocks are involved in all areas so that there is no stratigraphic evidence for ascribing different ages to the different trends. It appears that the local stress directions may have been the consequence of jostling between different basement blocks. The compression in the Mt. Lofty-Olary arc may be related to that along the main continental margin, while that in the North Flinders zone may be related t o that in the Amadeus Transverse Zone. The deformation may also be seen as peripheral to the Gawler Craton so that the fold trends swing to the west at both the northern and southern ends of the Torrens Hinge Zone. The folding deformation rapidly dies out across the hinge zone as shown in the sections of Fig. 5.13. REFERENCES Ambrose, G.J. and Flint, R.B., 1979. Precambrian and Palaeozoic geology of the Peake and Denison Ranges. South Aust. Dep. Mines and Energy Rep. 79/19 (unpubl.). Bell, T.H., 1978. The development of slaty cleavage across the Nackara Arc of the
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359 Preiss, W.V., 1974. The River Broughton Beds - a Willouran sequence in the Spalding Inlier. South Aust. Geol. Surv. Q. Geol. Notes, 49: 2-8. Preiss, W.V., 1977. The biostratigraphic potential of Precambrian stromatolites. Precambrian Res., 5 : 207-219. Preiss, W.V., 1978. Progress report o n t h e geology of part of the Worumba Anticline. South Aust. Dep. Mines and Energy Rep. 7816 (unpubl.). Preiss, W.V. and Kinsman, J.E., 1978. Stratigraphy and palaeoenvironmental interpretation of the Brighton Limestone south of Adelaide and its equivalents in the Orroroo region. South Aust. Geol. Surv. Rep. Invest., 49: 34 pp. Preiss, W.V., Walter, M.R., Coats, R.P. and Wells, A.T., 1978. Lithological correlations of Adelaidean glaciogenic rocks in parts of the Amadeus, Ngalia, and Georgina Basins. BMR J. Aust. Geol. Geophys., 3 : 43-53. Rutland, R.W.R., 1973a. Tectonic evolution of t h e continental crust of Australia. In: D.H. Tarling and S.H. Runcorn (Editors), Continental Drift, Sea Floor Spreading and Plate Tectonics: Implications t o the Earth Sciences. Academic Press, London, pp. 1003-1025. Rutland, R.W.R., 1973b. On the interpretation of cordilleran orogenic belts. Am. J . Sci., 273: 811-849. Rutland, R.W.R., 1976. Orogenic evolution of Australia. Earth-Sci. Rev., 12: 161-196. Rutland, R.W.R., Marjoribanks, R.W., Laing, W.P. and Glen, R.A., 1978. Tectonic deformations at Broken Hill, New South Wales, and their significance for interpretations of the ore environment. Inst. Min. Metall. Trans. Sect. B, 8 7 : B 172-176. Scheibner, E., 1973. A plate tectonic model of t h e Palaeozoic tectonic history of New South Wales. J. Geol. SOC.Aust., 20: 405-426. Schopf, J.W. and Barghoorn, E.S., 1969. Microorganisms from the late Precambrian of South Australia. J. Paleontol., 4 3 : 111-1 18. Schopf, J.W. and Fairchild, T.R., 1973. Late Precambrian microfossils: a new stromatolitic biota from Boorthanna, South Australia. Nature, 242: 537-538. Shaw, S.E., 1968. Rb-Sr isotopic studies of the mine sequence rocks at Broken Hill. In: M. Radmanovich and J.T. Woodcock (Editors), Broken Hill Mines - 1968. Australas. Inst. Min. Metall. Monogr., 3 : 185-198. Sprigg, R.C., 1946. Reconnaissance geological survey of portion of the western escarpment of the Mount Lofty Ranges. Trans. R. SOC.South Aust., 70: 313-347. Sprigg, R.C., 1947. Early Cambrian (?) jellyfishes from the Flinders Ranges, South Australia. Trans. R. SOC.South Aust., 71: 212-224. Sprigg, R.C., 1949. Thrust structures of the Witchelina area, South Australia. Trans. R. SOC.South Aust., 73: 40-47. Sprigg, R.C., 1952. Sedimentation in the Adelaide Geosyncline and the formation of the continental terrace. In: M.F. Glaessner and R.C. Sprigg (Editors), Sir Douglas Mawson Anniversary Volume, University of Adelaide, pp. 153-159. Stanton, R.L., 1976. Petrochemical studies of ore environment at Broken Hill, New South Wales, Australia: 4 - environmental synthesis. Inst. Min. Metall. Trans. Sect. B, 85: B221-233. Stewart, J.H., 1972. Initial deposits in the Cordilleran geosyncline: evidence of a Late Precambrian (<850 m.y.) continental separation. Geol. SOC. Am. Bull., 83: 13451360. Talbot, J.L., 1963. Retrograde metamorphism of the Houghton Complex, South Australia. Trans. R. SOC.South Aust., 87: 185-196. Thomson, B.P., 1966a. The lower boundary of the Adelaide System and older basement relationships in South Australia. J. Geol. SOC.Aust., 13: 203-228. Thomson, B.P., 1966b. Stratigraphic relationships between sediments of Marinoan age - Adelaide region. South Aust. Geol. Surv. Q. Geol. Notes, 20: 7-9.
360 Thomson, B.P., 1969. Precambrian basement cover - the Adelaide System. In: L.W. Parkin (Editor), Handbook of South Australian Geology. South Aust. Geol. Surv., Gov. Printer, Adelaide, pp. 49-83. Thomson, B.P., 1970. A review of t h e Precambrian and lower Palaeozoic tectonics of South Australia. Trans. R. SOC.South Aust., 94: 193-221. Thomson, B.P., 1976. Regional geology of the Gawler Craton. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5: 461-466. Thomson, B.P., 1980. Geological Map of South Australia, Scale 1 : 1,000,000. Geol. Surv. South Aust. Thomson B.P. and Coats, R.P., 1964. Pre-Burra Group sequences in the Adelaide Geosyncline. The Callanna Beds (new name). South Aust. Geol. Surv. Q. Geol. Notes, 9: 3-5. Thomson, B.P., Coats, R.P., Mirams, R.C., Forbes, B.G., Dalgarno, C.R. and Johnson, J.E., 1964. Precambrian rock groups in the Adelaide Geosyncline: a new subdivision. South Aust. Geol. Surv. Q. Geol. Notes, 9 : 1-19. Thomson, B.P., Daily, B., Coats, R.P. and Forbes, B.G., 1976. Late Precambrian and Cambrian geology of the Adelaide “Geosyncline” and Stuart Shelf, South Australia. 25th Int. Geol. Congr., Excursion Guide 33a, 65 pp. Tilley, C.E., 1921. The granite gneisses of southern Eyre Peninsula and their associated amphibolites. Q. J. Geol. SOC.London, 77: 75-134. Tuckwell, K.D., 1978. Stratigraphic subdivision and correlation of the upper part of the Willyama Complex, New South Wales. J. Geol. SOC.Aust., 25(5): 295-308. Uppill, R.K., 1979. Stratigraphy and depositional environments of the Mundallio Subgroup (new name) in the late Precambrian Burra Group of the Mount Lofty and Flinders Ranges. Trans. R. SOC.South Aust., 103: 25-43. Veevers, J.J. and McElhinny, M.W., 1976. The separation of Australia from other continents. Earth-Sci. Rev., 1 2 : 139-159. Veevers, J.J. and Cotterill, D., 1978. Western margin of Australia: Evolution of a rifted arch system. Geol. SOC.Am. Bull., 89: 337-355. Vernon, R.H., 1961. Banded albite rich rocks of the Broken Hill district, New South Wales. Mineragr. Invest. Tech. Pap. C.S.I.R.O. Aust., 3. Wade, M.J., 1968. Preservation of soft-bodied animals in Precambrian sandstones at Ediacara, South Australia. Lethaia, 1: 238-267. Webb, B.P., 1960. Diapiric structures in the Flinders Ranges, South Australia. Aust. J. Sci., 22: 9. Webb, A.W. and Horr, G., 1978. The Rb-Sr age and petrology of a flow from the Beda Volcanics. South Aust. Geol. Surv. Q. Geol. Notes, 66: 10-13. Webb, A.W. and Thomson, B.P., 1977. Archaean basement rocks in the Gawler Craton, South Australia. Search, 8 : 34-36. Webby, B.D., 1978. History of t h e Ordovician Continental Platform Shelf Margin of Australia. J. Geol. SOC. Aust., 25: 4 1 4 3 . Wells, A.T., Forman, D.J., Ranford, L.C. and Cook, P.J., 1970. Geology of the Amadeus Basin, Central Australia. Bur. Miner. Resour. Aust. Bull., 10: 222 pp. Whitten, G.F., 1966. Suggested correlation of iron ore deposits within South Australia. South Aust. Geol. Sum. Q. Geol. Notes, 18: 7-11. Whitten, G.F., 1970. The investigation and exploitation of the Razorback Ridge iron deposit. South Aust. Geol. Surv. Rep. Invest., 33: 165 pp. Williams, E., 1978. Tasman Fold Belt System in Tasmania. Tectonophysics, 48: 159-205. Wilson, A.F., 1952. The Adelaide System as developed in the Riverton-Clare region, northern Mount Lofty Ranges, South Australia. Trans. R. SOC.South Aust., 75: 131-149. Wopfner, H., 1966. Cambro-Ordovician Sediments from the North-Eastern Margin of the Frome Embayment (Mt. Arrowsmith, N.S.W.) J. Proc. R. SOC.N.S.W., 100: 163-177. Wopfner, H., 1972. Depositional History and Tectonics of South Australian Sedimentary Basins. Miner. Resour. Rev., South Aust., 133: 32-50.
361
Chapter 6
AUSTRALIAN PRECAMBRIAN PALAEOBIOLOGY M.F. GLAESSNER and M.R. WALTER
INTRODUCTION
Precambrian fossils have been known from Australia since 1914, when the stromatolites of the Bitter Springs Formation were first described (Chewings, 1914; Howchin, 1914). These and other stromatolites later described by Mawson and Madigan (see Walter, 1972) were the only Australian Precambrian fossils that were verified by subsequent studies, until the now famous, latest Precambrian Ediacara fauna of soft-bodied metazoans was discovered in 1947 (at that time it was conventionally included in the Cambrian; Sprigg, 1947). The Ediacara fauna is now recognized as one of the cardinal faunas of the fossil record (Glaessner, 1978). The next phase in the study of Australian Precambrian fossils began with the discovery (Barghoorn and Schopf, 1965) and description (Schopf, 1968) of the microscopic bacteria, cyanophytes and algae of the 8004300 Ma-old Bitter Springs Formation. Resulting from the diversity, excellent preservation and detailed description of this microbiota, these fossils have played a signal role in discussions of the course of evolution through the Precambrian. Concurrently with the work on microfossils, comprehensive studies of stromatolites commenced anew, initially with the aim of testing their biostratigraphic utility (Glaessner et al., 1969; Cloud and Semikhatov, 1969; Walter, 1972; Preiss, 1972; etc.), but more recently also to exploit their palaeobiologic and palaeo-environmental record. Intermittent study of trace fossils has now been given new impetus by their potential significance for the positioning of the PrecambrianCambrian boundary (Glaessner, 1969; Daily, 1972; Cowie and Glaessner, 1975). The rich and varied Precambrian fossil record of Australia is now being elaborated at an increasing rate. During the last 1 0 years the number of known microfossiliferous formations has increased from 1 t o 16. Stromatolites are almost ubiquitous, microfossils and trace fossils are widespread, and new finds of fossil soft-bodied invertebrates are being made. The best-known biotas, the Ediacara fauna, Bitter Springs microbiota, and
362 the stromatolites of the Adelaide Geosyncline and Amadeus Basin will not be described at length. The more recently discovered fossils, which have yet to be intensively analyzed, will be considered in more detail. Particular attention will be directed to distinguishing the major groups of microfossils. In the fossil record not only the distinction between phyla of animals and plants but also the most fundamental one between the anucleate, asexual bacteria and cyanophytes (“blue-green algae”), the procaryotes, and the nucleated algae and all higher organisms, the eucaryotes, most of which reproduce sexually, is still obscure. The evolution of the eucaryotes from the procaryotes through a series of radical innovations marked the inception of evolutionary processes as they are conventionally understood. The regular processes of sexual reproduction (mitosis and meiosis) occur only in the eucaryotes. In comparison, the reproductive processes of procaryotes result in vastly less genetic variability, and thus much slower speciation rates and diversification (Schopf et al., 1973). The study of Australian microbiotas plays an essential part in the continuing effort t o record the most significant evolutionary step which has led to the diversity of life as we know it. The study of Precambrian fossils must now include not only the cataloguing and describing of occurrences, esential though that is, but also interpretation and analysis of their evolutionary and environmental context. This is true of stromatolites as well as other fossils. Stromatolites are not “body fossils” but the traces of life activity. They are sedimentary structures built by bacteria, cyanophytes and microscopic algae, and only rarely do they contain remains of the organisms that formed them. Yet the biologically and environmentally informative different shapes and fabrics of microbial colonies are recorded in the shapes and fabrics of the stromatolites (Walter, 1977). The temporal framework for this review is that of Dunn et al. (1966), with some changes t o take account of new data. The most important change is that the base of the Adelaidean is probably much younger than the 1400 Ma originally proposed (Cooper, 1975) and so the Late Precambrian is divided for the purpose of this study at about 950 Ma (the base of the Upper Riphean defined on stromatolites) and about 750 Ma (the approximate age of the lower glacigenic sequences). These subdivisions are used because they are convenient for the following discussion. ARCHAEAN (>2500 Ma)
Black cherts are abundant in the greenstone belts of Western Australia but few have been palaeontologically examined. An assemblage of carbonaceous spheroids interpreted as microfossils has recently been discovered in cherts in barite deposits at North Pole in the Pilbara Block (Dunlop et al., 1978). These are probably about 3500 Ma old. The microscopic objects reported by Marshall et al. (1964) from the opaline cement of a quartzite are indis-
363 tinguishable from inorganic crystals and their aggregates. Moreover, the opaline cement may be much younger than Archaean. Muir et al. (1974a) discovered filamentous microfossils resembling fungi in these rocks; they believe the filaments were introduced during the Tertiary weathering cycle. No definite stromatolites are known from rocks of this age in Australia. Many of the Archaean sediments are carbonaceous and studies designed to elucidate the relative importance of biologic and metamorphic effects are presently under way ; these will allow comparison of the C-isotopic composition of this organic matter with that of the Barberton area of South Africa (Oehler et al., 1972). PROTEROZOIC (2500-570
Ma)
A . 2500-1 770 MU The oldest undoubted microfossils known from Australia (Fig. 6.1) are Those of the Frere Formation of the Nabberu Basin (Walter et al., 1976b); and the Duck Creek Dolomite, well dated at 2020 f 165 Ma (Compston and Arriens, 1968), of the Hamersley Basin (Knoll and Barghoorn, 1975b). The sediments of the Nabberu Basin are only poorly dated, but almost certainly predate the 1720 Ma-old metamorphic event (Compston and Arriens, 1968) which affected the sediments of the Hamersley Basin, and the microfossiliferous iron formations of the Frere Formation are plausibly correlated with the iron formations of the Hamersley Group, underlying the Duck Creek Dolomite in the Hamersley Basin. Thus the Frere Formation microfossils are most likely to be about 2000 Ma old. They occur in siliceous oncolites and peloids in a granular iron formation that is remarkably similar t o those of the Animikie Basin in North America. The microfossils are virtually identical with those of the Gunflint Iron Formation of the Animikie Basin (Barghoom and Tyler, 1965) but are not as well preserved. The assemblage is dominated by the mostly 1-2 pm-wide filaments of Gunflintia minuta Barghoorn, which seems t o have been responsible for the construction of the oncolites. Next most abundant is Huroniospora spp., which consists of subspherical to ellipsoidal unicells 3-12 E.tm wide, some of which were preserved during the process of division. Rosettes of fine filaments identified as Eoastrion simplex Barghoom are moderately frequent, and commonly occur in clusters within the laminae of the oncolites. Nonseptate 4-9 pm-wide tubes of Animihiea sp. nov. also occur moderately frequently. Among the other microfossils so far discovered the most significant is the distinctive Kahabehia u m bellata Barghoom, which consists of an “umbrella” connected by a stalk to a bulb. In the Frere Formation K. umbellata is rare. The Gunflint microbiota occurs in two facies: columnar stromatolites in which the dominant microfossils are G. minuta and Huroniospora spp., and stratiform cherts rich in E. simplex, E. bifurcatum and K. umbellata
364
Fig. 6.1. Fossil cyanophytes and bacteria from stromatolites of the Early Proterozoic Frere Formation, Nabberu Basin (2.5 km northeast of Camel Well). All are preserved as iron oxide casts in chert. Scale in a, 1 0 0 pm; in b, 1 0 p m ; in c, 20 p m ;that in c applies to photographs c-j. a and b, Gunflintia minuta Barghoorn. c and d, Animihiea sp. nov., in c with an internal G. minuta trichome; e, f and g, Eoastrion simplex Barghoorn; h-j, Huroniospora sp. Thin section BHP5319Dl. Reproduced by permission of the Editor of Nature.
(Awramik, 1976). In the Frere Formation these two microfossil assemblages occur together. Although stratiform cherts are present, those which have been searched are barren of microfossils. The mode of occurrence of the Frere microfossils has been interpreted as indicating that all elements of the microbiota were benthic (Walter et al., 1976b);in contrast, Awramik (1976) suggests that in the Animikie Basin, Eoastrion and Kahabehia were planktic. Knoll and Barghoorn (197513) state that the Duck Creek microfossils come from non-stromatolitic black cherts within stromatolitic dolomites
365 (iron formations are not known from this unit). They report small spheroids comparable to Huroniospora Barghoorn and somewhat larger forms which may exceed 20 pm in diameter. “Filaments are uncommon, but those present are referable to the Gunflint taxa Animihiea Barghoorn and Gunflintia Barghoorn. The most abundant entities are profusely branched budding bacteria morphologically identical to the modern iron and manganese precipitating bacterium Metallogenium Perfil’yev and Gabe. A similar non-stromatolitic facies of the Gunflint formation . . . contains the same suite of organisms in approximately the same relative abundances. ” The biological affinities of the Gunflint-Frere-Duck Creek suite of microfossils are not entirely clear. Gunflintia minuta is conventionally interpreted as a cyanophyte (“blue-green alga”) because of the occurrence in the Gunflint Iron Formation and the Pokegama Quartzite of filaments with enlarged cells which may be heterocysts and akinetes (Licari and Cloud, 1968; Cloud and Licari, 1972). Both these specialized cell types occur only in cyanophytes. Recent work on the preservation of microfossils (Knoll and Barghoorn, 1975a) shows the need for caution in interpreting morphological features. Until preservation is better understood it would be premature to accept the interpretation of the enlarged cells as heterocysts and akinetes. Filamentous bacteria which are morphologically indistinguishable from G. minuta (except in the lack of heterocysts and akinetes) are abundant in some Holocene stromatolites and are a common component of the stromatolite microbiota (Walter et al., 1972, 1976a; Golubic, 1976). It is not suggested that all individuals of G. minuta are bacteria, particularly as some occur within sheaths (Walter et al., 1976a; Awramik and Barghoorn, 1977). This is a common feature of extant cyanophytes although it also occurs in some bacteria. G. minuta may be an assemblage of a cyanophyte and a bacterium, particularly in view of the relatively wide size range of G. minuta trichomes. At least some of the Animihiea sp. nov. tubes are the sheaths of G. minuta. Huroniospora spp. are interpreted as cyanophytes (Schopf, 1970; cf. Darby, 1974). The very close resemblance between Eoastrion and the extant Fe and Mn oxidizing bacterium Metallogenium was first noted by Cloud (1965), and it is generally accepted that Eoastrion was a similar bacterium. An extant homeomorph of Kahabekia urnbellata has been described from soils: it lacks a membrane-bound nucleus and apparently also lacks photosynthetic pigments, although Siegel and Siegel (1968), who described this organism, do not explicitly state this latter point. Thus all of the microfossils from the Duck Creek Dolomite and the Frere Formation appear to be procaryotic. The ubiquitous 5-10 pm-wide spheres and spherulites in Precambrian iron formations, including those of the Hamersley Basin (La Berge, 1967, 1973), are most likely to be abiogenic and the products of crystallization of originally gelatinous sediment (J.H. Oehler, 1976). The more complex “Eosphaera-like” spheroids reported by La Berge (1967, 1973) from many iron formations including some in the Archaean of Western Australia cannot
366 be readily interpreted. The report by Karkhanis (1976) of fossil iron bacteria preserved in ferroan carbonate of the Brockman Iron Formation of the Hamersley Basin needs t o be confirmed.
Fig. 6.2. The stromatolite Alcheringa narrina Walter o n ripple marks. Pillingini Tuff, Fortescue Group, Mount Herbert, Pilbara Block. Scale bar 2 cm. This is t h e oldest known Australian stromatolite (2300 Ma).
367 Stromatolites are abundant in Early Proterozoic rocks of Western Australia and the Northern Territory (Walter, 1972). The oldest are Alcheringa narrina Walter (Fig. 6.2) and Gruneriu f. nov. from the Fortescue Group of the Hamersley Basin (Walter, 1972), and other forms from the possibly correlative lower units of the Nabberu Basin (Preiss, 197613). The Fortescue Group may be as old as 2329 k 89 Ma (Lewis et al., 1975). Stromatolites and “medusae” reported from the iron formations of the Hamersley Basin (Edgell, 1964) are abiogenic, diagenetic structures (Walter, 1970). The Duck Creek Dolomite of the Hamersley Basin is richly stromatolitic (Walter, 1972). The Stromatolites of the Hamersley and Nabberu Basins are of considerable biostratigraphic interest in that many are destinctive, unlike younger forms, and indeed occur on other continents in rocks of a similar age (Walter, 1972); but others are similar t o younger forms at group, i.e. form-genus, level (Preiss and Walter, 1975; Preiss, 197613). The palaeobiological interpretation of stromatolites is in its infancy but we can say that the fabric of Alcheringa narrina indicates construction by a microorganism with coarse filaments, or perhaps bundles of filaments, like some species of the cyanophyte Rivularia (Walter, 1972). Already by Fortescue Group time (-2300 Ma) and especially by Duck Creek Dolomite time (- 2000 Ma) stromatolite-constructing microorganisms were abundant and widespread and living in a variety of environments. From then on stromatolites were a normal component of Australian Proterozoic sediments; they will be discussed only when they are of special interest. B. 1770-1400 MU
Diverse and abundant microfossils are known from rocks of this age in the McArthur and Birrindudu Basins, and to a lesser extent in the Mount Isa Geosyncline. They occur in a variety of rock types, including shales and stromatolites of various kinds, and so are likely t o provide a good sample of the microbiota of the time. They contain the earliest evidence of eucaryotic organization and in this are consistent with observations on other continents - indeed, these particular microbiotas are playing a key role in the establishment of the time of origin of the eucaryotes. Sponge spicules have been reported from the McArthur Basin (Dunn, 1964); however, rocks similar to those containing the “spicules” from elsewhere in the basin are now known to be tuffites and in these the “spicules” are volcanic shards (Croxford and Jephcott, 1972). It is necessary first to examine critically the datings of the fossiliferous units that will be described. The Bungle Bungle Dolomite of the east Kimberley region is correlated with similar rock units in the region which form part of the Birrindudu Basin (Plumb and Derrick, 1975). Granites in the basement t o this basin are dated at 1820-1720 Ma, the youngest being 1720 f 8 Ma (Rb-Sr whole-rock dating: Page et al., 1976). The basal unit of
368 the Birrindudu Basin is the Gardiner Sandstone which contains glauconites. Datings derived from the glauconites are 1560 k 20 Ma (K-Ar) and 1500 k 108 Ma (Rb-Sr isochron) (Page et al., 1976). Unconformably well above the Bungle Bungle Dolomite is the Wade Creek Sandstone which has yielded a Rb-Sr whole-rock shale isochron dating of 1128? 110 Ma (Plumb and Derrick, 1975). Plumb and Derrick (1975) correlate the sediments of the Birrindudu Basin with the McArthur Group in the McArthur Basin. The microfossiliferous units of the McArthur Basin are within the McArthur and Roper Groups. The Cliffdale Volcanics, unconformably below the McArthur Group, give a Rb-Sr whole-rock isochron of 1770 k 20 Ma (Plumb and Derrick, 1975). The Hobblechain Rhyolite and Packsaddle Microgranite, immediately underlying the McArthur Group, are dated (Rb-Sr whole-rock isochron) at about 1575 Ma (Plumb and Sweet, 1974). A suite of glauconite datings from the Roper Group, unconformably above the McArthur Group, indicates a probable minimum age of 1390 Ma. The Roper Group is intruded by dolerite sills which yield K-Ar mineral dates of from 876 to 1280 Ma; the latter is regarded as a minimum age for the sills. The Roper Group is correlated with the South Nicholson Group which overlies the sediments of the Lawn Hill Platform, including the microfossiliferous Paradise Creek Formation. The rocks of the Lawn Hill Platform can be correlated with those in the Mount Isa Trough and the McArthur Basin; using the datings from these areas the Paradise Creek Formation can be shown to be younger than 1537 40 Ma (Sybella Granite). Widespread orogenesis, dated at 1400-1450 Ma, apparently postdates the Paradise Creek Formation and predates the South Nicholson Group (Plumb and Sweet, 1974; Plumb and Derrick, 1975). Thus there is a complex web of correlations and datings stretching across northern Australia; this is presented above in a much simplified form. It may be summarized as follows: The Bungle Bungle Dolomite, McArthur Group and Paradise Creek Formation are probably penecontemporaneous and 1500-1550 Ma old; the Roper Group is probably 1400 Ma old. Within the McArthur Group microfossils are known from the Amelia Dolomite, Emmerugga Dolomite, Cooley Dolomite, HYC Pyritic Shale Member of the Barney Creek Formation, and the Balbirini Dolomite. In the dolomites, the microfossils are found in secondary black chert lenses within carbonate stromatolites which are stratiform, nodular and bulbous. The microfossils of the HYC Pyritic Shale Member occur within the dolomitic shale itself and also within small chert nodules in the shale. The microfossils so far described from the Bungle Bungle Dolomite are also from secondary cherts in stratiform, nodular and bulbous stromatolites, while those from the Paradise Creek Formation are from a variety of silicified columnar, bulbous and nodular stromatolites. Overall, this suite of microfissils represents a wide range of environments.
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369
Fig. 6.3. Myxornorphu junecekii Muir, a possible cyanophyte from t h e 1500 Ma-old Amelia Dolomite, McArthur Basin (in core from Tawallah Pocket No. 1 a t a depth of 41.9 m ) . Scale bar 1 0 pm. Photograph by M.D. Muir, reproduced by permission of t h e Association of Australasian Palaeontologists.
Muir has described 19 taxa of microfossils from the Amelia Dolomite including 1 0 new species, 2 new genera and one new combination (Croxford et al., 1973; Muir, 1974, 1976). The assemblage is dominated by unicellular forms. Although Muir (1976) and Oehler et al. (1976) tentatively suggest that some of the micro-organisms may have been eucaryotic, the available evidence for this is weak and amenable to other interpretations. On the other hand, as Muir (1976) points out, many of the taxa are closely comparable with extant cyanophytes. The rarely septate t o nonseptate thick-walled filaments of Eomycetopsis have been considered t o be phycomycetous fungi (and thus eucaryotic), but they are also similar to certain algae and at present their systematic position is uncertain. Muir (1976), following Darby (1974), suggests that some specimens of Huroniospora microreticulata Barghoorn from the Amelia Dolomite “may be yeast-like” because of the presence on the unicells of apparent buds. However, in view of the irregularity in cell form of some extant unicellular cyanophytes (Desikachary, 1972) and the presence of bud-like protuberances resulting from nannocyte formation in the cyanophyte Aphanothece bullosa var. major Geitler (Desikachary, 1959, pp. 38 and 43), this suggestion is unwarranted. Muir (1974, 1976) has described from the Amelia Dolomite numerous examples of two unusually large and complex microfossils : A meliaphycus croxfordii and Myxomorpha janecekii (Fig. 6.3). These fossils are unknown
370 from elsewhere. Both have relatively large cells for organisms of this age (up to 18.6 pm long in M . janecekii). M . janecekii shows cellular differentiation, with subparallel to radially arranged elongate cells capped by spheroidal cells; some specimens have a non-cellular basal attachment structure. A . croxfordii forms a cylindrical t o cup-shaped thallus, up t o 300 pm wide, which stands perpendicular to the sedimentary layering. The basal part is noncellular. Although Muir (1976) suggested resemblances with eucaryotes, cellular differentiation is common in cyanophytes, and M. janecekii can be compared with chamaesiphonalean cyanophytes such as Dermocarpa. A . croxfordii resembles the entophysalidacean cyanophyte Cyanostylon. The last piece of evidence for the presence of eucaryotes in the Amelia Dolomite is the enveloped tetrahedral tetrad of cells described by Oehler et al. (1976). Its significance is discussed by the authors. While these configurations are best interpreted as mitotically produced green algal autospores, they may result from cell-shifting in originally non-tetrahedral tetrads, and then could be either cyanophycean or eucaryotic. So at present there is no compelling reason for interpreting any members of the Amelia Dolomite assemblage as eucaryotic. In the Amelia Dolomite, mucilage impressions lacking preserved cells are abundant and frequently they delineate stromatolite laminae. Where cells are preserved they are often within mucilage (Huroniospora psilata, Sphaerophycus reticulatum 2, Myxococcoides reniformis, M. minor, M. kingii, M . minutus, Palaeoanacystis vulgaris, P. plumbii), some of which forms multilayered sheaths (Sphaerophycus parvum, S. tetragonalis). As Muir (1976) points out, these cells are very similar t o extant chroococcalean cyanophytes. The Amelia Dolomite assemblage is closely comparable with modern chroococcalean stromatolitic mats such as those of the intertidal zone of Shark Bay, Western Australia (Logan et al., 1974; Golubic, 1976). The stromatolites also are similar (Fig. 6.4). Filamentous microfossils are generally rare, although locally abundant. Besides Eomycetopsis which has already been mentioned there are fine filaments cf. Gunflintia minuta Barghoorn and G. oehlerae Muir which can be compared with oscillatoriacean cyanophytes. An assemblage of manganiferous microfossils from the Cooley Dolomite Member has been briefly reported by Muir et al. (1974b). These include the bacterium Eoastrion. Well preserved microfossils are abundant in silicified stratiform stromatolites of the Balbirini Dolomite (D.Z. Oehler, 1978). The Balbirini assemblage very closely resembles that of the Amelia Dolomite. The microfossils from the HYC Pyritic Shale Member of the Barney Creek Formation represent an entirely different environment from the other McArthur Basin assemblages. Hamilton and Muir (1974) described possibly biogenic organic spheroids, and J.H. Oehler (1977) found abundant microfossils in chert nodules in pyritic carbonaceous shales of huge Pb-Zn sulphide ore bodies. Included in the assemblage are seven species of bacteria, twelve species of cyanophytes, and two species referable either t o the cyanophytes
371
Fig. 6.4. Stratiform stromatolites in the Balbirini Dolomite, McArthur Basin, were constructed by pustular mats of the cyanophyte Eoentophysalis belcherensis Hofmann. Thin section of a black chert lens in the stratiform stromatolites (from near Balbirini Homestead). a. “Pustular mat” fabric, scale bar 1 mm. b. E. belcherensis from within the laminae of the stromatolite, scale bar 20 pm. Photographs by M.D. Muir.
372
Fig. 6.5. Fossil bacteria preserved as pyrite casts in thin section of chert from the 1500 Ma-old HYC Pyritic Shale Member of the Barney Creek Formation, McArthur Basin (from drill core of the McArthur Pb-Zn deposit). A and B, Ferrimonilis uariabile J.H. Oehler, A in transmitted light, B in reflected light, C-G. Biocatenoides incrustata J.H. Oehler, transmitted light. Photographs by J.H. Oehler, reproduced by permission of the Association of Australasian Palaeontologists.
or eucaryotic algae. The study of this assemblage has contributed significantly to the understanding of ore genesis in the shales. About 50% of the fossils observed are bacteria (Fig. 6.5), an unusually high proportion compared t o other assemblages. The bacteria are pyritized, whereas the cyanophytes are organically preserved, which suggests that the bacteria inhabited anoxic, iron sulphide-rich muds at the bottom of the depositional basin. No mat-forming cyanophytic communities are preserved. The cyanophytes are interpreted as a “plankton rain” and detritus from shallow near-shore waters. Apparently the basin floor was not suitable for habitation by cyanophytes, suggesting that it lay below the photic zone (J.H. OehTer, 1977). The metabolic activities of the bacteria probably maintained reducing conditions within
373 the carbonaceous muds, thus allowing the preservation of fine-grained sulphide minerals precipitated in the mud. Further evidence of biological activity within the sediments comes from the sulphur isotope values for some of the pyrite in the shale. These suggest that the sulphide ions were derived from seawater sulphate bacterial reduction (Smith and Croxford, 1973). Thus there probably were sulphate reducing bacteria in the sediment. The presence of minute sulphide (pyrite?) crystals in many of the McArthur Group microfossils (Amelia and Balbirini Dolomites) leads to the same conclusion. Microfossils have been briefly described from cherts in stromatolitic dolomites of the Bungle Bungle Dolomite (Diver, 1974). The assemblage is similar to that of the Amelia Dolomite in that it is dominated by small coccoid cells. “The size, vegetative form, and morphology of most of the microbiota are comparable with living cyanophycean algae. Coccoid forms with layered walls resemble members of the Chroococcaceae, and the septate filaments look like trichomes of the Oscillatoriaceae. Living oscillatoriaceans possess the facility of gliding motility and the non-septate filaments may be abandoned sheaths of such algae” (Diver, 1974). Some of the coccoid cells contain dark spots about 1pm wide and others include “apparently membrane-bound structures 2-3 pm in diameter containing minute dark granules”. Diver suggests that these may be the preserved organelles of eucaryotic algae. Detailed studies, such as those of D.Z. Oehler (1976, 1977) on Bitter Springs Formation microfossils, will be required to check this highly significant possibility. To date there are insufficient data available t o allow the Bungle Bungle intracellular bodies to be distinguished from collapsed cytoplasm. The microfossils of the Paradise Creek Formation are poorly preserved but particularly significant because they occur within five different forms of stromatolites, three columnar or pseudocolumnar and two bulbous and nodular (Licari et al., 1969; Licari and Cloud, 1972). Each form of stromatolite appears t o contain a distinct assemblage of microfossils. Licari and Cloud (1972) identify the stromatolites as aff. Conophyton, Eucupsiphoru purudisu Cloud and Semikhatov, “weakly branching and much interconnected columnar stromatolites of ambiguous affinity”, “large domal Colleniu” and “small domal Colleniu”. The study of additional specimens of the “aff. Conophyton” by Walter (unpublished) has shown that it definitely is Conophyton and resembles C. gurgunicum Korolyuk (emend. Komar, Raaben and Semikhatov). The C. cf. gurgunicum contains segmented unbranched filaments arranged roughly parallel to the laminae. The filaments average 1.8 pm in width. These are interpreted by Licari and Cloud (1972) as nostocalean cyanophytes which constructed the conophytons. This is supported by comparison with modem conophytons, which are built by the 0.8-1.2 pm-wide filaments of the nostocalean (oscillatoriacean) cyanophyte Phormidium tenue var. grunuliferum Copeland (Walter et al., 1976a); in the modern examples the filaments are also parallel to the laminae. By com-
374 parison with the modern conophytons it can be confidently suggested that the filaments in C. cf. garganicum are of a phototactic cyanophyte, i.e., one capable of gliding motility in response to a stimulus by light. The columns of Eucapsiphora paradisa contain at their margins microfossils which are remarkably similar t o the extant chroococcacean cyanophyte Eucapsis. The 2-3 pm-wide coccoid cells of this fossil are arranged in a regular cubic lattice in colonies 12-18 pm wide. Licari and Cloud (1972) refer to this microfossil as Eucupsis(?). The third columnar and pseudocolumnar form of stromatolite contains very poorly preserved septate filaments averaging 2.2 pm in width. The large bulbous and nodular stromatolites (“Collenia”) contain coccoid cells averaging 4.4 pm in diameter. Similar but smaller stromatolites contain fragmentary segmented filaments and coccoid unicells 10--12 pm in width. Kline (1975) has very briefly reported the occurrence of microfossils similar to the Gunflint Iron Formation taxa Eoastrion simplex Barghoorn and E. bifurcutum Barghoorn in chert from the Paradise Creek Formation. These are well enough preserved to allow electron microscopic studies which have confirmed the branching of filaments in the form resembling E. bifurcatum. The radiating processes are variously septate and nonseptate, the
Fig. 6.6. Partially silicified Conophyton stromatolites which have weathered out of dolomite of the 1500 Ma-old Tooganinnie Formation, McArthur Basin, 10 km NNW of old Tawallah homestead. Scale 10 em. Reproduced with the permission of the Director, BMR; previously published by the AGPS in BMR J. Aust. Geol. Geophys.
375 former with coccoid cells. Rounded terminal cells 0 . 2 4 . 3 pm wide occur on some processes. Stromatolites are abundant in rocks 1770-1400 Ma old and include a wide range of forms. Few have been treated taxonomically: those that have include C o n o p h y t o n garganicum (Fig. 6.6) from various units in the McArthur Group (Cloud and Semikhatov, 1969; Walter, 1972) and Eucapsiphora paradisa from the Paradise Creek Formation. One of us (M.R.W.) has recently observed numerous forms of columnar branching stromatolites in the McArthur Group, including one which is tungussiform. Rocks in this age range contain the oldest C o n o p h y t o n in Australia, although older occurrences are known from other continents.
C. 1400-950 Ma Apart from as yet unstudied occurrences in ferruginous ooids and granules from the South Nicholson Basin (Walter, unpublished), and of “unicellular organic bodies of unknown origin” from the same basin (Randal, 1966), the only known microfossiliferous rock unit within this age interval is the McMinn Formation of the Roper Group (McArthur Basin). The dating of the Roper Group at about 1400 Ma was discussed in the previous section. Abundant and diverse microfossils have been found in core samples of the Maroak Sandstone and Kyalla Members of the McMinn Formation (Peat et al., 1978). They occur within black, pyritic, carbonaceous shale representing the sulphide facies of a chamositic ironstone. The microfossils are thought to be allochthonous, and t o have lived in oxygenated environments as plankton and benthos. The range of sedimentary facies present in the formation indicates fluviatile, transitional and near-shore marine environments. The microfossils are especially significant for two reasons: they derive from non-carbonate environments which are as yet little studied, and they are of an age when it appears that critical evolutionary events may have occurred, i.e. the advent of sexual, eucaryotic organisms. Peat et al. (1978) have recognized life cycles and degradational series among the plexus of intergrading forms of microfossils, in a study that will enable many improvements t o be made t o the taxonomy of such fossils. Spherical microfossils up t o 600 pm wide are present (Fig. 6.7) and are comparable with the Late Precambrian form Chuaria. They are very much larger than any previously described forms. The relatively enormous size of these fossils suggests that they are eucaryotic algae. This is consistent with Schopf’s (1975b) observations on the temporal distribution of cell sizes, in which he notes an abrupt increase in average cell diameters at about 1400 Ma; he suggests this may mark the origin of eucaryotic algae. Filamentous microfossils in the McMinn assemblage are also relatively gigantic, reaching a width of 120 pm (Fig. 6.7). These are larger than extant cyanophytes, but nonetheless closely resemble the living Lyngbya majuscula.
376
377 The McMinn assemblage is strikingly different from penecontemporaneous chert-carbonate microbiotas from around the world, and points t o the need for greatly increased effort in studying floras from siliciclastic rocks. McKirdy (in Peat et al., 1978) has reported on the organic geochemistry of the McMinn Formation, in a study of both the extractable (soluble) organic matter and the kerogen. This allows the correlation of diagenetic and metamorphic grade as determined from the organic matter with the state of preservation of the microfossils. The many parameters measured and observed indicate that there has been some mobilization of organic matter during maturation t o the stage of early mesodiagenesis. The shales are “oil prone”. Stromatolites of this age are poorly known in Australia. They include those of the approximately 1100 Ma-old Bangemall Group of Western Australia: C o n o p h y t o n garganicum australe Walter and Baicalia capricornia Walter (1972). This assemblage has been correlated with those of the Middle Riphean (1350 f 50-950 f 50 Ma) of the USSR (Preiss, 1977). Poorly preserved, finely filamentous microfossils have recently been found in these C o n o p h y ton (Walter, unpublished).
D. 950-750 Ma Rocks in this age range in Australia contain two of the cardinal microbiotas of Precambrian palaeobiology: those of the Bitter Springs Formation (Amadeus Basin), and the Skillogalee Dolomite and its correlatives (Adelaide Geosyncline). Other rock units in the Adelaide Geosyncline are also microfossiliferous (?Blyth Dolomite, Myrtle Springs Formation, Auburn Dolomite) and microfossils have recently (Walter, unpublished) been discovered in the rim rocks of the Madley Diapir, Officer Basin; Kemp (1976, p. 26) mentions other discoveries in the Officer Basin. Despite arguments t o the contrary, there is good evidence that these microbiotas include eucaryotic algae (see below). Stromatolites of this age are well known in Australia and have played a key role in discussions on the validity of stromatolite biostratigraphy. Many come from the microfossiliferous units and increasing attention is being paid to their palaeobiological significance. The dating of the Bitter Springs Formation reveals one of the successes of stromatolite biostratigraphy. Until Fig. 6 . 7 . Microfossils from black, pyritic, carbonaceous shale of the 1400 Ma-old McMinn Formation, McArthur Basin. a. cyanophyte trichome. b and c . Eucaryotic alga (?) with internal granular body. Scale bars, 10 pm. a and b in acid macerates; c in thin section; transmitted light. From 22.6 m level in diamond drill hole 67. Photographs by C. Peat; a is reproduced by permission of the Association of Australasian Palaeontologists; b is reproduced with the permission of the Director, BMR; previously published by the AGPS in BMR J. Aust. Geol. Geophys.
378 1972 all available radiometric data indicated that the Bitter Springs Formation is younger than 1370 Ma, possibly younger than 1200 Ma; and older than 740 f 30 Ma (see Walter, 1972, pp. 38-40). A study of the stromatolites led Walter (1972) to correlate the upper (Loves Creek) member of the formation with the Upper Riphean of the USSR. On this basis he reasoned that the Loves Creek Member is no older than 950 f 50 Ma, the age of the base of the Upper Riphean. This has now received strong support from a recent study (Marjoribanks and Black, 1974) that indicates that some of the basement metamorphics below the Bitter Springs Formation have an age of 1076 f 50 Ma (Rb-Sr whole-rock isochron). The minimum age of the Bitter Springs Formation is based on the correlation of the unconformably overlying glacigenic rocks (Areyonga Formation) with the lower tillites of the Kimberley region to the north, and the Rb-Sr whole-rock shale isochron dating of these at 740 f 30 Ma. There is also a Rb-Sr whole-rock sediment dating on the Pertatataka Formation, above the Areyonga Formation, of 730 45 Ma (Compston and Taylor, 1969). Neither of the last two datings can be accepted with confidence. The Bitter Springs Formation is therefore younger than 950 f 50 Ma and possibly older than 740 f 30 Ma. The Skillogalee Dolomite of the Burra Group in the Adelaide Geosyncline has until recently been considered to be about 1000-1100 Ma old, on the basis of isotopic dating and stromatolite data. New isotopic data (Cooper and Compston, 1971; Cooper, 1975) now indicate that it is probably younger than 867 f 32 Ma (Rb-Sr whole-rock isochron). The discovery and study of new stromatolitic sequences show that the previous correlation of Skillogalee Dolomite with the Middle Riphean is not justified, “and a Late Riphean (950 f 50-680 20 Ma) age is preferred for the basal beds and for the whole early Adelaidean” (Preiss, 1976a). The Sturtian tillite unconformably above the Skillogalee Dolomite is correlated with the tillite of the Kimberley region dated at 740 f 30 Ma. Thus the best interpretation of the presently available data is that the Skillogalee Dolomite is younger than 867 k 32 Ma and older than 740 k 30 Ma. The same age constraints apply to the other microfossiliferous units from the Adelaide Geosyncline mentioned in this section. The microbiota of the Loves Creek Member of the Bitter Springs Formation is very well known and has been reviewed frequently (Schopf, 1968, 1970, 1972, 1975b; Schopf and Blacic, 1971). It is known from several localities and everywhere occurs within diagenetic chert lenses in flat-laminated dolomites that may be stratiform stromatolites (Fig. 6.8). The dolomites are interbedded with red siltstones and silty dolomites in a sequence suggestive of supratidal deposition. The stratigraphically significant columnar stromatolites (Fig, 6.9) from elsewhere in the Loves Creek Member were formed subtidally or intertidally (Walter, 1972); despite extensive searching no microfossils have yet been found in them. However, McKirdy (1976) reports that the columnar stromatolite Acaciella australica Walter
*
*
379
Fig. 6.8. Microfossiliferous black chert lenses in dolomitic stratiform stromatolites of the 800--900 Ma-old Bitter Springs Formation, 1.2 km N N E of Ross River Tourist Chalet, Amadeus Basin. Scale 10 cm.
contains organic matter with an “alkane pattern (range n-CISto n-C,,, maximum n-C18)considered typical of blue-green algae”. Fifty species of micro-organisms have been described from the Loves Creek Member, including bacteria (3 species), filamentous (27 species) and coccoid (11 species) cyanophytes, eucaryotic algae (7 species) and probable fungi (2 species). The validity of the taxonomy and of some of the biological interpretations has been disputed by Knoll and Barghoorn (1975a) - this is discussed below.
Fig. 6.9. The stromatolite Acaciellu australica Walter, from the 800-900 Ma-old Bitter Springs Formation, Amadeus Basin, as it would have appeared during growth. Reproduced by permission of the Palaeontological Association, London, from the original figure in Special Papers in Paleontology, 11, (1972).
380
381 Schopf has repeatedly emphasized that “in ecologic setting, general morphology, detailed cellular anatomy, growth habit and mode of reproduction, many of the Bitter Springs cyanophytes bear striking resemblance t o living blue-green algae (Fig. 6.10, a-b). The Cyanophyceae provide an impressive example of evolutionary conservatism, a feature attributable t o the wide ecologic tolerance, versatile physiology and unusually stable genetic system characteristic of the class”. Furthermore, “in ecologic setting, laminated organization, and in general biologic makeup, the Bitter Springs microflora seems analogous t o modern mat-building algal communities that produce stratiform stromatolites . . .; several of the relatively common Bitter Springs cyanophytes are morphologically similar t o the dominant species of modern stromatolitic biocoenoses” (Schopf, 1972, p. 76). But as Schopf points out this conservatism may conceal radical and significant cytoplasmic and physiological evolutionary changes. Morphogenetic studies of the microfossil-bearing stromatolites may eventually reveal some such changes. There is now intense debate on the interpretation of the cellular morphology and ultrastructure of fossil micro-organisms. This centres on microfossils from the Bitter Springs Formation and concerns primarily the interpretation of some of them as eucaryotes. The debate is of great interest and significance because it deals with the fundamental problem of distinguishing primary (biological) from secondary (degradation, diagenetic and metamorphic) features. Or, it hinges the decision as t o whether there is any evidence for the origin of the eucaryotes well before the appearance of the oldest softbodied metazoan fauna in the terminal Precambrian of Ediacara and elsewhere. There are four main reasons for believing that the Bitter Springs and Skillogalee microbiotas include fossil eucaryotes: (1)the frequent presence in the Bitter Springs of spheroidal unicells with internal contents interpreted as organelles (procaryotes have no organelles) (Fig. 6.10, h and i); (2) the presence in the Bitter Springs of a tetrahedral tetrad of cells most probably resulting from either mitosis or meiosis (which do not occur in procaryotes) ; ( 3 ) the large average diameter of several taxa of Bitter Springs spheroidal cells (larger than expected in procaryotes); Fig. 6.10. Microfossils from black chert of the 800-900 Ma-old Bitter Springs Formation, Ellery Creek, 80 km west of Alice Springs, Amadeus Basin. a. Obconicophycus amadeus Schopf and Blacic. b, c and e , Cephalophytarion variabile Shopf and Blacic, d. Veteronostacale amoenum Schopf and Blacic, f and g. Caudiculophycus rivularioides Schopf, h and i. Glenobotrydion aenigmatis Schopf. Scale in a applies to a--g. a-h; thin sections, transmitted light; i ; transmission electron micrograph of a cell released by maceration (w = cell wall, o = organelle, c = cytoplasmic residue). Photographs a--g by J.W. Schopf and J.M. Blacic, reproduced by permission from the Journal of Paleontology, and the Society of Economic Paleontologists and Mineralogists. Photographs h and i by D.Z. Oehler.
382 (4) the presence in a probable correlative of the Skillogalee Dolomite of large, branching tubular filaments best interpreted as siphonaceous eucaryotic (green or golden green) algae. The main criticism has concerned the recognition of organelles within unicells : “The three-dimensional morphology of microfossils is generally not preserved by the cell wall as has been previously assumed, but rather by one or more mucoidal sheaths . . . The putative organellar remnants are in reality degraded protoplasm representing the entire contents of the cell” (Knoll and Barghoorn, 1975a). The arguments against this are discussed in detail by D.Z. Oehler (1976, 1977) and Schopf and Oehler (1976). Arguments related to the other three points listed above are presented by Schopf and Oehler (1976), and the interpretation of the tetrahedral tetrad is thoroughly examined by Oehler et al. (1976). In summary, it must be concluded that while Knoll and Barghoorn (1975a) rightly sound a note of caution, the weight of evidence still favours the interpretation that rocks in the time interval discussed here contain fossil eucaryotic, and probably meiotic, algae. The diverse and well-preserved microbiota from Boorthanna, in the Adelaide Geosyncline, from a probable correlative of the Skillogalee Dolomite, is particularly interesting - not only because it includes the siphonaceous alga mentioned above but also because it occurs within both stratiform and columnar branching stromatolites (Schopf and Fairchild, 1973; M.R. Walter, personal observations). The columnar stromatolites resemble Baicalia. This is one of very few known occurrences of a fossil biota which constructed columnar branching stromatolites. The Skillogalee microbiota described by Schopf and Barghoorn (1969) is apparently from stratiform stromatolites; the filamentous cyanophyte Archaeonema longicellularis Schopf seems to have produced the laminated fabric of the stromatolites. Palaeopleurocapsa wopfnerii Knoll et al. (1975) from the Skillogalee Dolomite is the oldest known fossil representative of the cyanophyte order Pleurocapsales. Schopf (197513) reports the discovery of new microbiotas from two other formations in the Burra Group (Myrtle Springs Formation and Auburn Dolomite) and one possibly from the Blyth Dolomite of the River Wakefield Group underlying the Burra Group, Adelaide Geosyncline. These have yet to be studied.
E. 750-570 M a Rocks in this age range in Australia are most remarkable for their record of soft-bodied metazoans. Stromatolites are less abundant than in older rocks, and only one algal microbiota definitely of this age is known. That is from a siliceous ooid grainstone of the Julie Member of the Pertatataka Formation in the Amadeus Basin. Several peloids in the grainstone contain poorly preserved small spheroidal unicells and nonseptate filaments (Walter, unpublished). Spheroidal microfossils were reported from the Tapley Hill
383 Formation of the Adelaide Geosyncline by Timofeev (1966, 1974). Megascopic algal (“seaweed”) fossils are known from sediments of this age on other continents, but the only comparable record here is the brush marks possibly made by flexible floating bodies which Milton (1966) described from the Winnall Beds of the Amadeus Basin. McKirdy’s (1976) study of the organic geochemistry of Precambrian stromatolites included samples of Tungussia inna Walter from the Wonoka Formation of the Adelaide Geosyncline. In this preliminary study, alkanes were extracted from six identified groups of stromatolites of various ages. The alkane distribution in T. inna was found t o be virtually identical to that in PT. wilkatanna from the Skillogalee Dolomite and different from that in the other specimens studied. If these results can be confirmed by further studies they will be highly significant, not only in indicating that the alkanes are truly indigenous to the rock, and thus providing a method of investigating organic chemical contamination, but also because they will aid the biological interpretation of stromatolites. The oldest metazoan body fossils occur in the geological record of the latest Precambrian, probably only tens of million years before the beginning of the Cambrian. The rocks containing them are undoubtedly younger than the occurrences of the youngest Precambrian glacigenic rocks, in Australia and probably elsewhere. This means that they are younger than the final stages of a prolonged epoch of widespread major glaciations. Both phenomena are well documented in the upper Precambrian rocks of Australia. They are the objects of continuing, detailed studies and they have attracted worldwide interest. A causal relation between glaciations and the first differentiation of marine metazoan faunas has been suggested (Harland and Rudwick, 1964; Rudwick, 1964) and disputed (Glaessner, 1972). It seems plausible that a connexion exists between major climatic variations and significant geotectonic and palaeogeographic changes. Such changes could have been the causes of accelerated faunal differentiation in the sea and particularly on continental shelves. Not enough is known at present about the most ancient faunas and their dating to support postulated causal relations between physical and biological (evolutionary) events of distant times in any other than a purely speculative manner. The discovery of Late Precambrian faunas has made it possible t o eliminate or question a number of speculative hypotheses about drastic changes in the oxygen content of the atmosphere, the composition of sea water or the nature of the inhabited ocean floor at the beginning of the Cambrian. This discovery is significant, but much further evidence has to be awaited before it can be fully evaluated. The environment and composition of the Ediacara fauna, the oldest rich and varied marine fauna in Australia, are well known and its relations to faunas of similar age from other continents are relatively close. The Ediacara fauna (Fig. 6.11) occurs in the youngest part of the Adelaidean of the Flinders Ranges in South Australia, the Pound Sandstone
384 (or Quartzite). This formation consists of the mainly red Bonney Sandstone Member in its lower part and the mostly white and grey Rawnsley Quartzite Member above. This Member contains in its basal portion lensing, sandy and silty beds which are fossiliferous in many places, throughout a thickness of up to 112 m. The composition of the fauna is shown in Table 6.1 as a list of species with an indication of the quantitative distribution of major taxa in a collection of some 1500 specimens. Some of these organisms lived near the places where they are found. The Conulata and “Petalonamae” were more or less firmly attached t o the substratum and could not have been moved far. In most of the annelid worms and arthropods, growth series of individuals of different sizes are preserved. The absence, on many bedding planes, of mechanical sorting according to size and the presence of delicate, minute, juvenile specimens suggests that their post-larval development took place nearby. The chondrophoran hydrozoans (Fig. 6.11b) are represented by their float chambers which are similar to those of living genera (Porpita, Velella). They live near the surface of the water and feed on its rich plankton content. The medusoids occur on certain bedding planes in large numbers, which may indicate that they drifted inshore in swarms, as their descendants do today. The environment indicated by the fauna and by the character of the sediments was one of near-shore
Fig. 6.11. Metazoan fossils from the Late Precambrian. a. Hallidaya brueri, Wade, a fossil jellyfish from central Australia, X 2 ; photography by M. Wade. b - e F r o m Ediacara, South Australia. b. Chondroplon bilobatum Wade, a hydroxoan float, X 0.4, c. Beltanella gilesi Sprigg, a medusoid, X 0.75, d. Glaessnerina grandis (Glaessner and Wade), a coelenterate fron resembling a sea pen, XO.4, e. Dichinsonia costata Sprigg, a polychaete worm, X 1; photographs by University of Adelaide.
385
TABLE 6.1 List of Ediacara Fossils (named species only) COELENTERATA (67%) Hydrozoa Chondrophora Chondroplon bilobatum Wade (Fig. 6.11b) Ovatoscutum concentricum Glaessner and Wade Eoporpita medusa Wade Scyphozoa Brachina delicata Wade Kim berella quadrata (Glaessner and Wade) Conulata Conomedusites lobatus Glaessner and Wade Medusoid fossils Cyclomedusa davidi Sprigg Cyclomedusa plana Glaessner and Wade Cyclomedusa radiata Sprigg Ediacaria flindersi Sprigg Beltanella gilesi Sprigg (Fig. 6 . 1 1 ~ ) ?Lorenzinites rarus Glaessner and Wade Mawsonites spriggi Glaessner and Wade Medusinites asteroides Sprigg) Rugoconites enigmaticus Glaessner and Wade “Petalonamae” (=Coelenterata, possibly related to Pennatulacea) Pteridinium cf. simplex Giirich Charniodiscus (=Arborea)arboreus (Glaessner) Charmiodiscus oppositus Jenkins and Gehling Glaessnerina grandis (Glaessner and Wade) (Fig. 6.1 I d ) Glaessnerina longa (Glaessner and Wade) Phyllozoon hanseni Jenkins and Gehling ANNELIDA Polychaeta ( 2 5 % ) Dickinsonia costala Sprigg (Fig. 6.11e) Dickinsonia brachina Wade Dickinsonia elongata Glaessner and Wade Dickinsonia lissa Wade Dickinsonia tenuis Glaessner and Wade Spriggina floundersi Glaessner Marywadea ovata (Glaessner and Wade) ARTHROPODA (5%) Trilobitomorpha o r Chelicerata Praecambridium sigillum Glaessner and Wade Crustacea Branchiopoda Parvancorina minchami Glaessner (Systematic position unknown) Tribrachidium heraldicum Glaessner ( 3 % )
deposition in embayments formed and eventually filled by moving sand waves. This occurred in the final stages of the Precambrian history of the Adelaide Geosyncline. Soon after the stratigraphic change from the lower t o the upper member of the Pound Sandstone there occurred a widespread,
387 though not ubiquitous, brief phase of relative quiescence or at least reduced energy regime. In the quiet embayments where finer silt and clay particles were deposited, soft-bodied organisms came t o rest and were covered before or during decay, or their clay moulds were infilled by periodically renewed influx of sand. This partly explains the preservation of abundant soft-bodied organisms and enables us to reconstruct their morphology by observing differences between more or less resistant parts of their bodies (Wade, 1968). No organic substances are preserved in the highly oxidizing, porous, waterflushed environment of the Pound Sandstone. Part of the explanation for the preservation of soft-bodied invertebrates is seen in the stage reached by the fauna in its morphological evolution. It appears to include only small-particle feeders which were unable t o consume the available large bodies of worms or coelenterates. They are found distorted, contracted, occasionally torn by waves, but not damaged by predators. Food must have been available in the form of detritus resulting from decay and also as phytoplankton. This is not likely to be preserved in the sandy and silty environment in which the Metazoa occur. A number of trace fossils produced by sediment-feeding and “grazing” worm-like animals are further indications of rich benthic life. They were less differentiated than the trace fossils of the Lower Cambrian (Glaessner, 1969; Wade, 1970; Daily, 1973) and only about eight different forms have been recognized. The Ediacara fauna, named after the first locality at which it was found, is now known from many localities throughout an area 100 X 200 km in extent. Two sandstones, the Punkerri Hills Sandstone in the north of South Australia (27”2’S, 130”30‘E; Major, 1974) and the basal Arumbera Sandstone at Deep Well, 80 km south-southeast of Alice Springs (Glaessner, 1969), have each yielded a single but unmistakable fragmentary impression of a frond (petalodium) of a fossil coelenterate similar to one described from Ediacara. These fossils were first assigned to the genus Rangea from SouthWest Africa but the description of new material from that area has shown that it had been misinterpreted and that it does not occur in Australia. The two fragments from central Australia and similar specimens from Ediacara are not at present specifically or generically identifiable and further finds must be awaited. In central Australia there is a fossiliferous horizon which is younger than the Ediacara fossil from the base of the Arumbera Sandstone but older than the oldest Cambrian. It contains the medusoid coelenterates Hallidaya brueri Wade (Fig. 6.11a) and Shinnera brooksi Wade (Wade, 1969). The first of these species was found in the Arumbera Sandstone well above its base but below the horizon of the first Lower Cambrian trace fossils. Both species were originally described from the Central Mount Stuart Beds in the vicinity of Mount Skinner, about 160 km north of Alice Springs, and H. brueri has recently also been found elsewhere in the same region. Another fossil of apparently coelenterate affinities, Arumberia banksi Glaessner and Walter,
388 was described from three localities 14-22 km southwest of Alice Springs from the level in the Arumbera Sandstone a t which Hallidaya occurs (Glaessner and Walter, 1975). Discoveries of fossil soft-bodied animals of Late Precambrian age have been made a t about 1 2 widely scattered localities in other continents (Glaessner, 1971). Most of them have yielded only few specimens; at two localities only annelid worms or arthropods have been found but elsewhere coelenterates occur exclusively or dominate the fauna. The oldest known, and, next t o Ediacara, richest fauna is that of the lower Nama Group (Richter, 1955; Germs, 1972). Other occurrences are in England, northern Sweden, northern Russia, The Ukraine, northern and central Siberia, China, southeastern Newfoundland and North Carolina. Those assemblages that have been described resemble the Ediacara fauna in general appearance and composition. As far as is known a t present, they have collectively only five genera and two species in common with it. Their age range is important for the dating of the Australian faunas whose ages cannot be determined directly by isotopic methods. The oldest reported fauna of Ediacaran type is that of Charnwood Forest, Leicestershire, England, which is believed t o be about 680 Ma old, though its age determination is based on the K-Ar dating of only one sample of a dacitic intrusive while other similar rocks gave younger dates by the same and Rb/Sr methods. All other Late Precambrian occurrences of fossil Metazoa appear t o be younger. Their ages, most of which are based on K-Ar dating of glauconites, fall into the interval from about 640 to 620, 600, and possibly 580 Ma. They represent a time span comparable with the length of the Tertiary Period. They predate fossiliferous Lower Cambrian rocks in Scandinavia, the U.S.S.R., where lithological correlations are supported by palynological studies from many bore samples, and in North America. A consideration of all available evidence results in placing the Australian Ediacara faunas within the same time span (Glaessner, 1977). As far as we know a t the present time, there are n o metazoan body fossils which are older than the Ediacara fauna. The trace fossil Bunyerichnus dalgurnoi Glaessner was found some 1800 m below the horizon of the first Ediacara fossils. Several claimed occurrences of other Australian Precambrian trace fossils have since been disproved. Skolithos from the Wessel Islands off the north coast of Australia was thought t o be of Precambrian age because it was found in sandstones containing glauconite which was dated a t 780 Ma old. This dating is now thought t o be spurious since Cambrian trilobites have been found recently in the same sedimentary sequence not far above the sandstone (Plumb et al., 1976). The supposed medusoid Protoniobia wadea Sprigg is an abiogenic concretion from the Mount John Shale Member of the Wade Sandstone of northwestern Australia which is 1 1 2 5 f 100 Ma old (Plumb and Derrick, 1975). The reported finds of Skolithos from the tillitic Areyonga Formation and of Syringomorpha from the Winnall Beds of the Amadeus Basin in central Australia are not fossils but abiogenic sedimentary structures (Walter, unpublished).
389 Notwithstanding these failures, the evolutionary level of the Ediacara fauna and the history of life in early Precambrian time suggest that the metazoan animals originated more than 650 Ma ago, but that they were small and much less resistant to decay than those which have been fossilized in the youngest Precambrian strata. There are reports from other continents, so far insufficiently documented, of traces of bioturbation and occurrences of fecal pellets in burrows in older sediments, up to 900 or 1000 Ma old. It is possible that, when careful attention is given t o microscopic biogenic configurations of this kind in the abundant unmetamorphosed sediments of the Australian Late Precambrian, the fossil record of animal life will be extended to strata of that age. What has been achieved so far in the study of Precambrian palaeontology, particularly in Australia, has tended to give a new direction to efforts to define rationally and unequivocally what is meant by the base of the Cambrian. Agreement on this point would necessarily define what is meant by Precambrian. At this stratigraphic boundary the emphasis shifts away from biostratigraphy as we go back in earth history, but biostratigraphy does not cease to operate at the base of the Cambrian. There is now substantial agreement (Cowie and Glaessner, 1975) that the base of the Cambrian should be defined in a continuous sequence of strata, within the limits set by those containing the Ediacara fauna below and those containing distinctive Lower Cambrian trilobite faunas above. This means that there is a consensus of opinion that the Ediacara fauna pre-dates the beginning of the Cambrian; however, this is to be defined. Agreement is still being sought on a definition of a horizon at the base, within, or at the top of the Tommotian Stage or its equivalents which are at present considered as basal Cambrian. The first occurrence of fossil animals, of body fossils, or of “shelly” fossils is not a sufficiently precise marker for a world-wide stratigraphic horizon. Further studies are being directed towards providing facts on which the selection of a boundary strato-type section can be based, whether in Australia or elsewhere, and what useful alternatives for a decision on the placing of the boundary by consensus are available. If the end of the Precambrian is to be represented by a precise stratigraphic horizon it is not a natural datum t o be discovered but a point in a continuum fixed by agreement. The Australian studies have brought such an agreement closer. They are contributing to the still contentious definition of other boundary points in the time scale of the Precambrian. CONCLUSIONS
The Precambrian fossils of Australia are remarkably diverse, abundant and well preserved. It is unlikely that this continent is uniquely endowed with fossils, but what fossils exist are being rapidly discovered and described. So far the Archaean has yielded only carbonaceous microspheroids of uncertain origin. Sediments of this age are well exposed in the Pilbara region of
390 Western Australia. Archaean fossils are known from Africa and Canada; of these the stromatolites are undisputably biogenic (although there is dispute over whether they are bacterial or cyanophytic) but the “microfossils” of that age are contentious and may be abiogenic (Schopf, 1976). In Australia, as on other continents, the undisputed microfossil record commences with the Early Proterozoic. Concurrently, stromatolitic carbonate sediments first occur extensively, and from that time on through the remainder of the Proterozoic not one of the many sedimentary basins in Australia lacks them. Cyanophytes are the main constituents of the Early Proterozoic microfossil record, although bacteria are also abundantly represented. With some exceptions, yet to be confirmed, Early Proterozoic cyanophytes are simple forms of that phylum. The bacteria are diverse and in addition there are fossils the biological affinities of which are obscure. The microfossil assemblages of the 1500-1550 Ma-old McArthur Group and its correlatives are the earliest in Australia t o contain forms which have been considered t o be eucaryotes, but only bacteria and cyanophytes can be confidently identified. By 1400 Ma the evidence for the presence of eucaryotes in the microbiota is mounting, and it seems probable that within a few years more convincing evidence for the advent of eucaryotes will be discovered in rocks of about this age. Because of its age, the Roper Group fossil assemblage is critically significant in this regard. No significant microfossils are known from Australia in the interval of 950-1300 Ma. Then in the 800---go0 Ma-old Bitter Springs Formation and Skillogalee Dolomite and their correlatives we have two of the most important microfossil assemblages known. They include a wide range of cyanophytes (among them some complex forms), and for the first time provide convincing evidence of the former presence of eucaryotes and probably also of sexual reproduction. Possibly as many as 4 phyla of eucaryotes are represented (Chlorophycophyta, ?Rhodophycophyta, ?Pyrrophycophyta, ?Eumycophyta). The stage was set for the appearance of the Metazoa, or perhaps they had already evolved but their fossils have yet t o be recognized. Increasing attention t o suitable rocks of this age may well produce evidence of the earliest metazoans. It is not until the latest Precambrian, in rocks 640---570 Ma old, that the first undisputed metazoan body fossils are found. They are large and diverse, and include representatives of at least 3 phyla (Coelenterata, Annelida and Arthropoda). Effective predation and scavenging of animals apparently did not begin until the Cambrian. Whether evolution progressed steadily through the Precambrian or by a series of “quantum jumps” is a question we cannot yet answer. The micro-organisms and early animals did not live independently of each other or of their environment. Stromatolites are to be seen not only as potential biostratigraphic tools, but as the sedimentary record of microbial communities, the repository of otherwise unavailable physiological information on the constructing microorganisms, and not least, a source of food and shelter (and perhaps oxygen) for the earliest metazoans (Fischer, 1965;
391 TABLE 6.11 Location of Australian Precambrian sedimentary basins and approximate age of fossiliferous formations discussed in the text
Arumbera Sondstone Julie Member Winnall Beds
Central Mount Stuort
P w n d Sandstone Wonoka Fm
700
Tapley Hill Fm Burro Gp
800
Auburn Dolomite Myrtle Springs Fm Shillogalee Dolomite
900
IOOO]
1100
Bongemoll Gp
IRoper Gp
South Nicholson Gp Mullero F m
McArihur Gp
Bungle Bungle Dolomite
Balbirini Dolomite Barney Creek Fm Cooley Dolomite Emmerugga Dolomiie Amelia Dolomite
Paradise Creek Fm
1700-
1800-
3 NABBERU BASIN
1900-
2000-
2100-
Duck Creek Dolomite
Brochmon Iron Frn
I
1?
2200-
2300-
i Frere Fm
Fortescw Gp
W
2400-
Walter, 1977). The interaction of grazing metazoans and benthic microbial communities plausibly accounts for the Late Proterozoic decline in the abundance and diversity of stromatolites (Garrett, 1970; Awramik, 1971).
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394 La Berge, G.L., 1973. Possible biological origin of Precambrian iron-formations. Econ. Geol., 68: 1098-1109. Lewis, J.D., Rosman, K.J.R. and De Laeter, J.R., 1975. The age and metamorphic effects of the Black Range dolerite dyke. West. Aust. Geol. Surv. Annu. Rep., 1974: 80-88. Licari, G.R. and Cloud, P.E., Jr., 1968. Reproductive structures and taxonomic affinities of some nannofossils from the Gunflint Iron Formation. Proc. Natl. Acad. Sci. U.S.A., 59: 1053-1060. Licari, G.R. and Cloud, P.E., Jr., 1972. Prokaryotic algae associated with Australian Proterozoic stromatolites. Proc. Natl. Acad. Sci. U.S.A., 69: 2500-2504. Licari, G.R., Cloud, P.E., Jr. and Smith, W.D., 1969. A new chroococcacean alga from the Proterozoic of Queensland. Proc. Natl. Acad. Sci. U.S.A., 62(1): 56-62, 4 figs. Logan, B.W., Hoffman, P. and Gebelein, C.D., 1974. Algal mats, cryptalgal fabrics and structures, Hamelin Pool, Western Australia. Am. Assoc. Pet. Geol. Mem., 13: 38-84. Major, R.B., 1974. The Punkerri Beds. Q. Geol. Notes Geol. Surv. South Ayst., 51: 2-5. Marjoribanks, R.W. and Black, L.P., 1974. Geology and geochronology of the Arunta Complex, North of Ormiston Gorge, central Australia. J. Geol. Soc. Aust., 21: 291299. Marshall, C.G.A., May, J.W. and Perrett, C.J., 1964. Fossil microorganisms: possible presence in Precambrian shield of Western Australia. Science, 144 : 290-292. McKirdy, D.M., 1976. Biochemical markers in stromatolites. In: M.R. Walter (Editor), Stromatolites. Amsterdam, Elsevier, pp. 163-191. Milton, D.J., 1966. Drifting organisms in the Precambrian sea. Science, 153: 293-294. M i r , M.D., 1974. Microfossils from the middle Precambrian McArthur Group, Northern Territory, Australia. Origins Life, 5: 105-118. Muir, M.D., 1976. The Amelia Dolomite microflora, McArthur Group, Northern Territory. Alcheringa, 1: 143-158. Muir, M.D., Hamilton, L.H., Grant, P.R. and Spicer, R.A., 1974a. A comparative study of modern and fossil microbes, using X-ray microanalysis and cathodoluminescence. In: T.A. Hall, P. Echlin and R. Kauffmann (Editors), Microprobe Analysis as Applied to Cells and Tissues. Academic Press, London, pp. 33-58. Muir, M.D., Spicer, R.A., Grant, P.R. and Giddens, R., 197413. X-ray microanalysis in the SEM for the determination of elements in modern and fossil microorganisms. 8th Int. Congr. Electron Microscopy Canberra, Vol. 2, pp. 104-105. Oehler, D.Z., 1976. Transmission electron microscopy of organic microfossils from the late Precambrian Bitter Springs Formation of Australia: techniques and survey of preserved ultrastructure. J. Paleontol., 50: 90-106. Oehler, D.Z., 1977. Pyrenoid-like structures in late Precambrian algae from the Bitter Springs Formation of Australia. J. Paeontol., 51 : 885-901. Oehler, D.Z., 1978. Microflora of the middle Proterozoic Balbirini Dolomite (Mc Arthur Group) of Australia. Alcheringa, 2: 269-309. Oehler, D.Z., Schopf, J.W. and Kvenvolden, K.A., 1972. Carbon isotopic studies of organic matter in Precambrian rocks. Science, 175: 1246-1248. Oehler, J.H., 1976. Hydrothermal crystallization of silica gel. Geol. Soc. Am. Bull., 87: 1143-1 15 2. Oehler, J.H., 1977. Microflora of the HYC Pyritic Shale Member of the Barney Creek Formation (McArthur Group), middle Proterozoic of northern Australia. Alcheringa, 1: 315-349. Oehler, J.H., Oehler, D.Z. and Muir, M.D., 1976. On the significance of tetrahedral tetrads of Precambrian algal cells. Origins Life, 7: 259-267. Page, R.W., Blake, D.H. and Mahon, M.W., 1976. Geochronology and related aspects of acid volcanics, associated granites and other Proterozoic rocks in the Granites-Tanami region, northwestern Australia. BMR J. Aust. Geol. Geophys., 1 : 1-13.
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Section B
SOUTHERN AFRICA
This Page Intentionally Left Blank
397
Chapter 7
STRUCTURAL FRAMEWORK D.R. HUNTER and D.A. PRETORIUS
A . STRUCTURAL PROVINCES
*
In South and Central Africa five ancient structural provinces constituting the cratonic nuclei are separated by mobile belts in which subsequent tectonothermal events have caused polyphase deformation and high-grade metamorphism (Fig. 7.1). The cratonic areas are composed of granitoid gneisses and greenstone belts constituting a shield overlain in places by Precambrian and Phanerozoic supracrustal rock sequences. The shields have been affected by several deformational events and regional plutonic activity before they became stable between -3.0 and -1.8 Ga, after which they have remained largely unaffected by subsequent tectonism. The intervening linear mobile belts mark the sites of repeated intense deformation and metamorphism. The origins of these linear zones are still imperfectly known but there is an increasing volume of evidence to suggest that they are ensialic in origin (Clifford, 1970; Shackleton, 1973; Mason, 1973; Kroner, 1977; Martin and Porada, 1977). Cratonic nuclei The Kaapvaal craton attained sufficient rigidity at -3.0 Ga t o sustain and preserve epicontinental sedimentary basins containing gently dipping, largely unmetamorphosed deposits. These sediments rest on a shield composed of various granitoids and gneisses with relicts of deformed greenstone belts (see Chapter 8). The age relationships between the greenstone belts and complexly deformed, grey tonalitic gneisses are still uncertain. Geochronological data have so far proved inadequate to resolve the problem. Mafic volcanics from the Onverwacht Group in the Barberton greenstone belt have yielded a Rb/Sr mineral isochron age of 3.5 ? 0.2 Ga (Jahn and Shih, 1974) whereas
* By D.R. Hunter.
398
Fig. 7.1. Simplified map of southern and central Africa showing t h e distribution of cratons and mobile belts. T h e Limpopo mobile belt is -2.6 Ga and older; t h e Ubendian (2) is -1.85 Ga; t h e Kibaran, Irumide, Kheis-Matsap, and Namaqua-Natal belts (3) are -1.1 Ga and t h e Damaran, Gariep, Malmesbury and Zambezi belts are -0.5 Ga. Solid black areas within t h e Damaran belt are inliers of older basement: A = Kafue anticline; B = Hook granite, C = Grootfontein inlier, D = Kamanjab inlier. T h e unlabelled area near t h e southern flank of t h e Damaran belt is t h e extensive inlier south of Windhoek.
a hornblende-biotite tonalite intrusion in the grey gneisses in Swaziland (both part of the Ancient Gneiss Complex) gives a Rb/Sr whole-rock age of 86 Ma (Davies and Allsopp, 1976). Diapiric plutons of tonalitic 3395 composition intruded the schist belts at -3.2 Ga. A widespread granitoid intrusive event at -3.1 Ga was followed by a long period of relative stability during which a succession of ensialic basins developed within the Kaapvaal province (Table 7.1). These basins were the sites of sedimentation and volcanism, the products of which were laid down uncon-
*
399 formably on erosion surfaces cutting deformed schist belts and the associated granitoid rocks. The depositional axes of the successive basins migrated from the southeast to the northwest accompanied by an increasing inundation of the granite-greenstone shield. The oldest basin, in which the Pongola Group accumulated between -3.1 and 2.9 Ga, is confined t o the southeastern corner of the Kaapvaal craton, and was succeeded by the Witwatersrand, Ventersdorp and Transvaal basins, volcanics in the uppermost group of the latter being dated at -2.2 Ga. Deposition in the Transvaal basin was followed by the emplacement of the Bushveld Complex, the mafic rocks of which have yielded a Rb/Sr age of 2095 ? 24 Ma (Hamilton, 1977). Emplacement of the Bushveld Complex appears to have caused perturbation of Rb/Sr and K/Ar systematics over a wide area of the Kaapvaal structural province. Sedimentological studies of these successive basins reveal that sediment was derived from the north or northeast. Unroofing of the Limpopo mobile belt played a significant role in supplying sediment during the later stages of the Transvaal sedimentation (Button, 1973). Deposition of the Waterberg Group was confined t o a more restricted basin under continental conditions in the northern parts of the Kaapvaal craton and overlapping onto the Limpopo mobile belt. The limits of the Kaapvaal structural province are not accurately known in the south and east where Mesozoic rocks obscure the Precambrian geology. To the north the boundaries of the province are not clearly defined due to a gradual increase in metamorphic grade and deformation into the Limpopo mobile belt (Graham, 1974). A similar gradation from the Rhodesian craton into the Limpopo belt is also reported (Coward and James, 1974; Coward et al., 1976; see Chapter 10). The Matsap beds overlying the Transvaal rocks are strongly deformed along the western margin of the craton and pass southwestwards with increasing grade of metamorphism into the Kheis Group which was formerly regarded as a much older rock sequence (Cornell, 1977). The Rhodesian craton consists primarily of a granite-greenstone terrane with younger Precambrian cover sequences preserved along the eastern (Umkondo) and northwestern margins (Piriwiri, Deweras, and Lomagundi). Recent geologic and geochronologic studies in Zimbabwe (see Chapter 8) have revealed a long history of crustal development extending back t o >3.5 Ga. There appears to have been a cyclical repetition of the development of gneisses, greenstone belts and both potassic and sodic granitic intrusions during the period between >3.5 and 2.6 Ga. The oldest granite-greenstone terrane is found in the southeastern part of the c p t o n and consists mainly of tonalitic gneisses with infolded greenstone belt remnants, the latter constituting, at least in part, the Sebakwian Group. The tonalitic gneisses and intrusive Mushandike granite have been dated at -3.5 Ga (see Table 8.111, Wilson, Chapter 8 for references). The main greenstone belts are younger. Wilson et al. (1978) have divided the younger greenstones into lower and
400 TABLE 7.1 Outline of principal lithologic units in cratons and mobile belts in Central and Southern Africa Ga
Anglo-Kasai
0.5
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pan-African tectogenesis
Damaran
Pan-African tectogenesis Nama Group
Damara Supergroup Klein Aub and Ghanzi F m s 1.0
NW Cape and southern SWA
Kuboos-Bremen igneous suite Pan-African tectogenesis Nama Group
Richtersveld igneous suite Gariep G r o u p and Malmesbury Formation
-_------
Bushimay/ Oendeolongo?
Doornpoort and Grauwater Formations
Extensive thermal event
Kunene Complex? (min. age )
Charnockites, enderbites, norites?
2.0
Ubendian tectogenesis
Pilanesberg intrusions
Namaqua granitoids
. . . . . . . . . . . . . . . . . . . . . . . . .
Widespread granitoid intrusions
T
4
Koras and Sinclair Groups
1.5
Kaapvaal
Huabian episode of granitoid intrusions
Vioolsdrif igneous suite Kheis and Namaqua supracrustals
______---__
1
WaterbergMatsap?
Mild reheating
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Huab, Abbabis and Marienhof Formations?
Bushveld Complex
401
~~~
Limpopo
Rhodesian craton
W and NW Zimbabwe
Zambia-Zaire
Malawi
Pan-African tectogenesis
Pan-African tectogenesis
Pan-African tectogenesis Songwe syenites
Thermal event
Upper Katangan (Kundelungu)
Sijarira and Tengwe River Groups
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . pegmatites
Lower Katangan (Roan, etc.) Kibaran tectogenesis
Lead ages (galena, Elbas Mine)
Irumide tectogenesis Mafingi Group
Muva sequence
Kibara Group (inc. Plateau) Waterberg Group
Dyke intrusion Umkondo Group
Mild rehe: ting?
Mashonaland doleri tes
Metamorphism
Shearing, uplift of belt?
Urungwe granodiorite
Granite (Kamativi area)
Nyika granite
Ubendian tectogenesis
Ubendian tectogenesis accompanied by granitoid intrusion
402 TABLE 7.1 (continued) Ga
Anglo-Kasai
Damaran
NW Cape and southern SWA
Lulua sedimen. tary/volcanic com p 1ex
Kaapvaal
Transvaal Supergroup
Luiza metasediments? Dibaya granite/migmatite
Skalkseput granite ?
Gaborone granite
Ventersdrop Supergroup Witwatersrand Supergroup Usushwana Complex
3.0
Draghoender granite
Pongola Group
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Kasai-Lomani gabbrolnorite and charnockite complex?
Granitoid intrusions
?
Tonalitic diapirism
Ancient Gneiss Complex-
Limpopo
Rhodesian craton
Schiel alkaline intrusion?
W and NW
Zambia-Zaire
Malawi
Zimbabwe
Lomagundi Group? Deweras Group? Kamativi, Tshontanda and Malaputese ForPiriwiri Group? mation (min. age ) -- - - - _ _ _ ---- - - . . . . . . . . Great “Dyke” Great “Dyke” and satellites and satellites Matok, Bulai Granitoid etc. granites intrusion Tonalitic diapirism Shamvian G r o u p ? Upper Greenstones
Rusizian metasedimen ts
Jemba River granulites
. . . . . . . . . . . . . . . . . .
Pre-Muvu gneisses
Mkushi granite complex?
?
Mashaba tonalite and gneisses Lower Greenstones?
Hook granite
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Charnockites and enderbites Mafic dykes Anorthosi tic and layered complexes Limpopo Group? Mont d’Or granite Mafic dykes
Macuville Group
Mushandike granite Older greenstones (including Sebakwian) and gneisses
404 upper groups, constituting an informal subdivision of the Bulawayan Group. The Upper Greenstones and the sedimentary Shamvaian Group rest on a basement of deformed and eroded Lower Greenstones, -3.5 Ga greenstones, the Mashaba-type tonalites (-3.0 Ga) and older tonalitic gneisses (-3.6 Ga). Radiometric dating of the Upper Greenstones and intrusive tonalites and granites show a cluster of ages around -2.7 Ga. The age of the Lower Greenstones is uncertain but field evidence clearly demonstrates that they were deformed and eroded prior t o the accumulation of the Upper Greenstones (see Wilson, Chapter 8). Wilson et al. (1978) have suggested that theMashabatype tonalites and associated gneisses together with the Lower Greenstones may form part of a third granite-greenstone terrane intermediate in age between those dated at -2.7 Ga and -3.5 Ga. The Great “Dyke” that bisects the Rhodesian craton was emplaced at 2514 t 1 6 Ma (Hamilton, 1977) and represents the last major plutonic event before the deposition of the Piriwiri, Deweras, Lomagundi and Umkondo Groups, now preserved only around the margins of the craton. The Rhodesian craton differs from the Kaapvaal craton in that younger thick ensialic sedimentary and volcanic sequences are preserved only in the east and northwest. It is noteworthy that while accumulation of sediments and volcanics in ensialic basins was taking place in the Kaapvaal craton, volcanism and granitoid intrusion of greenstone belt type continued in Zimbabwe, final stabilization there being attained some 400 Ma later than in the Kaapvaal craton. The arcuate belt of sedimentary rocks (the Piriwiri, Deweras and Lomagundi groups) in the northwestern portion of the Rhodesian craton were involved in the Pan-African tectonothermal event (-500 Ma) but south of latitude 1 7 ” s there is no evidence of their involvement in this event. Radiometric data (Chapter 9 ) have been inconclusive in attributing an absolute age t o these rocks. On lithologic grounds correlations with the Witwatersrand and Transvaal Supergroups have been proposed as well as with the Katangan of Zambia. In the east the Umkondo Group rests unconformably on the granite-greenstone shield and is relatively undeformed. When traced eastwards and northwards into Moqambique the sequence has been involved in the Pan-African deformation and metamorphism. Sedimentological studies suggest that the source area for the detritus that accumulated in the Umkondo basin lay to the west and south. The minimum age of this group is 1.8 Ga (Chapter 9). The geology of the small Zambian nucleus (Clifford, 1970), also known as the Bangweula Block, is imperfectly understood but appears t o be analogous t o a craton. It consists of a granitoid and gneissic basement much of which is overlain by the Plateau Series that is older than 1.8 Ga (Cahen and Snelling, 1966). The granitoids have not been dated radiometrically but they must be older than -2.0 Ga as they have been marginally affected by the Ubendian tectothermal event (Drysdall et al., 1972). Kroner (1977) has
-
40 5 remarked that the Zambian nucleus resembles a “cratonic island” within younger mobile belts. Drysdall et al. (1972) suggest that it has been largely unaffected by orogenic deformation since Archaean times. Another small segment of cratonic basement is exposed around Rufunsa lying between the Zambian craton in the north and the Rhodesian craton t o the south (Fig. 7.1). De Swardt et al. (1965) proposed a correlation of metavolcanic and metasedimentary rocks (the Mpanshya Group) and associated granitoids with the Rhodesian granite-greenstone terrane but Barr (1972) considered the Mpanshya Group t o be younger. The Rufunsa segment displays considerable Irumide overprinting but again radiometric data are lacking. The Tanzania craton in the extreme north consists of granitoids and acid gneisses with deformed and metamorphosed Dodoman greenstone relics, the last metamorphism of which is probably -2.6 Ga (Wendt et al., 1972). Nyanzian greenstones in northern Tanzania have a pre-2.8 Ga age based on the Rb/Sr isochron age of 2800 2 100 Ma for the post-Nyanzian Migori granite (Dodson et al., 1975). The Nyanzian greenstones are overlain by the arenaceous and argillaceous sediments and volcanics of the Kavirondian Group that is older than 2.5 Ga (Clifford, 1970). These radiometric ages from the Tanzania craton suggest that it is a very ancient cratonic nucleus similar t o the Rhodesian and Kaapvaal cratons. The recognition of reworked cratonic material within the adjacent Pan-African RJoqambique belt suggests that the Tanzania craton was formerly more extensive (Hepworth, 1972). The Angola-Kasai craton is less easy t o define due t o the presence of a blanket of cover rocks. An ill-defined Kasai nucleus comprises granitoid gneisses and pegmatites, microcline from the latter yielding ages of -3.5 Ga (Cahen et al., 1976). It seems probable that this Kasai nucleus may have been more extensive before being reworked by later tectonothermal events, the most recent being -2.0 Ga in Angola.
Mobile belts The east-trending Limpopo belt separates the Rhodesian and Kaapvaal cratons. Cox et al. (1965) recognized three zones within the belt. A central zone contains large-scale interference structures in which fold axes are disposed in general northeast and northwest and consists of gneisses and medium- to high-grade metasediments and metavolcanics. The zone is bounded by northern and southern marginal zones, characterized by granulite facies metamorphism and by the generally consistent alignment of the foliation of the gneisses, a feature that has been attributed t o transcurrent movement (Mason, 1973). Subsequent studies have shown that, although the essential symmetry of the belt is valid, the boundaries with the Rhodesian craton t o the north and the Kaapvaal craton to the south are diffuse. Deformation characteristic of the northern marginal zone extends into the Rhodesian craton as far north as Bulawayo (Coward et al., 1976). Similarly Graham
(1974) found evidence that the deformation in the southern marginal zone extends as far south as the Murchison greenstone belt (see Chapter 8). In the absence of a well-defined structural break, the orthopyroxene isograd has been used t o define the northern and southern margins of the belt (see Chapter 10). In the east the Limpopo belt is blanketed by Mesozoic rocks so that its relationship t o the Pan-African Moqambique belt is not known. Recent detailed work in Botswana suggests that the Limpopo belt does not continue westwards under the cover of younger rocks for any great distance, but dies out at about longitude 27”E (Key and Hutton, 1976). These authors relate the interference fold structures in the Limpopo belt of northeastern Botswana to shear movement along the “Tuli-Sabi Straightening Zone”, which weakens westward, and hence conclude that the Limpopo mobile belt must also die out in this direction (but see Chapter 10). Previous investigators have proposed a contemporaneous north-northwesterly aligned offshoot of the Limpopo belt (the Shashe mobile belt) around the southwestern margin of the Rhodesian craton, but Key and Hutton (1976) question the validity of this concept. It seems probable that the so-called Shashe mobile belt is another example of deformation extending into the cratonic domain. There is a gradual change from the low-grade granite-reenstone terrane of the cratons into the high-grade Limpopo belt gneisses some of which are regarded as the equivalents of greenstone lithologies in the cratonic domain. Recent metamorphic, structural, and radiometric studies indicate that the Limpopo belt was the site of a long-continuing sequence of tectonic and/or thermal events. Barton et al. (1977) have identified the presence of gneisses yielding Rb/Sr dates of 3.8 Ga followed by the intrusion of mafic dykes at 3.6 Ga. Metamorphic and intrusive events at -3.2 Ga followed the accumulation of the Limpopo Group. Further metamorphic events at -2.9 and 2.7 Ga have also been provisionally recognized (Table 7.1 and see Chapter 10). The last major tectonothermal event certainly predates the intrusion of the 2.5 Ga-old Great “Dyke” and its satellites which cut, and have chilled margins against the high-grade gneisses. Mineral ages of -2.0 Ga may indicate either (1)a stage of uplift and cooling of the Limpopo belt or (2) mild reheating associated with the intrusion of the Bushveld Complex in the Kaapvaal province or (3) a series of deformational events restricted t o the eastern end of the belt as suggested by Hickman and Wakefield (1975). The intrusion of the Bushveld Complex was contemporaneous with deformational events in the Ubendian mobile belt, which separates the Zambian and Tanzania cratons and is part of a continent-wide tectonic event at 1.8 and 2.0 Ga (Clifford, 1970). Concordia intercept ages of 1730 f 30 Ma and 1870 f 30 Ma have been yielded by zircons from the Fransfontein granitic suite that forms part of the Kamanjab inlier in the Damaran belt (see Fig. 7.1). Ubendian ages have also been reported from the Namaque belt further t o the south.
407
The relationship of t h e Ubendian belt to its adjacent cratonic domains is largely obscured by a cover of younger rocks. Recent investigations have shown that although the high-grade gneisses and granulites in northern Malawi are situated within the Moqambique belt, they are of Ubendian age. Furthermore, the granulite and associated charnockites were formed during an older event contemporaneous with granulite metamorphism of the Limpopo belt (-2.7 Ga). Within the Ubendian belt the older basement together with supracrustal pelitic, calcareous and psammitic sediments was deformed at -2.0 Ga, at which time synkinematic granitoids, presumably resulting from partial melting of the salic basement, were intruded. The Irumide belt trends north-northeastwards from southern Zambia t o northern Malawi transecting the older Ubendian belt. The validity of a discrete Irumide fold belt in southern Zambia has been questioned (Ramsay and Ridgway, 1977). These authors find little evidence for the existence of an Irumide orogeny of Kibaran age (-1.0-1.3 Ga). However, Ray concluded that deformation at 1.1Ga produced north-trending structures in northern Malawi that were later reorientated by Pan-African movements. Metasediments of the Mafingi and Muva sequences (quartzites, pelites and conglomeratic beds) rest unconformably on older gneisses that suffered cataclastesis during the Irumide folding of their cover rocks about north-striking axial planes. In western Zimbabwe (see Lockett, Chapter 1 0 ) four northeasterly trending belts of deformed supracrustal rocks suffered early high-grade (granulite facies) and later retrograde metamorphism. The earlier event was accompanied by intense deformation producing northeast trending isoclinal folds. The retrograde metamorphism is associated with a later phase of folding coaxial with the early folds. Radiometric data suggest that the highgrade metamorphic event is -2.0 Ga and was followed by a thermal event at 1.O Ga when mineralized pegmatites were generated. Lead ages of 1.3 Ga led Cahen and Snelling (1966) t o propose that the Irumide belt extends southwestwards along the frontier between Zambia and Zimbabwe. The twopyroxene granulites in the belt defined by Ramsay and Ridgway (1977) extending from the Zambia/Zimbabwe border near Kariba northeastwards into Zambia are considered by these authors to be possibly of Pan-African age. It is clear that the identification of a discrete Irumide belt is subject t o considerable uncertainty, although a thermal event at -1.0 Ga has apparently affected areas such as western Zimbabwe. Intense overprinting by t i e Pan-African event in southeastern Zambia contributes to the difficulty of defining the deformational history of this area. The 1.0 to 1.3 Ga Kibaran event affected extensive areas of the African continent. In central and southern Africa metamorphism and deformation of this age is preserved within the Kibaran belt of eastern Zaire and in the Namaqua-Natal belt of South West Africa/Namibia and South Africa. The Kibaran belt extends southwestwards from East Africa into Zai’re
-
-
-
408 before being covered by the Katangan sequences of Zambia which were subsequently involved in the Pan-African tectonothermal event. Some 10,000 m of clastic and chemical sediments, deposited on an eroded basement of rocks, dated at -2.0 Ga and older, constitute the Kibara Group and its correlatives in East and Central Africa. Thick sequences of sediments and basic t o acid volcanics, later subjected t o low-grade metamorphism, were deposited in several small basins a t 1.3 to 1.7 Ga and are now preserved in an series of basement inliers within and adjacent t o the Damaran belt of Namibia/South West Africa (Schalk, 1973). Evidence that they suffered tectonism is provided by mineral ages ranging from 1.3 t o 1.1Ga. A connection between the Kibaran and Namaqua belts of South Africa and South West Africa/Namibia was first suggested by Holmes (1951). Radiometric studies of the Namaqua belt have yielded mineral ages in the range 1.1 t o 1.25 Ga (Nicolaysen and Burger, 1965), but no thick sedimentary/volcanic sequences equivalent in age t o the Kibara Group have been positively identified in the Namaqua belt. There are, however, metasedimentary and metavolcanic rocks within the Namaqua gneissic terrane which on structural and lithologic grounds have been correlated with the Kheis Group. Formerly regarded as older than 2.9 Ga (Vajner, 1974) recent radiometric dating suggests that the Kheis Group may be -1.9 Ga (Cornell, 1977) and approximately equivalent in age t o the Matsap rocks, dated at 2.0 Ga (Crampton, 1974), that crop out along the western margin of the Kaapvaal craton. These rocks show an increasing degree of metamorphism when traced further southwestwards. Ages of 1.8 t o 2.0 Ga have also been reported from the Haib-Vioolsdrif metavolcanics and granites in the Namaqua belt near the Atlantic coast (Reid, 1976; and see Fig. 7.1). The Kibaran event in the Namaqua belt (see Joubert, Chapter 10) caused extensive reheating and was accompanied by granite intrusion and regional anatexis (Joubert, 1971; Jackson, 1976). Uplift of the belt has exposed high-grade metamorphic rocks (Pretorius, 1974; Kroner, 1977). In Natal Province of South Africa ages around 1.0 Ga have been reported (Nicolaysen and Burger, 1965) but a possible connection of this zone with the Namaqua belt is obscured by a cover of younger rocks. It is not yet possible to establish precise chronologic relationships. The most extensive outcrops consist of amphibolites, hornblende gneisses and various granitic gneisses. Granulites and charnockites are present in the southernmost exposures. At the northern margin of the belt t h e gneisses and amphibolites have been tectonically intersliced and thrust over similarly deformed Mfongosi metasupracrustals (see Chapter 10). That the Pan-African event affected a major part of t h e African continent is reflected by widespread radiometric ages within the range 0.75 t o 0.45 Ga. Clifford (1970) recognized two patterns of zones within the Pan-African belts :
409 (a) zones of orogenically deformed Upper Precambrian sedimentary sequences, e.g., the Katangan Supergroup, The Damaran Supergroup, the Gariep Group and the Malmesbury Formation; (b) zones of rejuvenated basement e.g., the Moqambique and Zambesi belts of Central and East Africa and the Malagasy Republic. Clifford (1970) interpreted the similarity of the age patterns, the geographical and structural continuity of the belts and the presence or absence of a thick sedimentary sequence as a reflection of exposure of different structural levels through a sinuous mountain chain. The Katangan rock sequence and its correlatives consist largely of arkoses, sandstones, shales, conglomerates and chemical sediments which are strongly folded within the confines of the Pan-African tectonothermal belts in contrast t o their flat-lying disposition on adjacent stable cratonic areas. There is a gradual southward increase in metamorphic grade and in abundance of granitic intrusions across the Pan-African belt in Zambia. Similarly in northwestern Zimbabwe there is progressive increase in regional metamorphic grade from south t o north across the Pan-African belt. The Pan-African belt in South West Africa/Namibia and its extensions into South Africa consist of two branches, the best preserved being the intracratonic branch, about 400 km wide, and probably connected beneath the cover of younger rocks with the Zambesi and Moqambique belts. The second branch trends subparallel to the present coast of southern Africa probably linking the Gariep and Malmesbury belts (Martin and Porada, 1977). The intracratonic branch contains deformed sediments, at least 10 km thick, in which can be distinguished a “mioge~synclinal’~ facies of low metamorphic grade and a “eugeosynclinal” facies that has been subjected t o high-grade metamorphism, anatexis and granitoid plutonism in the central zone of the fold belt (Martin and Porada, 1977). A marked structural asymmetry results from the progressive increase in deformation towards the southern margin, culminating in nappe and thrust tectonics (Martin and Porada, 1977; Bickle and Coward, 1977). The Gariep fold belt contains sediments that display rapid facies changes. A northward trending line of subvolcanic complexes (the Richtersveld Suite) of granitic t o syenitic composition dated at -0.9 Ga (Welke et al., 1973) seems to have intruded the hinge zone of the downwarp in which the Gariep sediments accumulated (Martin and Porada, 1977). The Gariep belt is distinguished from the main branch of the Damaran belt by the presence of the northeast-trending Kuboos-Bremen line of plutons dated at -0.5 Ga (Welke et al., 1973; Martin and Porada, 1977) and compositionally similar to the Richtersveld Suite. Further south in the Cape Province of South Africa, evidence that the Pan-African event affected the southernmost parts of the continent is found in the deformation of the Nama Group and the Malmesbury and Cango Formations. The Cape Granites intrude the Malmesbury Formation and are
410 dated at between 0.6 and 0.5 Ga. It has been assumed that all the Cape Granites are of similar age but it has been suggested that some of the granites may have formed the floor on which the Malmesbury Formation was deposited (Hartnady, 1969). The Zambesi and Moqambique belts represent examples of Pan-African rejuvenation of older crustal rocks. No supracrustal rocks reflecting sedimentation immediately preceding Pan-African tectogenesis are known. The Umkondo Group of Eastern Zimbabwe has a minimum age of 1785 f 80 Ma (Vail and Dodson, 1969). Thus it does not represent a period of sedimentation of Pan-African age although it is involved in deformation associated with the Moqambique and Zambesi belts and is pierced by mantled gneiss domes (Talbot, 1973). Archaean greenstones of the Rhodesian craton can be traced into the Moqambique belt through a transition zone of increasing tectonic and metamorphic overprinting (Johnson and Vail, 1965). Three structural domains have been identified (Johnson and Vail, 1965): (1)discrete non-penetration shear zones; (2) penetration shearing and shear-folding; (3) plastic isoclinal folding within which domain mantled gneiss domes have risen. The origins of the various mobile belts and their possible evolution with time remain problematical. The ensialic nature of the Precambrian mobile belts in south and central Africa has been noted by many authors (e.g., Clifford, 1970; Shackleton, 1976; Kroner, 1977; Martin and Porada, 1977). The available data suggest that large continental plates have existed since at least the earliest Proterozoic and have been rejuvenated and reworked at various times in linear mobile belts. The resultant pattern of mobile belts and cratons “gives the impression of younger orogens surrounding ancient nuclei” (Kroner, 1977). A plate tectonic model for the evolution of the southern African mobile belts is supported by the presence of: (a) triple junctions of mobile belts; (b) widespread granulite facies rocks in Pan-African belts suggestive of crustal thickening; (c) marginal positive gravity anomalies; (d) the isolated occurrence of glaucophane in the Gariep Group (Kroner, 1974); (e) a calc-alkaline magmatic arc south of the Damara belt (Watters, 1976); (f) a thrust complex at the northern margin of the Natal belt interpreted by Matthews (1972) as an obduction zone. However, there is other substantial evidence that cannot be reconciled with a plate-tectonic model: (1)palaeomagnetic data provide no evidence of significant relative motion of the cratonic areas through much of Precambrian time (Briden, 1976); (2) the absence of convincing ophiolites; (3) the continuity of earlier structures and lithologies across craton/mo-
411 bile belt boundaries (e.g., Shackleton, 1973). Two provisional hypotheses have been considered t o account for Precambrian crustal mobility: (a) limited relative motion between stable blocks within continental plates resulting in the formation of short-lived oceanic basins and, ultimately, mobile belts; and (b) intracrustal vertical movements induced by mantle upwelling causing weakening of the overlying crust, igneous activity and the development of mobile belts. Combination of these hypotheses has led t o the suggestion that African mobile belts are “either derived from aulacogens as a result of internal plate distortion and/or through taphrogenesis in Rhine- or East African-type graben systems” (Kroner, 1977). Whatever the causes of Precambrian crustal mobility, the occurrence of ancient reworked sialic crust within these linear belts seems t o support the concept that a major part of the continental crust formed at an early stage of the Earth’s evolution. B. ANALYSIS OF GRAVITY DATA
*
The regional gravity field over South Africa and South West Africa Although gravimetric methods of geophysical surveying were first systematically employed in mineral exploration in Southern Africa in the late 1930’s, their application was essentially t o local problems. The first regional gravity survey of the sub-continent was completed by the Geological Survey of South Africa in the second half of the 1950’s, and the results were published in 1962 (Smit, 1962). The next regional investigation was undertaken in South West Africa (Kleywegt, 1966). In subsequent years, smaller areas were studied, covering the whole of Swaziland and portions of Botswana, Zimbabwe, and Moqambique. In the mid-l97O’s, there was a resurgence of interest in obtaining total coverage of the gravity field over Southern Africa, and, subsequent to the compilation of the information described here, regional gravimetric data became available for the whole of Zambia, Malawi, and Botswana. The countries of the sub-continent, for which adequate knowledge of the gravity field is still lacking, are Zimbabwe, Angola, Moqambique, and Lesotho. For almost twenty years after the gravimetric information for South Africa was released, little thought was given to the significance of the regional component of the total variation in the field. Contoured maps of the Bouguer and isostatic anomalies were used primarily in the structural analysis of small areas of interest t o localized mineral exploration activities.
* By D.A. Pretorius.
412
Fig. 7.2. A moving-average, 2' X 2 O (latitude and longitude) surface of t h e regional component of t h e Bouguer anomaly field over South Africa and t h e southern part o f South West Africa.
The advent, in the late 1960's and, particularly, the early 19703, of computer-based, sophisticated methods of filtering areally distributed, quantitative data permitted, for the first time, the systematic differentiation of the total field into regional, or trend, components and local, or residual, components. The results described in this contribution t o the structural framework of Southern Africa were obtained through statistical techniques embracing polynomial and moving-average trend-surface analysis, In general, it was found that the main features of the tectonic fabric were more sharply defined through moving-average methods, employing blocks of data encompassing areas measuring two degrees of latitude by two degrees of longitude, than through polynomial trend surfaces of up to the sixth order. Such a moving-average surface is depicted in Fig. 7.2, which includes gravity observations over the whole of South Africa and the southern portion of South West Africa. It can be seen that the regional gravity field over South Africa has a roughly ovoidal shape, with the major axes trending WSW-ENE, from the Cape of Good Hope to the South Africa-Moqambique border, between Swaziland and Zimbabwe, and SSE-NNW, from Lesotho to the central part of Botswana. The field reaches its lowest intensity over Lesotho
4 13
which also represents the topographically highest area in Southern Africa, indicating that the east-central part of South Africa is still undergoing isostatic adjustment. Two minima - the Maseru and Vredefort gravity lows appear t o be off-set portions of the Kaapvaal low which is believed t o be a function of a comparatively thicker sialic crust, centred under Lesotho. The assumed thinning of such a crust is more gradual towards the southwest continental edge than towards t h e northeastern and southeastern edges, the gravity gradient in the last-mentioned direction being particularly steep. In the southeastern portion of South Africa, the coastal effect on the gravity field extends 1 5 0 km inland, while, in Moqambique, the effect is still discernible 250 km inboard from t h e coast. Oceanic crust and the upper mantle might thus be at a shallower depth under Moqambique than under South Africa. Two small, but conspicuous, gravity highs are detectable in the regional gravity field on the southwestern side of the dominating Kaapvaal gravity low. The Trompsburg high is known t o be associated with a major, differentiated, Late Proterozoic igneous complex of circular shape, concealed beneath a thick cover of Phanerozoic rocks. The geological feature responsible for the Strondkop high has not yet been clearly established, but it would seem likely that it might also be indicative of a concealed igneous intrusion of major dimensions, possibly blanketed by Late Proterozoic strata. The field over the southern part of South West Africa points t o the presence of the Suidwes gravity low in the central part of the border area between that country and Botswana. Because of the absence of adequate data for the latter territory, at the time of preparing this contribution, it was not possible t o determine whether the Suidwes gravity low could be regarded as a westwards displacement of the Kaapvaal gravity low, o r whether it constituted a discrete crustal feature, unrelated t o the areas of thicker sialic crust in South Africa. It is considered that the latter is the more likely possibility, indicating the presence of at least two focal points of crustal thickening in the region covered by Fig. 7.2. The main elements in the tectonic fabric, as revealed by the regional component of the gravity field, are shown in Fig. 7.3. Two axial directions of minimum gravity are coincident with two axes of crustal upwarping. Whether the upwarping is a product of a thicker sialic crust along the central portion of the Kaapvaal crustal fragment o r whether more sial development where upwarping and updoming, on a sub-continental scale, took place remains t o be ascertained. On all scales of folding, these two axial directions can be detected as having controlled the geometry of deformation. The configuration of outcrops and suboutcrops of strata of varying ages reflects the “vosfi”-pattern - the Vaal-Orange-superimposed-fold-interference pattern, with the Vaal trend being that orientated WSW-ENE and the Orange trend NNW-SSE. Regional plunges, on the sub-continental scale, are away from the Maseru and Vredefort gravity lows, towards the coastlines and under the
414
Fig. 7.3. The trends o f the major axes of arching and of the dominant lineaments revealed by the configuration of the contours of the regional component of the Bouguer anomaly field over South Africa.
Kalahari area of Central Botswana. The dominant fracture lineaments are parallel, or sub-parallel, to the Vaal and Orange trends. Elongation of the gravity highs at Trompsburg and Strondkop is in the NNW direction, providing evidence that the localization of emplacement of prominent igneous complexes is lineament-controlled. On the Trompsburg lineament, in addition to the Late Proterozoic intrusion, a number of significant diamond-bearing kimberlites, of possible Cretaceous age, have been injected, pointing to longevity of activity of, at least, some of the major Vaal- and Orange-trend fractures. That the ENE lineaments might have acted as transcurrent faults, with substantial horizontal movements, is suggested by the apparent right-lateral displacement of the Vredefort and Maseru portions of the Kaapvaal gravity low. The projection of this fracture to the northeast coincides with what might be a right-lateral displacement of the coastline of Moqambique. The geometry of the regional gravity field over South Africa is a largescale portrayal of the areal and local tectonic fabrics, favouring the concept that the parts of the Kaapvaal crustal fragment have not reacted t o the processes of deformation in a style different to that which shaped the gross structural morphology of the whole. The principal fold and fracture trends
415 follow the same two general directions, the interaction of which has produced an interference pattern which is the key t o understanding the configuration and areal distribution of the various sequences which make up the total stratigraphic assemblage of South Africa. The local, residual components of the gravity field show a spatial array of parallel or sub-parallel zones of alternating high and low gravity values which, for the most part, can be seen t o be coincident with synclines and anticlines, respectively. The residual trends on the Kaapvaal crustal fragment are aligned roughly west-east in the eastern portion of the fragment and n o r t h s o u t h in the western segment, thus defining a major regional flexuring over the Kaapvaal gravity low. This low reflects crustal upwarping and thickening of the sial along the Molomari arch, a NNW-trending structural feature which dominates the fabric of the Kaapvaal crustal fragment (Fig. 7.4).The pattern of the array is disturbed in the north, where the Limpopo gravity high is developed, and in the southwest, where the Karasburg gravity high is present. In the southeast, the possible presence of another gravity high is
Fig. 7 . 4 . The trends of the axes of zones of higher and lower Bouguer anomalies, revealed by the configuration of the contours of the residual component of the Bouguer anomaly field over South Africa and South West Africa, and the location and partial extent of four crustal fragments (cratons), which are the main elements of the structural framework, and the mobile belts separating them.
416 masked by the steep gradients of the coastal effect. At the time of the investigation, the absence of gravity data for Botswana and Moprnbique did not permit any conclusions regarding the nature of the gravity field at the northwestern and eastern extremities of the Kaapvaal crustal fragment. In South West Africa, two different patterns of gravity residual alignments are apparent. In the south, there is a concentric arrangement of alternating highs and lows about the Suidwes regional gravity low. In the north, a second array would seem t o be related to what, if gravity data were available for Angola, might emerge as the Ondasha regional gravity low. Changes in the direction of concavity of the curvilinear residual trends are strongly suggestive of the presence of different focal points of concentric configurations of alternating highs and lows on areal and local scales. This concentricity of gravity trends about regional gravity lows, also seen, partially, in the case of the Kaapvaal crustal fragment, and the well-established younging of successive Proterozoic stratigraphic sequences, radially away from the central zones of the regional gravity lows, are evidence of these lows representing foci of maximum crustal updoming and maximum sialic thickening, The two crustal fragments in South West Africa are separated by the Spitzkoppe gravity high. In summary, the regional gravity field over South Africa and South West Africa reveals that this part of the sub-continent is built up of at least four crustal fragments, generally ovoidal in shape, which are separated from each other by regional gravity highs that flank the fragments concentrically on all sides. The fragments, more-or-less the equivalents of what have been called cratons previously, are characterized by a central portion of gravity minima which are taken as indicating the presence of relatively thicker sialic crust under areas of maximum tectonic upwarping and doming. The intervening zones of high gravity have been labelled as mobile belts which have hallmarks of high-grade metamorphism and plastic deformation. The increase in the gravity field suggests an attenuation of the sialic crust beneath the mobile belts, with associated downwarping and rift-faulting. The relatively narrow widths of the regional gravity highs between the crustal fragments renders it unlikely that the whole of the western segment of South Africa, southwest of the Kaapvaal fragment, is a mobile belt - the NamaqualandBushmanland metamorphic terrane - as is presently accepted to be the case. The gravity field would preferentially favour the presence of another crustal fragment - the Sanama fragment - in the northwestern Cape Province, separated from the Suidwes gravity low by the Karasburg gravity high and from the Kaapvaal gravity low by a similar, narrow, mobile-belt-gravity high which has yet t o be defined and named. Fig. 7.4 shows that the central portion of Botswana, beneath the Kalahari sand-cover, could be a most interesting geological region in that it acts as host to the conjunction of four cratons - t h e Kaapvaal, the Rhodesia, the Suidwes, and the Ondasha crustal fragments.
417 The regional gravity field over Southern and Central Africa Despite the fact that, at the time the present contribution was compiled, there were only scattered data available on the nature of the gravity field in the rest of Southern Africa, beyond the boundaries of South Africa and South West Africa, an attempt was made t o extrapolate the trend-surface analyses into Botswana, Zimbabwe, and Zambia. The results are shown in Fig. 7.5. The portion of the sub-continent south of Angola-Zaire-Tanzania shows the presence of three first-order gravity lows - Luwama in Zambia, Suidwes in South West Africa, and Kaapvaal in South Africa.’Higher regional gravity values, on average, make the other known crustal fragments - Rhodesia in Zimbabwe, Ondasha in South West Africa, and Sanama in South Africa - of second rank, such a classification t o be regarded as of a tentative nature only, in view of the paucity of data considered outside of South Africa and South West Africa. Even on the very broad scale depicted in Fig. 7.5, it is apparent that the configuration of contours clearly shows zones of relatively high gravity values wrapping themselves round the peripheries of the crustal fragments-cratons defined by low regional gravity values. A profile has been drawn of the gravity field from the Tanzanian-Zambian border t o the southern coast of South Africa (Fig. 7.6). This shows that the gravity highs-mobile belts are coincident with the more pronounced drainage features intersected along the line of section - the Lake Rukwa rift valley, the Zambesi River valley, and the Limpopo River valley. The gravity minima occur over the most topographically elevated sections of the gravity lows-craton - the Mbala-Kasama high country of the Luwama crustal fragment, the Salisbury-Fort Victoria high country of the Rhodesia fragment, and the Lesotho high country of the Kaapvaal fragment. There is thus a good degree of correlation between regional gravity lows, the central portions of cratons, thick developments of sialic crust, maximum uparching and doming, and topographically elevated country, as well as between regional gravity highs, the peripheral portions of cratons, the inter-cratonic areas, attenuations of the sialic crust, maximum downwarping and rift-faulting, and topographically depressed country. The order of values in the various regional gravity lows might be a measure of the status of isostatic imbalance and the extent of vertical uplift and consequent erosion of cover-rocks of each of the crustal fragments. If this is the case, then, of the three first-rank crustal fragments, the Luwama craton is the one showing the youngest onset of uplift and the least amount of resultant erosion. The Suidwes craton would have suffered the greatest amount of uplift and erosion, while the Kaapvaal craton would be intermediate between the Luwama and Suidwes fragments. The second-rank cratons - Rhodesia, Ondasha, and Sanama - would all have undergone greater vertical elevation and erosion of higher-level crustal material, t o account for the
418
Fig. 7 . 5 . A sextic polynomial trend surface o f the Bouguer anomalies over central and southern Africa, showing the location of the three first-rank crustal fragments (cratons) in the structural framework and o f the line o f profile depicted in Fig. 7.6.
419
t,
,,
t Rukwa
IMbala I
I
Kasama
-
,
ZAMBIA
T
ZambesiR IMI Darwin I Umnlali R Fort Victorla BubiR llLouis Trichardt I Middelburq I Salisbury BeitlBridge Dlifants R I LimpapolR
I I
I
TANZANIA
1
I
Mpika Luanqwa R Chipatal
MOCAMBIOUE
RHODESIA
SOUTH AFRICA
\ I
50
\ \
RHODESIA
0
500
.
Coast SOUTH AFRICA
Kilometres
Fig. 7.6. A profile o f the regional component o f the Bouguer-anomaly field along a line extending from t h e Tanzania-Zambia border t o the southern coast o f South Africa.
possibly thinner sialic crust under their central portions, as suggested by the higher average regional gravity values than are present over the Kaapvaal, Suidwes, and Luwama crustal fragments. The analysis of the gravity data over Southern Africa leads to a conclusion - tentative until such time as the newly available data from Botswana, Zambia, and Malawi have been processed - that the structural framework of the sub-continent might be compared to an inverted egg-tray. The receptacles for the eggs would become the quasi-circular, updomed, crustal fragments, or cratons, underlain by thicker sialic crust, while the divisions between the receptacles would become the narrow, concentric, mobile belts, underlain by attenuated sialic crust. Such a regional pattern t o the framework would be analogous to that seen, on an areal scale, in classic Archaean granite-greenstone terranes, with diapiric granitic plutons encircled and separated by greenstone belts. Sub-continental, regional, areal, and local structural fabrics all seem t o reflect the same pattern. REFERENCES Part A Barr, M.C.W., 1972. Short report o n the Mpanshya Group o f central Zambia. Univ. Leeds Res. Inst. Afr. Geol. Annu. Rep., 16: 3-7.
420 Barton, Jr., J.M., Fripp, R.E.P. and Ryan, B., 1 9 7 7 . R b / S r ages and geological setting of ancient dykes in t h e Sand River area, Limpopo Mobile Belt, S o u t h Africa. Nature, 267: 487-490. Bickle, M.J. and Coward, M.P., 1977. A major thrust in t h e southern Damara belt, Namibia. Univ. Leeds Res. Inst. Afr. Geol. Annu. Rep., 20: 8-13. Briden, J.G., 1976. Applications o f palaeomagnetism t o Proterozoic tectonics. Philos. Trans. R. SOC.London, Ser. A, 2 8 0 : 405-416. Button, A., 1973. A Regional S t u d y of t h e Stratigraphy and Development o f t h e Transvaal Basin in t h e Eastern and Northern Transvaal. Ph.D. Thesis Univ. Witwatersrand, Johannesburg (Unpubl.). Cahen, L. and Snelling, N.J., 1 9 6 6 . T h e Geochronology of Equatorial Africa. NorthHolland Publ. Co., Amsterdam, 1 9 5 pp. Cahen, L., Delhal, J. and Lavreau, J., 1976. T h e Archaean of Equatorial Africa, a review. I n : B.F. Windley (Editor), T h e Early History o f t h e Earth. Wiley, London, pp. 489498. Clifford, T.N., 1970. T h e structural framework of Africa. In: T.N. Clifford and I. G a s (Editors), African Magmatism and Tectonics. Oliver and Boyd, Edinburgh, pp. 1-26. Cornell, D.H., 1 9 7 7 . A post-Transvaal age for t h e Marydale Formation, Kheis Group, southern Africa. Earth Planet. Sci. Lett., 3 7 : 117-123. Coward, M.P. and James, P.R., 1 9 7 4 . Deformation of t w o Archaean greenstone belts in Rhodesia and Botswana. Precambrian Res., 1: 235-258. Coward, M.P., James, P.R. and Wright, L., 1 9 7 6 . Northern margin of t h e Limpopo mobile belt, southern Africa. Geol. SOC.Am. Bull., 8 7 : 6 0 1 4 1 1 . Cox, K.G., Johnson, R.L., Monkman, L.J., Stillman, C.J., Vail, J.R. and Wood, D.N., 1 9 6 5 . T h e geology of t h e Naunetsi igneous province. Philos. Trans. R. SOC. London, Ser. A , 2 5 7 : 71-218. Crampton, D., 1 9 7 4 . A n o t e o n t h e age of t h e Matsap Formation of t h e northern Cape Province. Trans. Geol. SOC.S. Afr., 77 : 71-72. Davies, R.D. and Allsopp, H., 1 9 7 6 . Strontium isotopic evidence relating t o t h e evolution of t h e lower Precambrian crust in Swaziland. Geology, 4 : 553-556. De Swardt, A , , Garrard, P. and Simpson, J.G., 1 9 6 5 . Major zones of transcurrent dislocation and superposition of orogenic belts in part of central Africa. Geol. SOC.Am. Bull., 76: 89-102. Dodson, M.H., Cavanagh, B.J., Thatcher, E.C. and Aftalion, M., 1 9 7 5 . Age limits for t h e Ubendian metamorphic episode in northern Malawi. Geol. Mag., 1 1 2 : 403-410. Drysdall, A.R., Johnson, R.L., Moore, T.A. and Thieme, J.G., 1 9 7 2 . Outline of t h e geology of Zambia. Geol. Mijnbouw, 5 1 : 265-276. Graham, R.H., 1 9 7 4 . A structural interpretation o f t h e southern margin of t h e Limpopo belt, northern Transvaal. Univ. Leeds Res. Inst. Afr. Geol. Annu. Rep., 18: 63-69. Hamilton, P.J., 1 9 7 7 . Sr isotope and trace element studies of t h e Great D y k e and Bushveld Mafic Phase and their relation t o early Proterozoic magma genesis in southern Africa. J. Petrol., 18: 24-52. Hartnady, C.J.H., 1 9 6 9 . Structural analysis of some pre-Cape formations in t h e Western Province. Precambrian Res. Unit, Univ. Cape T o w n , Bull., 6 : 7 0 pp. Hawkesworth, C.J., Moorbath, S., O’Nions, R.K. and Wilson, J.F., 1 9 7 5 . Age relationships between greenstone belts and “granites” in t h e Rhodesian Archaean craton. Earth Planet. Sci. Lett., 2 5 : 251-262. Hepworth, J.V., 1 9 7 2 . T h e Mozambique orogenic belt and its foreland in northeast Tanzania: a photogeologically-based study. J. Geol. SOC.London, 1 2 8 : 461-500. Hickman, M.H. and Wakefield, J., 1 9 7 5 . Tectonic implications o f new geochronologic data from t h e Limpopo belt a t Pikwe, Botswana, southern Africa. Geol. SOC.Am. Bull., 86: 1468-1472.
421 Holmes, A., 1 9 5 1 . T h e sequence of Precambrian orogenic belts in south and central Africa. 1 8 t h Int. Geol. Congr., London, 1 4 : 254-269. Jackson, M.P.A., 1976. High-grade metamorphism and migmatization of t h e Namaqua Metamorphic Complex around Aus in t h e southern Namib Desert, S o u t h West Africa. Precambrian Res. Unit, Univ. Cape T o w n , Bull., 1 8 : 327 pp. Jahn, B. and Shih, C., 1 9 7 4 . O n t h e age of t h e Onverwacht Group, Swaziland Sequence, South Africa. Geochim. Cosmochim. Acta, 3 8 : 873-885. Johnson, R.L. and Vail, J.R., 1 9 6 5 . T h e junction between t h e Mocambique and Zambesi belts; northeast Southern Rhodesia. Geol. Mag., 1 0 2 : 489-495. Joubert, P., 1 9 7 1 . T h e regional tectonism of t h e gneisses of part of Namaqualand. Precambrian Res. Unit, Univ. Cape Town, Bull., 1 0 : 220 pp. Key, R.M. and Hutton, S.M., 1 9 7 6 . T h e western extremity of t h e Limpopo mobile belt. Precambrian Res., 3: 79-90. Kroner, A., 1 9 7 4 . T h e Gariep Group, Part 1: late Precambrian formations in t h e western Richtersveld, northern Cape Province. Precambrian Res. Unit, Univ. Cape T o w n , Bull., 13: 115 pp. Kroner, A , , 1 9 7 7 . T h e Precambrian geotectonic evolution of Africa: plate accretion versus plate destruction. Precambrian Res., 4 : 163-213. Martin, H. and Porada, H., 1977. T h e intracratonic branch of t h e Damara orogen in S o u t h West Africa 11. Discussion of relationships with t h e Pan-African mobile belt system. Precambrian Res., 5 : 339-358. Mason, R., 1973. T h e Limpopo mobile belt, southern Africa. Philos. Trans. R. SOC.Lond o n , Ser. A, 2 7 3 : 463-485. Matthews, P.E., 1 9 7 2 . Possible Precambrian obduction and plate tectonics in south-eastern Africa. Nature, 2 4 0 : 37-39. Nicolaysen, L.O. and Burger, A.J., 1 9 6 5 . N o t e o n a n extensive zone of 1000 million-year old metamorphic and igneous rocks in southern Africa. Sci. Terre, 1 0 : 497-516. Pretorius, D.A., 1974. T h e structural boundary between t h e Kaapvaal and Sonama crustal provinces. Econ. Geol. Res. Unit, Univ. Witwatersand, Johannesburg, Inf. Circ., 8 8 : 27 PP. Ramsay, C.R. and Ridgway, J., 1 9 7 7 . Metamorphic patterns in Zambia and their bearing o n problems of Zambian tectonic history. Precambrian Res., 4 : 321-337. Ray, G.E., 1974. T h e structural and metamorphic geology o f northern Malawi. J. Geol. SOC.London, 1 3 0 : 427-440. Reid, D.L., 1 9 7 6 . Radiometric age determinations in t h e Haib-Vioolsdrif region, lower Orange River. Precambrian Res. Unit, Univ. Cape T o w n , Annu. Rep., 13: 49-60. Schalk, K., 1973. Some late Precambrian formations in central S o u t h West Africa. Annu. Rep. Geol. Surv. S. Afr., 2 : 29-40. Shackleton, R.M., 1973. Correlation of structures across Precambrian orogenic belts in Africa. In: D.H. Tarling and S.K. Runcorn (Editors) Implications of Continental Drift t o t h e Earth Sciences, 2 . Academic Press, London, pp. 1091-1095. Shackleton, R.M., 1 9 7 6 . Pan-African structures. Philos. Trans. R. SOC.London, Ser. A , 280 : 491-497. Talbot, C.J., 1 9 7 3 . T h e geology at t h e eastern end of t h e Zambesi Orogenic Belt, Rhodesia. Trans. Geol. SOC.S. Afr., 7 6 : 113-133. Vail, J.R. and Dodson, M.J., 1 9 6 9 . Geochronology o f Rhodesia. Trans. Geol. SOC.S. Afr., 7 2 : 79-113. Vajner, V., 1 9 7 4 . T h e tectonic development of t h e Namaqua mobile belt and its foreland in parts of t h e northern Cape. Precambrian Res. Unit, Univ. Cape Town, Bull., 1 4 : 201 pp. Watters, B.R., 1 9 7 6 . Possible late Precambrian subduction zone in S.W. Africa. Nature, 259: 471-472.
422 Welke, H.J., Allsopp, H.J., Kostlin, E., Corner, B. and Kroner, A., 1973. Review of age measurements from the Richtersveld Complex and its environs. Abstr. 15th Congr. Geol. SOC.S. Afr., p. 92. Wendt, I., Besang, C., Harre, W., Kreuzer, H., Lenz, H. and Muller, P., 1972. Age determinations of granitic intrusions and metamorphic events in the early Precambrian of Tanzania. 24th Int. Geol. Congr., Montreal, 1: 295-314. Wilson, J.F., Bickle, M.J., Hawkesworth, R.J., Martin, A., Nisbet, E.G. and Orpen, J.L., 1978. The granite-greenstone terrains of the Rhodesian Archaean craton. Nature, 271: 23-27. Part B Kleywegt, R.J., 1966. The Gravity Survey of South West Africa. M.Sc. Thesis, Univ. Natal, Durban, 1 0 2 pp (unpubl.). Smit, P.J., 1962. The gravity survey of the Republic of South Africa. Handb. Geol. SUIT. S. Afr., 3: 477 pp.
423
Chapter 8
THE GRANITIC+NEISS
GREENSTONE SHIELD
C.R. ANHAEUSSER and J.F. WILSON
INTRODUCTION
The Early Precambrian (Archaean) granite-eenstone terrane that forms part of the southern African shield constitutes one of the oldest known fragments on the Earth’s surface. Underlying the eastern part of Botswana, the northeastern segment of South Africa, and Zimbabwe, the shield consists of two cratons - the Kaapvaal and the Rhodesian cratons - separated by the Limpopo high-grade metamorphic belt. Ranging in age from -2700 t o 3500 Ma (or even older), the region acts as basement t o a unique series of cratonic basins which formed under epicontinental conditions as far back in time as 3000 Ma ago. A great variety of volcanic, sedimentary, and granitic elements make up the Rhodesian and Kaapvaal cratons but much of the southern African Archaean is covered by younger formations which obscure approximately 70% of the cratonic crustal area. The exposed Archaean consists predominantly of granitic rock types with the greenstone belt relics constituting a relatively low proportion of the ancient basement (Table 8.1). As a consequence of their economic importance, the greenstones have generally been studied in greater detail than the areally more extensive granitic rocks surrounding them. Despite the multiplicity of rock types and other geological phenomena associated with the Archaean granite-eenstone terrane in southern Africa it has been possible to distinguish characteristic styles of tectonism, metamorphism, volcanism, plutonism, and sedimentation which have been employed in various models, the latter having been formulated t o account for the evolution of the greenstone-granite pair that constitutes the oldest shield areas (Macgregor, 1951; Anhaeusser et al., 1968, 1969; Engel, 1968; Viljoen and Viljoen, 1970a; Stowe, 1971; Anhaeusser, 1973a; Talbot, 1973; Wilson, 1973; Hunter, 1974a, b). In this chapter details are provided of the Archaean geology of Angola, South Africa and Zimbabwe. Specific attention is devoted to aspects of vol-
424 TABLE 8.1 Areal extent and relative proportions of Archaean and younger sequences developed on the Rhodesian and Kaapvaal cratons
Area (km’) Exposed Archaean (km’) Exposed granites (km’) Exposed greenstone belts (km’) Younger cover (km’) Exposed Archaean (%) Exposed granites (%) Exposed greenstone belts (%) Younger cover (%)
Rhodesian craton
Kaapvaal craton
Combined Rhodesian and Kaapvaal cratons
312,000 195,762 162,027 33,735 116,238 62.7 51.9 10.8 37.3
585,000 80,662 73,778 6,884 504,338 13.8 12.6 1.2 86.2
897,000 27 6,4 24 235,805 40,619 620,576 30.8 26.2 4.6 69.2
~~
After Anhaeusser (1976a).
canism and sedimentation in the greenstone belts, including structure, mineralization, metamorphism and geochronology. Consideration is also given t o the various granitic events recognized on the cratons, and the nature of the pre-greenstone crust is debated. A. South Africa
*
THE GNEISS TERRANE
Granitic rock types constitute over 91% of the exposed Archaean basement on the Kaapvaal craton - the balance of approximately 9% being occupied by the larger greenstone occurrences indicated in Fig. 8.1. Although much of the granitic terrane remains t o be more fully investigated, reconnaissance studies suggest that a large proportion of the area is mainly underlain by a variety of gneisses. To the southeast of the Barberton greenstone belt, and underlying the central area of Swaziland, are a suite of gneissic rock types considered by Hunter (1970, 1973, 1974a) t o be part of what he termed the Ancient Gneiss Complex. These gneisses were divided into three main types, namely: (1) a bimodal association of interlayered leuco-tonalitic gneisses and amphibolites; (2) homogeneous, gneissic biotite-hornblende tonalite; and (3) metamorphites that include quartzites, quartzo-feldspathic gneisses, siliceous and biotite-rich garnetiferous gneisses, quartz-diopside and diopside-plagioclase granulites, quartz-magnetite-grunerite gneisses, iron-formations and
* By C.R. Anhaeusser.
425 biotite-hornblende gneisses. The metamorphites attain their maximum development (underlying 200 km’) in south-central Swaziland, where they constitute a deformed, layered sequence. Smaller outcrops of quartzofeldspathic gneiss, diopside-plagioclase and quartz-diopside granulites, and biotite-hornblende gneisses are interlayered within the tonalitic gneissamphibolite unit as isolated relics. According t o Hunter (1970) the mineral assemblages displayed in the metamorphites were formed at temperatures of at least 600°C and pressures of -3 kb placing them in the high-temperature, low-pressure cordieriteamphibolite facies of the Abukuma-type metamorphic facies series (Winkler, 1967). It has been suggested that some of the banded metamorphites composed of quartz, plagioclase, biotite, hornblende, and diopside, and which outcrop in west-central Swaziland, may represent retrograded granulite facies rocks (Hunter, 1974a). The characteristic granulitic texture and platy nature of the quartz are preserved but no relic minerals diagnostic of granulite facies metamorphism have so far been identified. The tonalitic gneisses of the Ancient Gneiss Complex are medium-grained, banded rocks, the banding being caused by the alternation of layers of different colour varying from dark t o pale grey. The banding is commonly conformable with interlayered amphibolites, some large bodies of which measure up to a kilometre wide and up t o 8 km long (Hunter, 1970,1974a). The effect of the banding is accentuated in the more migmatitic gneisses where biotite and hornblende-bearing quartz-feldspar gneisses are interlayered with the tonalitic gneisses. The gneisses flanking the Barberton greenstone belt in the Transvaal include biotite- and hornblende-bearing tonalitic granites and gneisses, passing locally into trondhjemitic, dioritic, granodioritic and quartz-dioritic equivalents, as well as a variety of metamorphosed xenoliths, the latter unequivocally derived from the fragmentation of the main greenstone mass. Included in this category are diapiric plutons emplaced around the northwestern and southern margins of the Barberton belt (Fig. 8.3) and which are responsible for the superimposed structural deformations in these areas (Anhaeusser, 1975,1976d;Viljoen and Vil.joen, 1969d). The overall similarities (including major element geochemistry, texture and petrology) of the ancient tonalitic gneisses and their inclusions with the rocks in Swaziland classed in the Ancient Gneiss Complex by Hunter (1970) have been discussed by Anhaeusser (1973a) and Viljoen and Viljoen (1969d, e). Traceelement geochemistry (Condie and Hunter, 1976: Glikson, 1976; Hunter et al., 1978) permits recognition of five distinct types of siliceous gneiss within the bimodal association, and comparison with the diapiric plutons. The essential features are summarized in Table 8.11. The amphibolites interlayered with the siliceous gneisses are olivine-nonnative, have low K/Rb ratios, and flat REE patterns showing slight or no depletion in La and Ce. Siliceous gneisses similar t o Type 1 (Table 8.11) having low K/Rb ratios are reported from the Johannesburg dome (Anhaeusser, 1973b), but it is
-
426 TABLE 8.11 Summary of the geochemical characteristics of siliceous gneiss in the Ancient Gneiss Complex, the tonalitic diapiric plutons and the biotite-hornblende tonalite gneiss Major elements
Trace elements
low-K20; S i 0 2 < 73%; A l z 0 3 > 14%.
K/Rb ratio low 1 5 0 ; light REE’s enriched; heavy REE’s depleted; little or n o Eu anomaly K/Rb ratio 200; enrichment of all REE’s except for pronounced negative anomalies of Eu K/Rb ratios very low t o 1 7 5 ; relatively low abundances of all REE’s; mild depletion of heavy REE’s
low-K2O; SiOz 76-77%; A1203 1 1 ~ 1 5 . 5 % . low-KzO; S i 0 2 57-70%; A l z 0 3 17--15.5%. Dioritic t o leucotonalitic gneisses K 2 0 1.9%;Si02 < 73%; A1203 > 14%. Migmatitic gneisses quartz monzonite gneiss K 2 0 4.9%; Si02 74%; A1203 1 4 % low-K2O; SiOz < 73%; A1203 > 14%
-
Diapiric plutons
Biotite-h orn blende tonalite gneiss
low-K20; S i 0 2 < 69%; A1203 > 1 5 %
-
-
-
K/Rb ratios vary from 300 t o 1037; very steep REE patterns with small t o moderate positive anomalies of Eu 180; REE pattern similar t o K/Rb Type 1 but more depleted in heavy REE’s K/Rb 280; REE patterns similar t o 0.1, lower than Type Type 1;Rb/Sr l(-0.2) 227; REE pattern similar t o K/Rb Type 1 except for gentler slopes of heavy REE’s
-
-
-
-
not known whether they have similar R,EE patterns. Whether hornblende was involved in the generation of the siliceous gneisses with unusually low K/Rb ratios or whether this reflects migration of K and/or R,b during metamorphism is not yet established. The biotite-hornblende tonalite gneisses occupy a discrete geographical area in southwestern Swaziland. They lack pronounced banding and are coarser grained than rocks of the bimodal suite. Their K/Rb ratios are consistent and have a linear disposition about the mean value of 227 (Hunter et al., 1978). These gneisses are possibly intrusive into the bimodal suite and have yielded an Rb/Sr age of 3395 Ma and an initial 87Sr/86Srratio of 0.7006 (Davies, 1971). Even if this is regarded as the age of emplacement implying a greater age for the bimodal suite, the age relationship of the Ancient Gneiss Complex and the Barberton greenstones is not resolved. Quartzofeldspathic gneisses and amphibolites from the metamorphite suite in Swaziland have geochemical characteristics that distinguish them from the bimodal association. The siliceous gneisses, with one exception, have Rb/Sr and K/Fa ratios >1,and their R,EE patterns show enrichment in light REE’s, pronounced negative Eu anomalies, and slight depletion or flat
427 slopes of heavy REE’s (Hunter et al., 1978). One siliceous gneiss from the metamorphic suite, having similar geochemical characteristics t o the biotitehornblende tonalite gneiss, may represent an intrusive into the metamorphites, its metamorphism and deformation being contemporaneous with that of the metamorphite suite. The amphibolites in the metamorphite suite differ from those of the bimodal association in their quartz-normative tholeiitic character, low Rb contents (<4 p.p.m.), high K/Rb ratios (>lo00 p.p.m.) and REE patterns which lie between the fields for low-K tholeiite and continental tholeiite and are anomalous (Hunter et al., 1978). These patterns contrast with those reported from mafic and ultramafic rocks of the Swaziland Supergroup (Herrman et al., 1976). The relationships of this group of metamorphic rocks to the bimodal suite and the Swaziland Supergroup are uncertain. Preliminary structural study suggests the earliest recognizable phase of deformation in the bimodal suite did not affect the metamorphites. As will be discussed later the basal members of the Swaziland Supergroup (Tjakstad Sub-Group) accommodate all the lithological varieties that constitute the metamorphites of the Ancient Gneiss Complex. In the opinion of Anhaeusser (1973a) and Viljoen and Viljoen (1969e) rocks from the lower units resisted, in places, the intrusion, assimilation, and granitization accompanying the destruction of the lateral (and basal) equivalents of the Barberton greenstone belt stratigraphy in Swaziland and elsewhere on the Kaapvaal craton. However, differences in traceelement chemistry place constraints on a direct correlation of the metamorphites with the Swaziland Supergroup. In the central part of the Kaapvaal craton studies of the Johannesburg granite dome have revealed tonalitic and trondhjemitic gneisses identical to those in the eastern Transvaal and Swaziland (Anhaeusser, 1973b). Support for this contention is based on geochemical and petrological similarities that exist between the gneisses of these regions as well as the preliminary U-Pb isotopic age measurements which range between 3400 and 3200 Ma (A.J. Burger, pers. commun., 1976). Mafic and ultramafic xenoliths also occur in the gneisses providing a bimodal association similar t o that described earlier. Migmatites and banded contorted gneisses are likewise commonly encountered, particularly in areas flanking xenolithic greenstone remnants. South of Johannesburg studies carried out on the Vredefort structure have revealed further gneisses, the latter forming part of a domical, near-circular mass of basement granite. Strata which were lying on this granite have been uplifted sharply, so that they form successive rings around the central mass. Vredefort has attracted wide attention for the so-called “shock metamorphism” in the collar strata and theories of origin linking the development of the structure t o meteorite impact are popular but are not generally supported in South Africa. Geochemical data (W.F. Slawson, pers. commun., 1971) led t o the view that a major portion of the updomed basement crust
428
is “standing on edge”. Multidisciplinary studies are currently being undertaken t o investigate the claim that a radial traverse across the Vredefort basement, from perimeter t o centre of the “ring”, permits the examination of a near-vertical profile through about 1 5 km of the Earth’s granitic crust. The discovery of granulite facies rocks on the dome (D. Stepto, pers. commun., 1975) as well as progressively older age measurements from the perimeter t o centre (ranging from 3000 Ma to approximately 3500 Ma, H. Welke, pers. commun., 1975) appears to add support t o the “crustal section” hypothesis. Exposures of the gneiss terrane on the western half of the Kaapvaal craton are generally poor and as a consequence little is known about them. In the northern and northeastern Transvaal exposure varies from poor in the area north and northwest of the Pietersburg belt t o very good in the escarpment regions west of the Sutherland belt and extending from Soekmekaar, Mooketsi, Duiwelskloof and beyond towards Tzaneen (Fig. 8.6). Most of the well-exposed terrane in these areas, as well as in the regions flanking the Murchison greenstone belt, consists of tonalitic or trondhjemitic gneisses and banded migmatites, together with countless mafic and ultramafic xenolithic remnants. North and west of the Sutherland greenstone belt the gneisses and their inclusions show signs of higher-grade metamorphism with garnets being particularly prevalent in all rock types. The higher metamorphic grade in this region is due to the influence of the thermal events associated with the Limpopo mobile belt t o the north. Many of the gneissic bodies in the regions flanking the greenstone belts in the northern and northeastern Transvaal appear t o be diapiric in character having an appearance, composition, and mode of occurrence very like the tonalitic gneisses described from the Barberton-Swaziland areas and on the Johannesburg dome. It is evident from the foregoing descriptions and discussions on t h e gneissic terrane of the Kaapvaal craton that unanimity has not been achieved as t o what the ancient gneisses represent with respect t o the origin and nature of the primitive crust. The starting premise, for many models of crustal evolution, has been that an assumed sialic crust presumably accreted upon the cooling of the outer shell of the Earth and was probably in existence by about 3500 Ma ago. The main support for the prior existence of such a sialic crust on t h e Kaapvaal craton has come from Hunter (1970, 1973, 1974a) who maintains that evidence derived from structural analysis, metamorphic style, geochemistry and geophysics leads to the conclusion that the Ancient Gneiss Complex predated the development of the Swaziland Supergroup (the latter stratigraphy being regarded as common t o all the Archaean greenstone occurrences in South Africa). The opponents of the view that the Ancient Gneiss Complex represents a pre-greenstone crust postulate that no sialic basement existed beneath the lower ultramafic units of the various greenstone belts, not only on the Kaap-
429 vaal craton but also on the Rhodesian craton (Engel, 1968; Viljoen and Viljoen, 1969d, e; Anhaeusser, 1973a, 1978a). Instead, these units, which are common t o the lowermost stratigraphic divisions of all the greenstone belts, as well as the innumerable xenoliths in the granitic terranes, are regarded as remnants of a once extensive ensimatic crust which predated the nucleation of granites. Apart from the uniquely distinctive geochemical characteristics of the high-magnesian basaltic and peridotitic volcanic rocks (komatiites) and the layered ultramafic complexes (primitive ophiolite complexes?) no sial-derived detritus is known t o be associated with the lower greenstones. All indications suggest that the greenstone volcanic sequences have characteristics that are in most respects comparable with modern abyssal oceanic and island-arc situations. The origin of the ancient gneisses remains enigmatic. Recent geochemical and isotopic data (Barker et al., 1976; Hunter et al., 1978) have led to the formulation of provisional models. That preferred for the generation of the Type 1 gneisses (Table 8.11) involves partial melting of wet amphibolite under hydrous conditions, according t o the scheme outlined by Barker and Arth (1976). Constraints on the conditions under which Type 2 gneisses could be generated are too imperfectly known t o permit modelling but the absence of andesitic rocks appears to rule out hypotheses involving fractionation of basaltic liquid. A possible mechanism could involve a small degree (-10%) of partial melting of a basaltic parent of granulite-facies mineralogy but remains to be tested. The quartz monzonite gneiss (Type 5) is regarded as a partial melt of hornblende- and/or garnet-bearing gneisses of the bimodal suite, melting having been initiated where dewatering occurred upon transition from upper amphibolite facies t o granulite facies. The quartz monzonitic liquids coalesced t o form the lenticular masses now found within the bimodal suite. An isotopic age of 3203 Ma and an initial s7Sr/86Srratio of 0.7047 have been reported for these gneisses (Davies, 1971). It is perhaps significant that this age is within the range of dates given for the emplacement of the tonalite diapiric plutons which intrude the Swaziland Supergroup. The quartzofeldspathic gneisses from the metamorphite suite plot on the Qz-Ab-Or diagram close t o the low-pressure, water-saturated, minimum-melt compositions. This and their other geochemical characteristics suggest that they could have been derived from partial melting of the low-A1203siliceous gneisses and amphibolites of the bimodal suite, metamorphosed t o granulite facies and composed largely of feldspar, pyroxene and quartz (Hunter et d., 1978). The amphibolites of the bimodal suite have REE patterns similar t o abyssal tholeiites but differ in respect t o other trace-element concentrations. The geochemical data from the amphibolites in the metamorphite suite are anomalous and are not amenable t o modelling. An origin for the post-Swaziland, tonalitic diapirs by partial melting of
430 basic rocks is supported by geochemical data (Condie and Hunter, 1976; Glikson, 1976). The best match exists for all elements in models involving 10--20% melting of eclogite of average Archaean tholeiite composition (Condie and Hunter, 1976). Deformation of the greenstone belt must have placed much mafic rock in regions of higher temperature inducing melting, the liquids so produced rising due to gravitational imbalance through the volcanic pile. No models can be proposed for the remaining varieties of gneiss due t o lack of adequate data. However, the controversy as t o the nature of the early crust may now have been partly resolved as it appears that tonalitic and/or trondhjemitic liquids were derived most likely from mafic rocks which inverted t o amphibolites and/or eclogites. Having established that a basic parent was the most probable starting material for the development of sialic crust the problem now revolves around the timing of such an event. Hunter (1974a) considers some parts of the Ancient Gneiss Complex of Swaziland t o be representative of a cycle of bimodal volcanism and metamorphism predating the Barberton greenstone pile. Opposing views have been expressed because, it is claimed, no justifiable criteria are available t o distinguish the gneisses in Swaziland from those elsewhere on the Kaapvaal craton. Clearly the only satisfactory hope of settling the problem must now be sought with the assistance of the geochronologist. GREENSTONE BELTS
Unlike the Rhodesian craton, which has approximately 63%of its surface area consisting of exposed Archaean rocks, the Kaapvaal craton is largely covered by younger platform deposits composed principally of Proterozoic sediments and associated volcanic sequences. In addition, the various mafic and felsic components of the Bushveld Complex contribute to the obscuring of all but approximately 14% of the Archaean basement, the greatest development of which is exposed in the northern, northeastern and eastern Transvaal, and parts of Swaziland (Fig. 8.1). In areas west of Johannesburg further domical “windows” of ancient granitic basement appear through the cover sequences but with the exception of the Johannesburg granite dome these occurrences are generally poorly exposed and little is known about them. In South Africa there are six principal areas of greenstone development. From north t o south these include the Sutherland, Pietersburg, Murchison, Barberton, Muldersdrif and Amalia occurrences (Fig. 8.1). Of these the Barberton greenstone belt is by far the largest, best exposed, least altered, and sequentially ordered of all, and, by virtue of its high degree of preservation, has provided an unrivalled opportunity t o study practically every aspect of Archaean granite-greenstone geology. In southern Africa in particular, and, indeed, in most Archaean terranes the evolutionary models devised for the Barberton region have proved of assistance in establishing
431
LEGEND
GREENSTONE BELTS
Younger cover Maflc phase phase ~ u s h v e l dIgneous Complex Felsic Archean granites gneisses mlgmatites GREENSTONE BELT LITHOLOGICAL UNITS SOUTH AFRICA SWAZILAND Sedimentary unit Mafic-to-felsic unit Ultrarnafic-inafic unit
Figtree Moodies
1
SUTHERLAND
2
PIETERSBURG
3
MURCHISON
4 5 6
BARBERTON MULDERSDRIF AMALIA
Onverwacht
Fig. 8.1. Regional geological map o f the Kaapvaal craton in South Africa showing the distribution of the exposed Archaean granite-greenstone terrane.
stratigraphic and tectonic order. In the sections which follow accounts will be provided of the geology of the greenstone regions mentioned above. Emphasis will be given to some of the more important aspects that have emerged from the Barberton studies, and, where appropriate, the similarities and differences that exist between this region and the other greenstone occurrences on the Kaapvaal craton will be outlined.
432 Barberton Mountain Land St ra t igraphy As its name suggests the Barberton Mountain Land comprises a rugged tract of country exposed in the Lowveld of the eastern Transvaal and Swaziland. The greenstone belt consists of a wide variety of volcanic, igneous and sedimentary rock types, collectively labelled the Swaziland Sequence or Supergroup, and which, in turn, is composed of three major subdivisions that are, in ascending order, the Onverwacht, Fig Tree, and Moodies Groups. The Onverwacht Group, at the base of the sequence, attains a thickness of approximately 1 5 km, and has been subdivided by Viljoen and Viljoen (1971) into six formations (Fig. 8.2). The lower three formations constitute the Tjakastad Sub-Group, comprising an approximately 7 km-thick succession characterized by abundant pillowed and massive flows and sills of peridotite and basalt (komatiites, high-Mg basalts, tholeiites) together with very subordinate quantities of interlayered, often aluminous, felsic pyroclasts (tuffs and agglomerates). Discontinuous bands and lenses of carbonaceous shaly chert, banded chert, banded iron-formation (magnetitequartzamphibole rocks), and calc-silicate rocks (quartz-diopside-plagioclase-garnet assemblages) have been recorded from various parts of the succession, which is also referred t o as the lower ultramafic-mafic unit. In addition, small intrusive bodies of Na-rich quartz-feldspar porphyry occur in the Komati Formation of the sub-group (Anhaeusser and Viljoen, 1965; Viljoen and Viljoen, 1969a). A persistent sedimentary horizon, termed the Middle Marker (Viljoen and Viljoen, 1969b), occurs at the top of the Tjakastad Sub-Group, and heralds an abrupt change in the nature of the volcanicity occurring in the upper three formations (Geluk Sub-Group) of the Onverwacht stratigraphy. Also collectively known as the mafic-felsic unit, the upper formations are represented by a typically calc-alkaline igneous series of lavas consisting of cyclically alternating mafic and intermediate-to-acid volcanics, together with a wide variety of pyroclastic rocks. Individual volcanic cycles usually commence with wide zones of tholeiitic basalt, which are generally overlain by thinner zones of dacitic-to-rhyodacitic lavas. These felsic lavas are commonly capped by pyroclastic units and banded cherts, some attaining thicknesses of up t o 180 m. Interlayered with the massive and pillowed lavas, pillow breccias, agglomerates, and tuffs, are subordinate ultramafic bands and lenses, as well as chemical sedimentary units some of which contain evidence of primitive life forms (banded carbonaceous shales and cherts and sideritic carbonate beds), aquagene tuffs, and pelagic carbonate-chlorite sediments. Associated with some of the felsic volcanics are consanguineous intrusive bodies of quartz-feldspar porphyry (granodiorite-quartz diorite-quartz monzonite). Although most of the ultramafic and mafic rocks in the Onverwacht
Fig. 8.2. Stratigraphic column of the Swaziland Supergroup as it is developed in the Barberton Mountain Land.
434
Fig. 8.3. Simplified geological map of the Barberton greenstone belt and surrounding intrusive diapiric gneiss plutons.
succession are sub-aqueous lava flows, a number of plutonic, sill-like, differentiated bodies are also developed at various stratigraphic levels throughout the column. The majority of these are layered ultramafic-mafic complexes, and are most common in the lower ultramafic unit (Fig. 8.3), where they are considered t o represent igneous masses developed penecontemporaneously with the surrounding peridotitic and basaltic lavas (Anhaeusser, 1976c; Viljoen and Viljoen, 1970b). Overlying the predominantly volcanic successions of the Onverwacht Group and occupying the central core of the Barberton greenstone belt, is an assemblage comprised mainly of detrital sediments, with subordinate volcanic and pyroclastic members. This succession has been subdivided into an argillaceous sedimentary unit known as the Fig Tree Group, and an arenaceous sedimentary unit referred t o as the Moodies Group. These
435 stratigraphic units display disconformable, as well as unconformable relations with each other and with the Onverwacht Group. The Fig Tree assemblage consists mainly of greywackes and shales, together with siliceous chemical precipitates (banded cherts, banded ferruginous cherts) and minor trachyandesitic lavas, agglomerates, and tuffs (Visser, 1956; Condie et al., 1970). The Fig Tree Group, which attains a thickness of 2-3 km, has been subdivided into three formations (Fig. 8.2). The Moodies Group, at the top of the Swaziland succession, attains a thickness approaching 4 km (Visser et al., 1956), and has its most complete development in the Eureka Syncline. Also subdivided into three formations, the Moodies Group represents a cyclically repetitive assemblage composed predominantly of conglomerates, quartzites, sub-greywackes, sandstones, and shales, together with minor volcanic horizons, jaspilite, and banded iron-formation (Visser, 1956; Anhaeusser, 1976d). The various facies in the Moodies Group are interpreted palaeoenvironmentally in terms of interacting alluvial, estuarinedeltaic, beachoffshore and tidal flat depositional systems (Eriksson, 1977). The abundance of orthoquartzites in the Moodies Group suggests a relatively stable crust on which extensive reworking of sediments occurred. The proposed meso- and only local macro-tidal ranges imply a fairly narrow shelf width during deposition (Eriksson, 1977). The conformably underlying Fig Tree sediments are considered t o be the deep-water facies equivalent of the Moodies Group which prograded northwards over these deep-water sediments.
Structure The broad structure of the Barberton greenstone belt is that of a synform with a pronounced NNE-ENE trend. The belt furthermore maintains a directional orientation similar t o that of a number of other Archaean greenstone belts developed on the Kaapvaal craton (Fig. 8.1). Within this broad synform, there exist a number of tight synclinal folds with steeply dipping, often overturned, limbs separated by major strike faults or slides, the latter effectively dividing the Swaziland Supergroup into a number of narrow, ENE-trending segments (Anhaeusser et al., 1968). Many of these faults show evidence of having been reactivated as transcurrent faults and numerous examples of dkcollement or detachment have been identified. The intrusion of tonalitic granite diapirs of various sizes, particularly evident on the northwest flank of the Barberton greenstone belt as well as in the southwest (Fig. 8.3), has been responsible for disrupting the broad regional pattern. This has resulted in the development of numerous narrow arcuate zones of intensely deformed Swaziland succession rocks being wedged between granite domes. Flattening, stretching and attenuation of the greenstone remnants is common, and migmatites are frequently developed in areas adjoining xenoliths rafted off the parent greenstone body.
436
Me ta morph ism The Onvenvacht volcanic assemblages, in particular, have been extensively altered, in all but a few localities, by dynamothermal as well as contact metamorphism and are represented by a wide range of metamorphic rock types. The metamorphic zoning includes an amphibolite facies adjacent t o granites and pegmatites, grading away from the contacts into various subfacies of greenschist metamorphism. The telescoped nature of the metamorphic aureoles at the contacts with the younger granites, the bimodal nature of the mineral facies and the absence of blueschist mineralogy in Archaean greenstone terranes distinguishes these environments from those of later times and suggests the emplacement of the Barberton rocks, and others like them, on a crust much thinner than the present (Engel, 1968). G eo ch ro nology Isotopic age determinations carried out in the Barberton Mountain Land and Swaziland have confirmed that the majority of rock types in these areas exceed 3000 Ma in age (see Anhaeusser, 1973a, for summary of ages). The oldest ages yet reported include an Rb-Sr isochron date of 3500 f 30 Ma for basaltic komatiite from the Tjakastad Sub-Group (Jahn and Shih, 1974). This age, together with the lead ages of between 3400 and 3800 Ma reported by Saager and Koppel (1976) exceeds those so far obtained for the Ancient Gneiss Complex (Rb-Sr isochron ages of 3341 f 50 Ma and 3395 L 96 Ma) (Davies, 1971). An Rb-Sr age of 3375 f 20 Ma for the Middle Marker (Hurley et al., 1972) and U-Pb ages ranging between 3220 and 3310 t 40 Ma for some of the diapiric plutons intruding the lower formations of the Onverwacht Group (Oosthuyzen, 1970) further confirm the great age of the primitive pile (7 km thick) of basaltic and peridotitic komatiites of the Tjakastad Sub-Group. Zircons from quartz porphyry lavas in the Hooggenoeg Formation of the overlying Geluk Sub-Group have yielded an U-Pb age of 3360 f 100 Ma (Van Niekerk and Burger, 1969) but felsic lavas from this part of the stratigraphic column have also defined a Rb-Sr linear isochron yielding an apparent age of 2620 t 20 Ma. This result, which is incompatible with previously established +3000 Ma ages for granites intruding the Swaziland Supergroup (Allsopp et al., 1968), has been interpreted as the time of isotopic homogenization of the Rb-Sr system in these rocks. It was suggested by Allsopp et al. (1973) that hydrothermal alteration of the volcanics was the probable cause of the anomalous result. Higher in the stratigraphy, Fig Tree shales, dated by Allsopp et al. (1968), gave values of 2980 f 20 Ma by the Rb-Sr isochron method. This reflects a minimum age only, as both the Fig Tree and Moodies sedimentary assemblages were deformed by tonalitic diapirs such as the Kaap Valley Granite pluton emplaced 3310 2 40 Ma ago (Oosthuyzen, 1970). As mentioned earlier, the granitic rocks of the Barberton Mountain Land
437 have a variety of ages ranging between 3390 and 3220 Ma for the early tonalitic and trondhjemitic gneisses (including those of the Ancient Gneiss Complex and the diapiric plutons), and 2600 Ma for some of the late potassic plutons. Between these age extremes are some of the migmatites and porphyritic granites and gneisses developed in the Nelspruit and Swaziland areas. These have yielded ages of approximately 3160 Ma (Allsopp et al., 1962, 1968, 1969; De Gasparis, 1967; Oosthuyzen, 1970; Davies, 1971), and are gradational into the -3000 Ma-old homogeneous potassic hood granites of the highveld plateau regions. Low initial ratios (R,= 0.7006 f 0.0012) obtained by Davies (1971) from hornblende tonalitic gneisses in Swaziland are indicative of little or no crustal contamination during the evolution of these rocks. Lastly, one of the oldest syenite plutons known in the Archaean occurs south of the Barberton greenstone belt. Dated at 3130 f 30 Ma by the U-Pb method (Oosthuyzen, 1970), the Boesmanskop syenite is intruded into trondhjemitic gneisses, and, because of its low initial ratio ( R , = 0.70103, B. Ryan, pers. commun., 1975), it is considered t o be essentially of mantle derivation.
Mineralization The most important mineral occurrences in the Barberton Mountain Land include deposits of gold, chrysotile asbestos and iron-ore. In addition, tin, magnesite, talc, and barite have been mined on a small scale. Occurrences of antimony, mercury, nickel, and corundum are also known but have so far not been exploited (Viljoen and Viljoen, 1969c; Anhaeusser, 1976a). An estimated 7 million ounces of gold have been extracted from the Barberton greenstone belt. Most of this production (-70%) has stemmed from four gold mines located o n the NW flank of the belt near the town of Barberton. The deposits are mainly of the quartz-lode type with structural controls playing an important role in the localization of the gold into exploitable concentrations. Arsenopyrite, pyrrhotite and pyrite commonly accompany the gold ore which is mainly of a refractory nature. Although most of the mineralization has been hydrothermally redistributed there are indications suggesting that some of the ore may originally have been of a stratiform syngenetic type (Anhaeusser, 1976b). Rocks of the lower ultramafic unit are host t o the occurrences of magnesite and talc, as well as the ornamental stone verdite. In addition, the dunites of the sill-like layered differentiated complexes are important contributors of chrysotile asbestos mineralization. The Barberton production, combined with that from Zimbabwean greenstone occurrences, places southern Africa third in the world, after Canada and the U.S.S.R., as a supplier of chrysotile asbestos fibre (Anhaeusser, 1 9 7 6 ~ ) . Ironare has been mined at the Ngwenya deposit in Swaziland where ferruginous cherts, shales and banded iron-formations of the Fig Tree Group
438 have been converted, by processes of supergene enrichment, t o high-grade hematitic ore (Bursill et al., 1964). Banded iron-formations and jaspilite occur elsewhere in the Barberton stratigraphic column but d o not constitute ore deposits. Murchison greenstone belt
The Murchison belt consists of an isoclinally folded, narrow, elongated, sequence of volcanic and sedimentary rocks intruded by Archaean granites and gneisses (Figs. 8.1 and 8.4). The belt has a strong ENE trend and can be traced, from beneath the Drakensberg escarpment, formed by the Transvaal Supergroup in the west, for a distance of approximately 140 km t o the east, before it disappears under Karoo cover rocks which form part of the Lebombo Range on the Moqambique border. Extensive invasion of the belt by granites has caused widespread fragmentation in the southwest between Ofcolaco, Gravelotte, and Mica, as well as in the east where numerous large greenstone xenolithic remnants occur north of Phalaborwa and extend into the Kruger National Park (Fig. 8.4). Along much of the northern flank of the Murchison belt is a layered mafic body known as the Rooiwater Igneous Complex. This is essentially a gabbroic body, with titaniferous magnetite bands developed in places. A hornblende-granite phase stratigraphically overlies the gabbros. The Complex, which has been metamorphosed and deformed together with the volcanic rocks to the south, appears t o represent an ancient intrusive body possibly broadly penecontemporaneous with the Murchison greenstone successions (Van Eeden et al., 1939; Viljoen et al., 1978). Stra tigmphy Although the degree of exposure and preservation of the Murchison belt is relatively poor it has been possible to reorganize and divide the successions in the area into a lower ultramafic-mafic unit and an upper mafic t o felsic unit using the broad greenstone belt model developed in the Barberton Mountain Land. Sediments, although present, are mainly found in the mafic t o felsic unit, there being no extensive sedimentary accumulations equivalent t o either the Fig Tree or Moodies rocks of the Barberton model. Rocks constituting the lower ultramafic-mafic unit are mainly developed along the southern flank of the Murchison belt as well as in many of the xenoliths t o the south (Fig. 8.4). The principal components of this assemblage include a wide variety of magnesian metabasalts and serpentinites, together with their metamorphosed counterparts : talc, talc-carbonate, chlorite, tremoliteactinolite, and hornblende schists. Subordinate developments of banded iron-formation and felsic tuffs, the latter mainly altered to quartzsericite (fuchsite) schists, have also been encountered (Minnitt, 1975).
-- -_- -- - ----___ -_-_-_-_-_-_-
Kilometres
Hornblende Granite
+
Phalaborwa
LEGEND Granite, gneiss, migmatite, pegmatite
x x x x z\
,KX
X
X
X
X
X
X
X
X
X
r
__--
Murulampana Granite
Quartz porpnyry and sericite schist
some intermediate to felsic iavas Chloritic schists, quartzitic and sericitic interlayers, iron -formation, carbonate-rich zones Magnesian metabasalts, talcose amphibolites, serpentinite
at (1939);Engelbrecht eta1 (1962), Minnitt (1975)and Viljoen eta1 (1978)
Fig. 8.4. Simplified geological map of the Murchison greenstone belt, northeastern Transvaal.
440 Rocks classed into the mafic t o felsic unit occupy most of the central sector of the Murchison belt. In addition, many of the xenoliths in the east consist of a variety of mafic t o felsic volcanic and volcanoclastic rock types, including metabasalts, amphibole, chlorite, chloritesericite, sericite, and quartzsericite schists, and quartz porphyries. Sedimentary interlayers associated with the volcanic rocks include a range of carbonate-rich rocks, phyllites, banded cherts and iron-formations (sulphide and oxide facies), as well as quartzites, grits and conglomerates. Cyclically repetitive assemblages, commencing with mafic lavas and amphibolitic or chloritic schists, pass upwards into sericitic and quartzitic schists and iron-formations, carbonaceous shales, and dolomitic rocks often terminate individual cycles (Minnitt, 1975; Viljoen et al., 1978). The sequences constituting the Murchison greenstone belt have been intruded mainly by leuco-tonalitic and trondhjemitic gneisses, many of which have been emplaced as diapiric plutons. These have been responsible for the widespread fragmentation of the greenstone belt and the development of xenoliths. Intruded into the gneisses on the southern side of the Murchison belt is a younger potassium-rich granite pluton, termed the Mashishimala granite (Fig. 8.4).This pluton has a prominent pegmatite-rich aureole in which are developed numerous “pegmatite-type” mineral deposits. A second body of granite southwest of Mica (Fig. 8.4), referred t o as the Murulampana pluton, consists of a two mica leuco-trondhjemitic gneiss. No ages are available for this granitic body but it is possible that it represents the central core of a large diapiric pluton, the gneissic margins of which abut against the southern flank of the Murchison greenstone belt between Ofcolaco and Mica.
Structure The most important regional strhctural feature associated with the Murchison belt is the prominent ENE trend of the linear zone. Evidence suggests that this linear belt has been intermittently active for a period of geologic time probably exceeding one billion years (Button, 1973). Not only was the zone active during the evolution of the Murchison greenstone belt at a period probably predating 3000 Ma ago but it subsequently acted as a zone of accelerated subsidence and influenced the sedimentation of the Wolkberg Group, the Black Reef Quartzite, and the Olifants River Group, which together form the lowermost members of the Transvaal Supergroup. Subsequently, in pre-Pretoria Group times, the zone was once again active, this time as a belt of tectonic disturbance which resulted in the linear, ENEtrending, Mhlapitsi fold-belt developing in the Transvaal cover sequences WSW of Ofcolaco. In the isoclinally folded Murchison synclinorium the prominent ENE trending schistosity, cleavages and fold axes are locally influenced by several cross trends, the chronology of which remains t o be established. These inter-
441 secting cross trends have played an important role in localizing mineralization along the “Antimony Line”, a linear, apparently stratigraphic, zone in the central part of the Murchison belt along which a number of important antimony deposits are situated (Van Eeden et al., 1939; Viljoen et al., 1978). As is the case in the Barberton greenstone belt the numerous diapiric granite plutons surrounding and intruding the Murchison belt have contributed t o the development of the superimposed fold deformations found in the area. Metamorphism Despite the narrowness of the Murchison belt the metamorphic grade prevalent in the region is that belonging t o the greenschist facies. Only along some of the granite-eenstone contacts, as well as in xenolithic remnants, does the metamorphic grade increase in places t o the amphibolite facies. Geochronology Few radiometric age measurements are available from the Murchison granite-greenstone terrane. Saager and Koppel (1976) tentatively concluded that the maximum age of the antimony mineralization is 3020 f 50 Ma. This age, they state, would also represent a minimum age for the Murchison greenstone belt. These authors also reported the least radiogenic lead so far recorded and which was found in the ore of the Old Star gold mine, a vein deposit located 10 km WSW of Gravelotte (Fig. 8.4). Nicolaysen (1954) reported an isolated age of 2820 Ma (Rb-Sr mineral age) on mica from a pegmatite 5 km SE of Gravelotte and Holmes and Cahen (1957) provided what appears t o be a minimum age of 2060 Ma (U-Pb) for the intrusive Palabora Igneous Complex located south of the Murchison belt near the western border of the Kruger National Park. Mineralization The Murchison Range has numerous small gold deposits some of which appear t o be epigenetic in origin and generally take the form of gold-quartz veins occupying shears, fractures and faults. A second group of gold deposits, of probable syngenetic origin, can be linked t o particular parts of the stratigraphy, usually banded iron-formations, black shales or sericitic schists, and are considered t o be mainly of volcano-sedimentary (exhalative) origin (Anhaeusser, 1976a, b; Viljoen et al., 1978). The gold deposits are found in a wide variety of host rocks but the more important occurrences are often associated with the important antimony deposits located along the so-called “Antimony Line”, an assemblage of talc schists, talcose carbonates, chloritic carbonates, siliceous green carbonates and cherty carbonate rocks that can be traced for approximately 30 km from Gravelotte in an ENE direction along the central core of the Murchison belt. The Consolidated Murchison
442
mine, located in this region, constitutes one of the most important single occurrences of antimony mineralization in the world, and is certainly the largest deposit of its kind anywhere in the Archaean. The mineralization consists mainly of the sulphides stibnite, berthierite, tetrahedrite, and chalcostibnite. Gold, pyrite, arsenopyrite and cinnabar occur locally, the mercury having been mined in the past, whereas gold is still recovered as a by-product of the antimony mining. Past descriptions of the orebodies state that the mineralization is epigenetic (hydro thermal) in origin but studies in progress suggest that at least some of the antimony and associated sulphides may be volcanogenic in origin (exhalative, fumarolic). Immediately south of the Murchison belt is a prominent pegmatite area referred t o as the Gravelotte-Mica pegmatite field (Anhaeusser, 1976a). Important deposits of mica, beryllium, lithium, tantalum and columbium, corundum, feldspar and silica occur in the area. Emeralds, together with beryl are mined from pegmatites intruded into biotite schists (Van Eeden et al., 1939; Anhaeusser, 1976a; Viljoen et al., 1978). Pie ters burg greens tone belt
The Pietersburg belt, located approximately 275 km NNE! of Johannesburg (Fig. 8.1), forms an irregular, elongate, greenstone remnant averaging 9 km in width and having a strike length of over 110 km. The isoclinally folded successions have a dominant NE trend coinciding with the similarly orientated Sutherland greenstone belt situated 50 km further eastwards (Fig. 8.6). The southwestern portion of the Pietersburg belt extends beneath the Transvaal cover sequence, whereas the northeastern extremity dissipates into a zone of intense migmatite development. Archaean granites, gneisses and migmatites surround and intrude the belt which is generally poorly exposed but for a few prominent ranges of hills, the highest of which rises approximately 440 m above the surrounding flats. Stratigraphy Few details are known of the Pietersburg greenstone belt as a whole. Only the southwestern portion of the belt, between Potgietersrus and P'etersburg, has been more fully investigated, mainly because most of the small gold mines and prospects are located in this region. Mapping by Grobler (1972) resulted in the subdivision of the stratigraphy into various formations classed under the general heading of the Pietersburg Sequence. Two major cycles of dominantly ultramafic metavolcanics and sills, together with quartz porphyries and intercalated metasediments, form the lower Landsberghoek Formation. This is conformably overlain by the Eersteling Formation which consists mainly of metatholeiitic lavas as well as subordinate ultramafic and felsic metalavas and intrusive quartz porphyries. Together these formations,
.__--.To Pietersburg
PIETERSBURG SEQUENCE Conglomerates, quartzites, shales (Uitkyk Formation)
Transvaal Cover
Banded iron-formation, banded chert Basaltic komatiites, tholeiites, various mafic schists, felsic tuffs Peridotitic komatiites serpentinites, various ultrarnafic schists
-
Fig. 8.5. Simplified geological map of the Pietersburg greenstone belt, northcentral Transvaal.
444
which constitute the Mount Robert Volcanic Group, show a gradation from dominantly ultramafic rocks at the base of each major cycle, and also in the group as a whole, to mafic and felsic type upwards. The Uitkyk Formation, consisting of a dominantly arenaceous succession of conglomerates and quartzites as well as shales, overlies the Mount Robert Group unconformably, and forms the uppermost unit of the Pietersburg Sequence. The simplified geological map of part of the Pietersburg greenstone belt (Fig. 8.5) shows the general distribution of the main rock types in the area. Grobler (1972) found that most of the ultramafic and mafic rocks are chemically identifiable with the komatiite varieties reported from the Barberton greenstone belt and he contended that the Pietersburg Sequence shows a close resemblance to the broader aspects of the Barberton model outlined earlier. Thus, from the descriptions provided by Grobler (1972), it would appear that the schistose ultramafic and mafic volcanic and plutonic rocks, together with the subordinate iron-formations, cherts, felsic schists and porphyries, could all qualify for inclusion in the lower ultramafic-mafic unit. Absent, by comparison with the Barberton model, are any mafic-tofelsic volcanic rocks as well as the “Fig Tree-type” argillaceous turbidite successions. However, the dominantly arenaceous Uitkyk Formation resembles the Moodies Group but does not contain volcanic horizons or jaspilitic iron-formations. Furthermore there is a distinct cyclicity in the sedimentation of the Moodies Group which is absent in the Uitkyk Formation. Structure The Pietersburg greenstone belt, like all the Archaean remnants located on the Kaapvaal craton, has been involved in several stages of superimposed deformation, commencing with the regional development of the main northeasterly orientated folding, faulting and schistosity. This early trend has been modified by the emplacement of granite plutons and stocks, many of which are embayed into the flanks of the greenstone belt as well as occurring as intrusive bodies within the schist belt itself (Fig. 8.5). According t o Grobler (1972) three periods of deformation were operative in the region. The first period preceded the deposition of the sediments of the Uitkyk Formation. During the second phase of tectonic disturbance both the Uitkyk and older volcanogenic formations were isoclinally overfolded t o the northwest, and at the same time most of the pronounced slaty cleavage and schistosity found in the area was developed. The third deformation manifests itself as a crenulation of existing foliation in the less competent schistose rocks. Details of the first phase of deformation, which affected not only the Pietersburg rocks but also the other NE-orientated greenstone belts on the Kaapvaal craton, remain obscure. Granite diapirism, on the other hand, appears t o be largely responsible for the development of the superimposed structural events common to the Archaean greenstone belts (Anhaeusser, 1975).
445 The Pietersburg belt has been affected by prominent north-striking normal faults, the latter also displacing rocks of both Transvaal and pre-Transvaal age. A major fault with about 8 km left-lateral strike-shift cuts obliquely across the SW portion of the belt between Potgietersrust and Pietersburg (Fig. 8.5). Known as the Ysterberg Fault this NE-orientated zone of displacement has an obscure origin. It is considered by Grobler (1972) t o represent either a wrench fault genetically associated with the greenstone belt, but with reactivation in Bushveld times (c. 2000 Ma ago), or it represents an obliqueslip fault associated with the formation of the Bushveld basin. Metamorphism The rocks of the Pietersburg greenstone belt have been affected by lowgrade regional metamorphism as well as superimposed contact metamorphism. Most of the sequences display low-grade greenschist facies metamorphism except at some granite contact zones and in xenoliths where amphibolite facies assemblages prevail. Contact metamorphism, accompanied by a range of metasomatic effects, occurs in the greenstones and granite gneisses flanking the Uitloop Granite body (Fig. 8.5). This coarse-grained, potash-rich granite has been correlated with the Bushveld granites (Van Rooyen, 1947; Grobler, 1972).
M inera1iza t ion The first reef gold in South Africa was discovered on the farm Eersteling 17KS by Messrs. Button and Pigg in 1871. Other discoveries followed in the district but the goldfield as a whole did not prove t o be very significant. Willemse (1938) recorded a gold recovery of only 1010 kg for the period 1907-1937. Since then mining activity has ceased except for sporadic, small-scale efforts. Practically the entire production has stemmed from gold reefs in the greenstones but a small production has recently emanated from the Uitkyk Formation where gold occurs together with pyrite in the polymictic conglomerates. Apart from the gold, small-scale occurrences of banded iron-formation, chrysotile asbestos, and copper-nickel are known in the region but have not been exploited. The chrysotile is associated with serpentinized ultramafic rocks, whereas the Cu-Ni occurrence appears to lie along a shear zone in the granites adjacent t o the greenstone contact. The banded iron-formations in the region are locally influenced by supergene enrichment, as near the Ysterberg fault, but insufficient tonnages are available for mining. Sutherland greenstone belt The Sutherland greenstone belt, located in the far northeastern Transvaal (Fig. 8.1) remains the least documented of the main Archaean remnants on the Kaapvaal craton. Preliminary detailed studies of parts of the region are
Fig. 8.6. General geological map of the Archaean granite-greenstone terrane in th e northern and northeastern Transvaal showing t h e numerous xenoliths and areas of intense migmatite development.
447 currently underway but no written accounts have yet been made available. The greenstone belt, as mentioned earlier, appears t o represent the far eastern extension of the Pietersburg greenstone belt and is separated from the latter by a zone of intense migmatite development which occupies the terrane north of Duiwelskloof (Fig. 8.6). The western portion of the Sutherland belt is split into two arms which partly encircle a diapiric granite pluton situated immediately SW of Giyani (Fig. 8.7), the principal town of the Gazankulu homeland. In the east the belt extends into the Kruger National Park where it is dismembered by intrusive granites in much the same way as the eastern portion of the Murchison belt (Fig. 8.4) (I.C. Schutte, pers. commun., 1976).
Stru tigrup hy The Sutherland belt is, in general, poorly exposed. Nevertheless, it is apparent from the mapping, both of the Geological Survey of South Africa (1971) and M. Prinsloo (pers. commun., 1976) as well as the writer’s own observations, that the vast majority of rocks in the area comprise various schistose and massive mafic and ultramafic rocks together with subordinate interlayers of banded iron-formation, banded chert, and quartzsericitic schists. A number of massive serpentinite bodies have also been recorded both t o the east, west and southwest of Giyani (Fig. 8.7). An isolated exposure of felsic volcanic material occurs approximately 10 km east of the Louis Moore gold mine. Geochemical studies carried out on the basic and ultrabasic rocks in the area indicate affinities with the komatiite assemblages reported from the Barberton Mountain Land (M. Prinsloo, pers. commun., 1976) and it is evident that by far the bulk of the successions developed in the Sutherland greenstone belt would qualify for inclusion within the ultramafic-mafic unit as defined for the Barberton model. Apart from chemical sediments (banded iron-formations, cherts, calcsilicates) there are no stratigraphic units comparable t o the argillaceous and arenaceous members of the Fig Tree, Moodies, and Uitkyk formations of the Barberton and Pietersburg belts. Structure The Sutherland belt appears t o be the eastern extension of the Pietersburg belt and consists of isoclinally folded successions with a dominant NE trend. This early regional trend has been modified, just as in the case of the remaining major greenstone occurrences on the Kaapvaal craton, by the emplacement of diapiric tonalitic or trondhjemitic leuco-gneisses. These plutons which vary in size and degree of exposure are responsible for what has probably best been described as “balloon tectonics” (Ramsay, 1975). The emplacement mechanism envisaged is one of viscous granitic magma inflating and stretching the ductile greenstones, causing them t o either wrap
448
0
20
10
k
I
kilometres
,
/
,
, ,
7
/
/ '
Y
Mafic schists with interlayers of banded iron-formation
banded cherl and sericitic schists.
and amphibole and talc schists
Fig. 8.7. Simplified geological m a p of t h e Sutherland greenstone belt, northeastern Transvaal. (Reproduced in terms of t h e Government Printer's copyright authority 5 7 0 3 o f 1.7.1976 in modified form from t h e official sheet 2 3 3 0 B Shangoni).
around the pluton margins or to extend as tapering wedges or schist belt protuberances into the granites flanking the greenstone belts (Figs. 8.3-5 and 8.7). A more advanced stage of diapiric granite invasion results in widespread fragmentation or dismembering of the flattened or distended greenstone
449 relics, the process culminating, in many instances, in extensive assimilation, granitization, and migmatization of remaining greenstone material.
Metamorphism The Sutherland Range and its satellite xenoliths is situated close t o the southern area of influence of the Limpopo mobile belt (Fig. 8.1) and as a consequence there is a notable difference in the average regional metamorphic grade of the formations in the area. Whereas the lower greenschist facies is dominant in greenstone belts elsewhere on the Kaapvaal craton this has increased t o middle or even upper greenschist facies in the Sutherland belt. As elsewhere, amphibolite facies assemblages occur in most xenoliths and along greenstone-granite contacts, particularly the northern contact, juxtaposed t o the Limpopo high-grade metamorphic zone. Mineralization The Sutherland belt has not been prominent from the mineralization viewpoint. A number of small gold prospects and mines occur widely scattered throughout the area, many of them being associated with banded iron-formations. By analogy with examples elsewhere in the southern African Archaean the iron-formations and their contained gold and sulphide mineralization may be of volcanogenic (fumarolic, exhalative) origin (Anhaeusser, 1976b). The massive ultramafic bodies west of Giyani provide the only other exploitable mineralization, the latter consisting of small occurrences of magnesite. Geochronology No isotopic age dating has yet been attempted in either the Pietersburg or Suth erland granit e-gre ensto ne environments . Preliminary age det erminations from the adjacent Limpopo mobile belt indicate a complex history for this area dating back beyond 3400 Ma ago (B. Ryan, 1975 and J. Barton, pers. commun., 1976). The writer interprets this region t o be the reworked equivalent of a granite-greenstone terrane shredded with xenoliths similar t o that illustrated in the Louis Trichardt-Bandelierkop area (Fig. 8.6). By analogy, the dominantly developed ultramafic-mafic units of the Pietersburg and Sutherland greenstone belts may be equal t o or even older than the ages being reported from the Limpopo area. OTHER GRANITE-GREENSTONE OCCURRENCES
On the western half of the Kaapvaal craton (Fig. 8.1) a number of isolated Archaean inliers appear as “windows” through the younger cover rocks. With few exceptions exposures in these inliers are poor and little detail is known of their geology. The Johannesburg granite dome is relatively well-exposed and studies of
450 the granitic rock types in the area show that they range from trondhjemitic and tonalitic gneisses and migmatites t o homogeneous and porphyritic granodiorites and adamellites (Anhaeusser, 1973b). Rb-Sr age measurements carried out on some of the granites yielded an age of 3200 Ma (Allsopp, 1961), but preliminary U-Pb studies now suggest there is a range of ages for the granites between approximately 2900 to 3400 Ma (A. Burger, pers. commun., 1976). The greenstone remnants in the Muldersdrif area (Fig. 8.1) as well as elsewhere on the Johannesburg dome again consist predominantly of a wide variety of mafic and ultramafic massive and schistose volcanic and plutonic rocks types (Anhaeusser, 1977, 1978b), the latter being comparable in their field relations, petrology, and geochemistry t o the lowermost units of the Barberton stratigraphic model. No units of the mafic-to-felsic stratigraphy are present. It is presumed that such rocks did exist but that these were stripped by erosion and fed into the developing Witwatersrand basin t o the south leaving only remnant keels of ultramafic and mafic rocks behind. In the vicinity of Amalia, in the far western Transvaal (Fig. 8.1), Van Eeden et al. (1963) reported that the oldest rocks consist mainly of greywackes, grits, shales, banded iron-formations, lavas and tuffs. These were correlated with the Kraaipan Formation t o the north and the Swaziland successions (mainly the Fig Tree Group) of the Barberton greenstone belt. The Amalia greenstones, which are poorly exposed, were found t o be folded, metamorphosed and mineralized during the emplacement of the Archaean granites in the area (Van Eeden et al., 1963). Geochronological studies, involving Rb-Sr whole-rock dating of granites in the Western Transvaal, suggest two periods of granite formation in this region, the first at 3200 Ma and the second 2700 Ma ago (Allsopp, 1964). The remaining Archaean exposures in South Africa include the Makoppa dome which straddles the Botswana border, 200 km NW of Johannesburg (Fig. 8.1), and the Vredefort dome 100 km south of Johannesburg. Both these areas are very poorly exposed but geochronological studies of the Vredefort granites suggests ages exceeding 3400 Ma (H. Welke, pers. commun., 1976). Small greenstone xenoliths are preserved in both areas but have not been examined in detail. Archaean granites and greenstone remnants also extend south of Barberton and Swaziland into northern Fatal and Zululand where they are best seen in the dissected country north of the Tugela river. Various metamorphosed volcanic and sedimentary rocks (chlorite-actinolite-hornblende schists, serpentinites, talc schists, banded iron-formations, cherts, phyllites, grits, and cherty quartzites) were reported by Du Toit (1931) and Matthews (1959) but poor exposure, coupled with complex tectonic disturbances along the southern boundary of the Kaapvaal craton, in the vicinity of the Mtingwe thrust zone, add t o the difficulty of resolving the geological history of the area. Most of the greenstone remnants discussed in this section appear t o con-
451 tain components belonging t o the essentially volcanic lower ultramaficmafic division of the Archaean stratigraphic model. From the available literature it appears that the remnants in the southern and western parts of the Kaapvaal craton may also have components equivalent t o the Fig Tree argillites of the Barberton area. Successions of this type appear t o be absent in the greenstone remnants in the far northern part of the Kaapvaal craton suggesting that the belts exposed in this region may have undergone a greater degree of uplift and erosion relative t o their counterparts in the south and so ut h west. Mineralization in the remnants removed from the larger greenstone belts has centred mainly around gold (Amalia greenstones, Van Eeden et al., 1963), and chrysotile asbestos (Muldersdrif layered, differentiated ultramafic complex, Anhaeusser, 1 9 7 6 ~ ) .In addition, a high-grade ( 5 0 4 0 % Cr,03),low-tonnage, unworked chromite deposit (Willemse, 1948) occurs in a xenolith west of Bandelierkop (Fig. 8.6) and some of the banded ironformations offer potential as sources of iron ore, particularly the Kraaipan magnetite-quartzites in the western Transvaal, as well as others in the northern Transvaal. In the granitic terrane, deposits of corundum, associated with pegmatites, show a clustering in the region north of the Pietersburg and Murchison greenstone belts (Anhaeusser, 1976a). These deposits, described by Hall (1920), yielded large quantities of corundum and made South Africa one of the world’s leading producers of this mineral during the first quarter of this century. Small graphite deposits also occur in this region in metamorphosed carbonaceous xenoliths in the Archaean granites and gneisses. GRANITIC EVENTS POST-DATING THE GREENSTONE BELTS
The timing of some of the granitic events on the Kaapvaal craton remains open t o various interpretations depending upon which model is preferred for the origin and evolution of the Archaean granite-reenstone basement of southern Africa. As discussed earlier one school of thought regards some of the ancient tonalitic gneisses as pre-greenstone in a>e whereas a second school would argue that all the granites and gneisses so far identified on the Kaapvaal craton are intrusive in origin and should therefore be regarded as essentially later than the greenstone belts. The tonalitic and trondhjemitic diapirs (and intrusive quartz porphyries) represent the oldest intrusive granitic rocks recognized so far on the Kaapvaal craton. Details of these ancient gneisses have already been discussed in an earlier section and the influence of their diapiric mode of emplacement on the structure of the greenstone belts into which they intrude has been emphasized. The sodic character of these early gneisses contrasts markedly with the granitic events postdating the development of the tonalitic gneisses. A granite series has been recognized which demonstrates a progressive
452 increase in the availability of potassium feldspar and accounts for the development of granodiorites, adamellites and granites. Migmutites which are also commonly encountered in the basement granite terrane appear to have diverse origins, and it can be demonstrated that they form in most cases as a result of the interaction of intrusive granites (no matter what their original composition) and pre-existing greenstones (no matter what origin these may have been accorded). The resulting migmatites may therefore be compositionally and texturally variable and studies aimed at quantifying the different characteristics, including isotopic variability, are being undertaken. Preliminary findings suggest that the earliest migmatites are associated with the tonalitic/trondhjemitic intrusive events whereas some of the potassic granites have been responsible for later migmatite development as well as the reworking of pre-existing migmatites. The granite series so far identified stems mainly from work carried out in the eastern Transvaal Lowveld and Swaziland and has been supplemented by additional observations from the Johannesburg granite dome and the granite terrane in the Murchison, Pietersburg and Sutherland greenstone areas (Visser, 1956; Viljoen and Viljoen, 1969d, e; Anhaeusser, 1973a, b; Hunter, 1973,1974a; Minnitt, 1975; L. Robb, pers. commun., 1977). In Swaziland, Hunter (1973, 1974a) recognized an assemblage of coarsegrained plutonic rocks ranging in composition from ultrabasic t o acid which he referred t o as the Granodiorite Suite. This differentiated suite, consisting of three discrete intrusives elongated parallel t o the NNE-ENE trend of the Barberton greenstone belt, has not been dated by isotopic methods. Their post-Ancient Gneiss Complex age is demonstrated by intrusive and discordant contacts with the older gneisses. The suite is nowhere in contact with the Swaziland Supergroup so that its age relationship t o the greenstones cannot be established. Geochemically the members of this suite are distinguished from the Ancient Gneiss Complex and the tonalite diapiric plutons by low Rb contents ( 1 0 - 6 1 ppm), high Sr contents (414-1219 ppm), a wide range of K/Rb ratios up t o 786, and their unusual REE patterns, which are like that shown by the gabbro-trondhjemite suite of the Kalanti area, Finland (Arth et al., 1974) except that abundances of light REE’s are several times higher in the Granodiorite Suite (Hunter et al., 1978). Presumably the Granodiorite Suite was generated, like the Finland Suite, by differentiation of relatively wet, mildly alkaline basaltic liquid in which hornblende and biotite were the major precipitating phases. Similarities in petrology and major element chemistry of many of the Granodiorite Suite rocks t o those in the Ancient Gneiss Complex led Viljoen and Viljoen (1969d) t o consider that there were grounds for including them within the general category of ancient tonalite gneisses. Along the northern flank of the Barberton belt there exists a migmatite gneiss terrane consisting of a complex assemblage of gneisses, migmatites, homogeneous and nebulitic granodiorites, adamellites and pegmatites ranging
453 in age between 3160 and 2990 Ma (De Gasparis, 1967; Allsopp et al., 1968; Oosthuyzen, 1970). Similar migmatites and gneisses occur in parts of Swaziland and in areas southwest of the Barberton greenstone belt. North of Nelspruit a large potassic granite body (the Nelspruit Porphyritic Granite) as well as a smaller granodiorite pluton (the Hebron Granodiorite) have been recognized (L. Robb, pers. commun., 1977). Preliminary age determinations indicate that these bodies are approximately 3200 Ma old and predate a further extensive area of relatively homogeneous tonalite (Cunning Moor Tonalite) developed to the northeast and which appears to be approximately 2800 Ma old. South of Barberton and forming part of the Lochiel plateau and the Transvaal and Swaziland highveld, are the “Homogeneous Hood Granites” (Viljoen and Viljoen, 1969d; Hunter, 1973). The name Lochiel Granite has since been proposed by Hunter (1974a) for these 3070 Ma-old, potassic granites (granodiorites, adamellites) that have a distinctive horizontal disposition, forming a hood or carapace over the older tonalitic gneisses. Granitic rocks compositionally and texturally similar to the potassic granodiorites and adamellites north and south of Barberton and in Swaziland also outcrop on the Johannesburg dome, around the flanks of the Witwatersrand basin in the western Transvaal, and near Marydale on the southwestern side of the Kaapvaal craton. In, addition, boreholes have intersected granites of similar age (-3200-2900 Ma) at Glen north of Bloemfontein (Hunter, 1974b). The latest granitic events that appear t o terminate the acid magmatic activity of the Archaean, resulted in the intrusion of a number of mediumt o coarse-grained, often porphyritic potassic granites and syenite plutons. Two ages of pluton emplacement have been recognized - an older variety ranging from 3100-2800 Ma and a younger variety emplaced approximately 2600 Ma ago (Allsopp et al., 1962; De Gasparis, 1967; Oosthuyzen, 1970). These plutons are distinctive by virtue of their topographic expression, texture, and mode of emplacement. They sharply transgress all earlier assemblages, including granites, and have limited effects metamorphically and structurally on formations into which they intrude. Geographically these plutons are best developed along the eastern and northern flanks of the Kaapvaal craton and may owe their origin t o thermal influences connected with reheating stages in the adjacent mobile belts. Several plutons of this nature have been recognized northwest of the Pietersburg greenstone belt where reactivation in the Limpopo belt has taken place on numerous occasions (J. Barton, pers. commun., 1976). West of the possible influences of the Moqambique mobile belt, flanking the eastern rim of the Kaapvaal craton, is the Mashishimala pluton (Fig. 8.4) and the numerous plutons described in the Barberton area and Swaziland (Viljoen and Viljoen, 1969d; Hunter, 1973). No plutons of the above type are known in the central areas of the Kaapvaal craton.
454
B. Zimbabwe
*
INTRODUCTION
The Archaean shield of Zimbabwe consists of the sedimentary and volcanic piles of the greenstone belts, the associated gneisses and granitic rocks, and certain ultramafic layered intrusions and mafic dyke swarms. This total assemblage constitutes the Basement Complex which, together with the Great “Dyke” (p. 572), forms the Rhodesian Archaean craton (Bliss and Stidolph, 1969; Wilson, 1973b, and 1979). The Great “Dyke” cuts across the Basement Complex and, with an age of approximately 2500 Ma, marks the upper limit of the Archaean. Later cover rocks ranging from Proterozoic to Recent obscure much of the northwest sector, but the craton is bounded on at least three sides by mobile belts. To the north and northwest is the Zambezi belt; to the east is the Moqambique belt; and t o the south is the Limpopo belt, which separates the Rhodesian craton from the Kaapvaal craton of South Africa. The boundary of the Limpopo belt has been arbitrarily taken as the northern limit of granulite facies metamorphism (e.g. I.D.M. Robertson, 1973a, b) but deformation of Limpopo age clearly extends northwards well into the cratonic area (Coward et al., 1976a), Macgregor (1947, 1951), on the evidence of similar lithologies and certain unconformities, divided the greenstone belts into three Systems - the Sebakwian, followed by the Bulawayan and finally the Shamvaian. The recent (Wiles, 1971) and current (Stagman, 1977, 1978) nomenclature of the Zimbabwe Geological Survey retains these names for the major subdivisions with the substitution of “Group” for “System”, as a lithostratigraphic terminology which retains the relative time connotations of the Macgregor scheme. Save for recent attempts t o define formation stratotypes in the Belingwe greenstone belt (e.g., Nisbet et al., 1977), the nomenclature is entirely informal. Broadly the Shamvaian Group is the final major sedimentary unit of certain of the larger greenstone belts, whereas the much more widespread Bulawayan Group embraces the underlying volcanic pile together with some sedimentary rocks. The Sebakwian Group covers various supracrustal and other mafic-ultramafic rocks, all at various metamorphic grades, which are considered older than the Bulawayan Group. Published opinions have differed on what should be assigned to these three subdivisions and, in particular, there are major differences between the Bulawayan and Sebakwian Groups as delineated by the Zimbabwe Geological Survey (e.g., Wiles, 1971) and by authors such as Anhaeusser (1976) and Viljoen and Viljoen (1969). These differences arose largely from the applica-
* By J.F. Wilson.
455 455 tion to to Zimbabwe, Zimbabwe, by by these these authors, authors,of of conclusions conclusionsreached reached from fromtheir theirwork work tion on the the Kaapvaal Kaapvaal craton craton (epitomized (epitomized in in the the Barberton Barberton model), model), and and from fromthe the on related controversconcerning the existence of gneissic terrane predating related controvers- concerning the existence of gneissic terrane predating the main main greenstone greenstone belts. belts. the Stowe (1971) and Wilson (1973b) (1973b)have haveargued arguedstrongly stronglyfor forsuch suchaagneissic gneissic Stowe (1971) and Wilson of earlier (“Sebakwian”) greenstone basement incorporating the remains basement incorporating the remains of earlier (“Sebakwian”) greenstone belts. Anhaeusser Anhaeusser (1973) (1973) and and Viljoen Viljoen and and Viljoen Viljoen (1969) (1969)on onthe theother otherhand, hand, belts. have argued equally strongly that the high-magnesium (“komatiitic”) mafic have argued equally strongly that the high-magnesium (“komatiitic”) mafic and ultramafic ultramafic rocks, rocks, both both of of the the greenstone greenstonebelts belts and and remnants, remnants, predate predateall all and the granitic rocks and gneisses. They assigned all these high-magnesium rocks the granitic rocks and gneisses. They assigned all these high-magnesium rocks t o the the Sebakwian Sebakwian subdivision subdivision and and their their work work carries carries the the implication implication that that to these rocks rocks and and those those of of the theLower LowerOnverwacht Onverwachtof of the theBarberton Barbertongreenstone greenstone these belt are are broadly broadly contemporaneous contemporaneous (cf. (cf.Anhaeusser, Anhaeusser, 1976). 1976). belt resolve the the problems. probiems. Field Field investigainvestigaRecent studies studies have have done done much much tto Recent o resolve tions in in the the Belingwe Belingwe greenstone greenstone belt, belt, in in the the southern southern part part of of the the craton, craton, tions have confirmed an unconformity between a greenstone belt succession, have confirmed an unconformity between a greenstone belt succession, which includes includes peridotitic peridotitic flows, flows, and and adjacent adjacent gneisses gneisses (Bickle (Bickle et etal., al., 1975). 1975). which Geochronological results, particularly Rb-Sr data (Table 8.111), have also Geochronological results, particularly Rb-Sr data (Table 8.111), have also done much to clarify the position and it is now apparent that granite-greendone much t o clarify the position and it is now apparent that granite-greenof different different ages ages exist exist in in the the Rhodesian Rhodesian craton craton (Wilson (Wilsonet etal., al., stone terranes terranes of stone 1978). 1978). The oldest oldest granite-greenstone granite-greenstone terrane terrane recognized recognized isis about about 3500 3500 Ma Ma old old The and occurs occurs in in the the southeastern southeastern part part of of the the craton. craton. Its Its greenstone greenstone belt belt remremand nants, and and presumed presumed c. c. 3500 3500 Ma Ma remnants remnants elsewhere, elsewhere, constitute constitutethe the SebakSebaknants, wian Group (Stagman, 1978). The main greenstone belts are younger and wian Group (Stagman, 1978). The main greenstone belts are younger and embrace the Bulawayan and Shamvaian Groups of Stagman. Wilson (1979 embrace the Bulawayan and Shamvaian Groups of Stagman. Wilson (1979 and Wilson Wilson et et al., al., 1978) 1978)divides divides the the Bulawayan Bulawayan Group Group into into the the widespread widespread and Upper Greenstones and the more problematical Lower Greenstones. Inview view Upper Greenstones and the more problematical Lower Greenstones. In of the present rather confused transition state of the terminology, Wilson’s of the present rather confused transition state of the terminology, Wilson’s terms are are retained retained here. here. Effectively Effectively the the Lower Lower Greenstones Greenstones constitute constitute an an terms informal Sub-Group Sub-Group of of the the Bulawayan Bulawayan Group, Group, and and may may well well come come to t o be be informal viewed as as aa separate separate Group Group within within aa Bulawayan Bulawayan Supergroup. Supergroup. In In the the west west viewed of the the craton craton the the Upper Upper Greenstones Greenstones are are overlain overlain by by the the and southwest southwest of and Shamvaian Group. Group. Shamvaian Field 8.111),indicate indicate Field relationships, relationships, coupled coupled with with the the Rb-Sr Rb-Sr results results (Table (Table 8.111), that that the the Upper Upper Greenstones Greenstones are are about about 2700 2700 Ma Ma old; old;the thelow lowinitial initial s7Sr/s6Sr 8’Sr/86Sr ratios of of around around 0.701 0.701 are are consistent consistent with with this this being being an an age age of of eruption eruption ratios Hawkesworth Hawkesworth et et al., al., 1975). 1975). This This age age isis in in agreement agreement with with aa recent recent Sm-Nd Sm-Nd result result of of 2640 2640 140 140 Ma Ma from from Upper Upper Greenstones Greenstones volcanics volcanics (Hamilton (Hamilton etet al., al., 1977). 1977). This This youngest youngest greenstone greenstone belt belt development development has has not not been been recogrecognized nized in in the the Kaapvaal Kaapvaal craton. craton. The The age age of of the the Lower Lower Greenstones Greenstones isisstill stillobscure. obscure.The Thefew fewRb-Sr Rb-Srdates datesso so far ratios far obtained obtained are are also also c.c. 2700 2700 Ma, Ma, but but initial initial 87Sr/86Sr 87Sr/86Sr ratios are are higher higher than than the the near near 0.701 0.701 of of the the Upper Upper Greenstones. Greenstones.Whether Whetherthese thesehigher higher initial initialratios ratios
**
TABLE 8.111 Rb-Sr whole-rock isochron ages and initial 87Sr/86Srratios from the Rhodesian Archaean craton. 87Rb decay constant 1.39 X lo-’’ yr-’. All errors a t 2 sigma level Rock unit
87Sr/86Srinitial
Great Dyke Late Granites4hilimanzi Suite Chilimanzi 2 * 3 Zimbabwe 2 . 3 Victoria Porphyritic 2 * 3 All three as one population Late Granites-Sesombi Suite Sesombi Somabula Upper Greenstones * Maliyami Formation, Que Que area Bulawayo greenstone belt Salisbury greenstone belt Mixed Upper * and Lower * Greenstones Belingwe greenstone belt “Lower Bulawayan”, Que Que area Lower Greenstones * Hokonui Formation, Belingwe 7,8 Gwenoro gneisses, southwest of Selukwe Gneisses, Umniati River, Rhodesdale batholith, northeast of Que Que Mashaba tonalite 7 * 8 Chingezi Gneisses, Belingwe Mont d’Or formation granite, Selukwe Mushandike granite Gneisses, Shabani area Tokwe Gneisses, Mashaba area
25145 16
0.7026 f 0.0004
2620k 2610_C 2660f 2625t
80 60 70 25
0.7035 2 0.0080 0.7044 f 0.0018 0.7025 t- 0.0030 0.7040 ? 0.0010
2690 2 140 2650t- 80
0.7008 t- 0.0008 0.7012 f 0.0004
2720 t- 140 2540 f 180 2 7 3 0 f 60
0.7010 t- 0.0004 0.7015 ? 0.0004 0.7012 5 0.0007
2 7 6 0 f 70 2530 f 280
0.7029 t- 0.0002 0.7034 t- 0.0012
2620 f 120 2 7 8 0 5 60
0.7056 2 0.0005 0.7011 f 0.0002
2 7 6 0 2 80 2970 2 160 2 8 8 4 f 92 3420 ? 120 3520 k 260 3570 f 120 3580 t- 400
0.7015 f 0.0004 0.7013 f 0.0008 0.7017 f 0.0005 0.711 4 0.002 0.7017 f 0.0030 0.7000 f 0.0010 0.701 f 0.002 ~~~
* Bulawayan Group.
’
References: Hamilton (1977); Hickman (1974b); Hickman (1976); Hawkesworth e t al. (1975); Moorbath e t al. (1977); Jahn and Condie (1976); Hawkesworth and Bickel (1976); Hawkesworth e t al. (1979); Moorbath e t al. (1976); l o Hickman (1974a).
reflect a primary feature or result from metamorphism of rocks significantly older than the Upper Greenstones is still debatable. What is clear, however, is that the Lower Greenstones were deformed and eroded prior t o the deposition of the Upper Greenstones. The Rb-Sr data (Table 8.111) strongly support the earlier findings of D.K. Robertson (1973) who, on the basis of model lead ages on galenas from various mines, postulated two major events within the craton, at 3500 Ma and 2700 Ma, respectively. The earlier event he associated with the formation of granitic crustal segments and “pre-Bulawayan” lams; the later event he related t o the formation of the main greenstones belts and associated
457 granite intrusions. Two of the most recent Rb-Sr results, however (the Mashaba tonalite and the Chingezi gneisses), show the presence also of c. 2900 Ma-old granitic rocks. These form part of the pre-Upper Greenstones basement and, on regional grounds (Wilson, 1979 and Wilson et al., 1978), it is possible that, together with the Lower Greenstones, they form part of a third granite-greenstone terrane within the craton, a terrane which may also include the c. 2800 Maald Gwenoro and Umniati River gneisses. Jahn and Condie (1976) have recently offered an age of 3080 f 60 Ma for the Bulawayo-Que Que greenstone belt. It is difficult t o place much geological significance on this result which is based on a somewhat arbitrary selection of data from a wide geographic spread and from a stratigraphic range now known t o cover Lower and Upper Greenstones. Although a considerable number of major element chemical analyses are available for selected greenstone belt rocks, granites and gneisses (Phaup, 1973), only recently has any detailed systematic study of major trace and rare earth elements been attempted. Preliminary results, mainly on greenstone belt volcanics, have been presented in two recent papers: Condie and Harrison (1976) and Hawkesworth and O’Nions (1977). In addition Nisbet et al. (1977) deal in detail with the geochemistry of the Reliance-Zeederbergs succession as here described. Although aspects of these papers are discussed in this present account of the Rhodesian craton, no chemical data are listed and readers are referred to these references for details. At the time of writing a universally accepted definition of the term komatiite is still wanting. In the Rhodesian context, many of the rocks of komatiitic affinity have CaO/Al,03 ratios less than 1.0. Thus where rocks of this affinity are implied in this account, they are prefixed by the general term high-magnesium; where periodotitic, this is stated if the context requires it. THE c . 3500 Ma TERRANE
The granites and gneisses
The c. 3500 Ma rocks known at present have their largest development in the southeastern part of the craton within a roughly triangular area which has Selukwe, Fort Victoria and Shabani near the approximate corners (Fig. 8.8). This crustal segment consists mainly of gneisses such as the Tokwe gneisses west of Mashaba (Fig. 8.9) and those around Shabani (Table 8.111).Most are tonalitic and highly deformed. Together with their infolded Sebakwian greenstone belt remnants they form what appears t o be the remains of an ancient northerly trending mobile belt. Between Fort Victoria and Mashaba the deformed gneisses are intruded by the largely granodioritic mass of the near 3500 M a d d Mushandike “granite” (Wilson, 1973a; Hickman, 1974a). Southeast of Shabani, the gneisses immediately adjacent t o the Lower Green-
458
30°E Later cover r o c k s Shomvoian Group Upper Greenstones Lower Greenstones
1/1
Sebakwian Group
0 ........... *
Bu lawayan Group
Granites and Gneisses vorious age
Approximate division between western and eastern successions o f Upper Greenstones Stromotolites
Fig. 8.8. Subdivisions of greenstone belts in central area of Rhodesian Archaean craton. (Simplified after Wilson, 1979).
459
Fig. 8.9. Highly flattened banded migmatitic gneisses in Tokwe River 7 km southwest of Mashaba. Samples give Rb-Sr “errorchron” of 3580 i 400 Ma (Table 8.111).
460 stones are cut by equally old, but much smaller mass of tonalite, and elsewhere by younger granitic rocks (Hawkesworth and Bickle, 1976; Wilson, 1979). All but one of the Zimbabwean 3400-3600 Ma-old granites and gneisses for which isotopic data are available have initial 87Sr/86Srratios compatible with an origin from basic lithosphere or upper mantle (Table 8.111). The one marked exception is the near 3400 Ma-old Mont d’Or formation granite which, as discussed below, intrudes the Sebakwian rocks at Selukwe. This has an initial ratio of 0.711 which is higher than any of the dated Zimbabwean Archaean granites and gneisses, irrespective of age. Moorbath et al. (1976, 1977) suggest an origin for this granite in partial melting of earlier sialic crust with high R b S r contents similar t o those of the near 3600 Ma-old Shabani gneisses, and speculate on the possibility that the Selukwe area may be underlain by similar such crust. To what extent the granite-gneiss terrane adjacent to presumed Sebakwian rocks elsewhere in the central cratonic area is also c. 3500 Ma old is not known. Lithologically these ancient granites and gneisses are indistinguishable from those which have yielded ages in the 3000-2800 Ma bracket. The Sebakwian Group
Within the triangular crustal segment the Sebakwian Group consists of the various metasedimentary and metavolcanic rocks, and related intrusions, which together constitute what remains of the c. 3500 Ma greenstone belts. The metavolcanic rocks and intrusions are ultramafic to mafic. Felsic varieties have not been recognized with certainty but would be difficult to distinguish within the gneisses. Elsewhere in the central cratonic area, lithologically similar rocks of presumed c. 3500 Ma age are also assigned t o the Sebakwian Group. Most of these have previously been variously termed “preSebakwian”, “Sebakwian” and “pre-Bulawayan” by different authors (Bliss and Stidolph, 1969; Stowe, 1971; Wilson, 197313). Of the larger developments of Sebakwian rocks, the best documented is that of the Selukwe area (Stowe, 1968a, b y1974; Cotterill, 1969,1976, and 1979). Here the greenstone belt succession is inverted and forms the lower limb of a large recumbent fold structure known as the Selukwe nappe. The igneous rocks range from peridotitic t o basaltic and are largely metavolcanic, but also include the chromitite-bearing intrusion of the Selukwe ultramafic complex. Half-way up the succession is the sedimentary Wanderer formation, the basal conglomerates of which rest with marked unconformity on the deformed and eroded ultramafic intrusion. Table 8.IV summarizes the stratigraphic succession in the Selukwe greenstone belt with the omission of the Mont d’Or formation which outcrops in a roughly oval-shaped area, some 1 2 km across, south of Selukwe. Most of this formation consists of quartzofeldspathic rocks many of which are highly
461 TABLE 8.IV Stratigraphic succession in Selukwe area (right way u p ) simplified after Cotterill ( 1 9 7 6 ) and Stowe (1974). Mont d’Or formation omitted.
Tholeiitic metabasalt flows, metadolerite sills; minor banded iron-formation Banded iron-formation and ferruginous metapelites Conglomerates, grits quartzites, metapelites; rapid facies variation
1000 m
0-600
Nomenclature Cotterill ( 1 9 7 6 )
Nomenclature Stowe (1974)
Tibilikwe formation
Upper Greenstone formation
Wanderer formation
Wanderer formation
m ’
0-500 m
UNCONFORMITY Ultramafic intrusions, some with chromitite
Magnesium-rich metabasalt flows; metadolerites. Minor interbedded metapelites and banded iron-formation
0-800
0-600
m
m
Selukwe Ultramafic formation (including Selukwe U1tramafic Complex)
Selukwe Sequence
Lower Greenstone formation (with serpentinites)
Selukwe Greenstone formation
sericitized and this has led t o dispute as t o their original nature. Stowe, while describing the presence of irregular intrusive bodies of granitic rock, favours a largely sedimentary origin from a granitic provenance and places the formation as the oldest stratigraphic unit in the nappe structure. Cotterill (1976, and 1979) on the other hand, regards the greater part of the formation as a tonalite intrusive into the nappe structure. He recognizes meta-sedimentary enclaves, but interprets them as equivalents of the Wanderer formation. The dated samples of Mont d’Or formation (Table 8.111) are from an area where both authors agree t o the presence of later granitic rocks, and thus, on either interpretation, the nappe structure can be assigned t o the c. 3500 Ma events (Moorbath et al., 1976). Stowe (1974) described the nappe as having slid northwards from a zone of origin in gneissic terrane parallel to, and about
462 50 km south of, the northwest-trending Gwelo greenstone belt. A source from the east, however, in the northerly trending ancient mobile-belt-like terrane seems a possible alternative. The time-span represented by the unconformity at the base of the Wanderer formation is not known, but complex tectonic events predated the inversion of the Selukwe greenstone belt succession. The Selukwe ultramafic complex was deformed, with the formation of chromitite boudins, and in part metasomatised prior to the deposition of the Wanderer formation sediments (Cotterill, 1976, 1979). The phenoclasts in the Wanderer formation conglomerate include chromitite and silicified talccarbonate rock from the stratigraphically underlying Selukwe ultramafic complex More significantly they also include gneissic tonalite and granodiorite, massive adamellite and pegmatite, which indicate complex granitic events prior t o this unconformity. Whether this early sialic crust predates the whole of the Selukwe greenstone belt succession is not clear. Stowe’s interpretation of the origin of the Mont d’Or formation implies such an early granitic basement. Cotterill’s interpretation leaves the matter not proven, but the meta-pelites of his Selukwe greenstone formation (Table 8.IV), while largely derived from magnesium-rich metabasalts, do contain some quartz grains. To the east in the Mashaba area (Fig. 8.8) is a further major development of the Sebakwian Group, and a re-appraisal of the work of Wilson (1968a, 1973a) in the light of the findings of Cotterill (1976) and Hickman (1974a) indicates similarities t o the Selukwe succession. Ultramafites (some chromitite-bearing), quartz-magnetite ironstones, amphibolites and lesser amounts of diopside quartzites occur as a vast number of remnants infolded with the Tokwe and Tokwe-type gneisses. The widest single development of ultramafic rocks and ironstones, which are also at a lower metamorphic grade than the smaller remnants, is close t o Mashaba village and here the geology is further complicated by the later intrusion of the Mashaba Igneous Complex (p. 570). Southeast of Mashaba is a dominantly clastic sedimentary unit with interbedded banded iron-formation. Meta-arkoses within the unit contain small clastic grains of sodic and potash feldspars, whereas intraformational conglomerates contain clasts of microcline-bearing granite. Wilson, in describing this sedimentary unit, regarded it as a formation representing the lowest part of the Bulawayan Group of the Fort Victoria greenstone belt. The Mushandike granite however, dated by Hickman at c. 3500 Ma (Table 8.111), intrudes this formation; and Cotterill (1976) has recently provided evidence for an unconformity between part of this formation and the adjacent chromitite-bearing serpentinite which Wilson (1968a, b) described as unit 1 of the Mashaba Igneous Complex. The balance of the evidence now favours Cotterill’s suggestion that some, if not all, of this layered chromitite-bearing serpentinite, which is confined t o one relatively small area, is not part of the Mashaba Igneous Complex proper, but is part of
463 the Sebakwian Group and comparable with the Selukwe Ultramafic Complex. The sedimentary formation in this area thus becomes comparable with the Wanderer formation at Selukwe. The succession at Mashaba however is not inverted. THE c. 2900 Ma TERRANE
Conventional K-Ar ages on biotites and some hornblendes from various granites and gneisses in the central cratonic area led Wilson and Harrison (1973) to postulate two widespread pre-Great “Dyke” granitic events, at c. 2900 Ma and 2600 Ma, respectively. Wilson (197313) extended this and concluded that the 2900 Ma event marked a minimum age for the main greenstone belts. Subsequent geochronological work disproved this conclusion, and also showed the K-Ar data t o be quite unreliable as indicators of age of intrusion. No consistency of age was found among granites originally assigned t o this 2900 Ma event on the basis of the K-Ar results and later dated by Rb-Sr. Two of the most recent Rb-Sr results, however, do confirm the presence of granitic rocks of this age within the craton. In both cases the low initial 87Sr/86Srratios indicate no long previous crustal history and that the rocks are not the result of reworking of known c. 3500 Ma granites and gneisses. The results are from the northerly trending, homogeneous, slightly foliated Mashaba tonalite, which occurs between Mashaba and Selukwe and effectively separates the c. 3500 Ma rocks of the two areas; and the highly flattened tonalitic--granodioritic Chingezi gneisses on the west side of the Belingwe greenstone belt (Table 8.111). These c. 2900 Ma rocks predate the Mashaba Ultramafic Suite and Mashaba+hibi dykes and thus, as discussed later, can be regarded as part of the pre-Upper Greenstones basement. On similar geological arguments, other intrusions west of the Belingwe belt can also be regarded as part of this basement (Orpen, 1976,1978) and, although as yet undated, may also form part of this c. 2900 Ma terrane. These intrusions cover a range of composition and include diorite, an extensive area of adamellite, and the Chingezi tonalite. The Chingezi tonalite, which is lithologically similar t o the Mashaba tonalite, intrudes the Chingezi gneisses and Hokonui formation of the western development of Lower Greenstones in the Belingwe belt. The Bulawayan Group This is the most extensive unit of the main greenstone belts and, in any evaluation of its stratigraphy, the Belingwe greenstone belt is a key area. Using as a basis the succession so far established in this belt from the published and unpublished work of M.J. Bickle, A. Martin, E.G. Nisbet and J.L. Orpen, Wilson (1979, and in Wilson et al., 1978) has attempted to corre-
464
late the major stratigraphic units found at Belingwe across the central cratonic area.
Belingwe greenstone belt Here two sequences can be recognized, the Lower and Upper Greenstones, separated by an unconformity. The Upper Greenstones have been folded into a major northerly trending syncline; the lower Greenstones flank this syncline on the western and southeastern sides (Fig. 8.8). In the southeastern development of Lower Greenstones (the Brooklands formation), the succession is one of quartzites, chloritic grits and conglomerates, high-magnesium lavas interbedded with phyllites and banded ironformation. The lavas are locally pillowed and the conglomerates contain tonalite clasts. The base of the Brooklands formation is not exposed, but its relatively undeformed greenschist facies rocks abut on t o complexly deformed gneisses intruded by a small tonalite stock dated at c. 3500 Ma (M.J. Bickle, pers. commun., 1978; Wilson et al., 1978). The western development of Lower greenstones is more extensive. Orpen (1978) recognizes three formations. The lowest Hokonui formation consists predominantly of various dacitic pyroclasts and some flows, minor mafic sills and a few banded iron-formation horizons; some andesitic flows and pyroclasts may also be present. The apparently unconformably overlying Bend formation consists of alternating pillowed mafic lavas, spinifex-textured peridotites and intercalated banded iron-formation. These rocks are capped by the Koodoovale formation which is dominantly conglomeratic, but also contains a felsic agglomerate member. Boulders in the conglomerates include granitic rocks. The grade of metamorphism in all three formations is low, and in places primary textures and relict primary mineralogy are easily recognisable in the igneous rocks. Stratigraphic thicknesses are of the order of 1-2 km for the Hokonui, and 2 km and 1km, respectively, for the Bend and Koodoovale formations. Beneath the Hokonui formation is a furthern unit, the poorly exposed Bvute formation, consisting of various amphibolites, tremolite--chlorite schists and some banded calc-silicate rocks. The status of this formation is not clear. It is infolded with the c. 2900 Ma Chingezi gneisses and is more highly deformed than the Bend and Hokonui formations. Following Wilson et al. (1978) it is included, for convenience, as part of the Lower Greenstones on Fig. 2.8 but the more ‘recent work of Orpen (1978) indicates that it may well be part of the Sebakwian Group. The base of the unconformably overlying Belingwe Upper Greenstones is marked by the Manjeri formation which at its maximum development is only 250 m thick. It consists of local occurrences of clastic shallow-water sediments and some limestones, capped by a persistent horizon of predominantly sulphide-facies banded iron-formation. On the southeast side of the belt the Manjeri formation transgresses the Brooklands formation t o rest
465
Fig. 8.10. Outcrop of unconformity (arrow) between c. 3500 Ma tonalite gneiss (left) and basal beds of the Manjeri formation (Upper Greenstones).
with wellexposed unconformity (Bickle et al., 1975; Martin, 1978; Moorbath et al., 1977) on the near 3600 Ma-old Shabani gneisses (Fig. 8.10). On the western side it rests unconformably on the Hokonui (Martin, 1978), and Bend and Koodoovale formations (Orpen, 1978). Succeeding the Manjeri formation is a near 6-7 km-thick volcanic pile of massive and pillowed flows with some subsidiary sills, the whole largely devoid of sediments. The lowermost kilometre of this pile, the Reliance formation, is dominated by highmagnesium mafic and ultramafic rocks including spinifex-textured and pillowed peridotite flows; the remaining thicker part, the Zeederbergs formation, consists of tholeiitic basalts and metabasalts with minor high-magnesium varieties. (Bickle et al., 1975; Nisbet et al., 1977; Orpen, 1978). Capping the Zeederbergs formation, and forming the topmost unit in the Belingwe Upper Greenstones succession, is the Cheshire formation. This is a near 2 km thickness of dominantly pelitic shallow-water sediments with locally developed banded iron-formation and limestones.
Lower Greenstones - regional considerations Elements of the Belingwe Lower Greenstones succession can be recognized in the central cratonic area (Fig. 8.8) but detailed comparisons with Belingwe are difficult (Wilson, 1979). A thick unit of felsic volcanic rocks
466 and derived schists, lithologically similar t o the Hokonui formation, is a predominant feature in all occurrences and directly underlies the Upper Greenstones in much of this sector. On the other hand, variations of BendKoodoovale-Brooklands type lithologies are more characteristic of the linear (Mweza) greenstone belt south of Belingwe (Worst, 1956), and the possible Lower Greenstones at the southwestern end of the Fort Victoria belt. Northwest of Gwelo the Lower Greenstones lie with marked unconformity on presumed Sebakwian rocks (Macgregor, 1937; Wilson, 1979). Elsewhere their lowest preserved sequences are intruded by granitic rocks.
Upper Greenstones - regional considerations Wilson’s correlations for the Upper Greenstones in the central cratonic area are more detailed and more specific than those for the Lower Greenstones. Several are controversial and others remains t o be tested by further field work, but they allow the production of a simplified, perhaps oversimplified, picture of the Upper Greenstones story. In much of the southwestern part of the central cratonic area (Fig. 8.8) the base of the Upper Greenstones is indicated by a thin Manjeri-type marker horizon. In places this can be traced for many kilometres and is recognizable in different greenstone belts. Succeeding this marker horizon is a volcanic pile comparable t o the Zeederbergs formation in type and in thickness. It is a conspicuous feature in all the major greenstone belts and equates also with what Macgregor (1947, 1951) called the “basaltic greenstones” of his Bulawayan System. The basal 1-2 km of this tholeiitic pile, in several areas, is a high-magnesium Reliance-type volcanic assemblage. Thereafter, however, above this total volcanic assemblage, the central cratonic western succession differs from that of the eastern succession, and the difference does not seem to be merely a reflection of erosion level and granitic intrusion. In the larger, eastern greenstone belts, such as those of Gwelo, Fort Victoria, Shangani-Fort Rixon, the volcanic rocks are capped by Cheshire-type sediments which are overlain by, and in part interbedded with, a further development of Zeederberg-type lavas not seen at Belingwe. In the west, however, the basaltic pile passes up into repeated mafic-felsic volcanic sequences of different thicknesses. In these, tholeiitic and some highmagnesium (basaltic and peridotitic) spinifex-textured flows alternate with felsic flows and a variety of related pyroclasts. Various sediments, derived from the shallow-water reworking of these felsic rocks, some banded ironformation and a few limestones are also a feature of this bimodal volcanic succession. Where metamorphosed and deformed the felsic rocks and some of the derived sediments are sericitic schists. The eastern pelitedominant sedimentary unit and its overlying and interbedded basalts can be traced laterally into the western bimodal pile so that the two are, in effect, broad time equivalents (Wilson, 1979). Succeeding the western bimodal pile is an
467 andesitedominant, calc-alkaline volcanic assemblage, ranging from basalt t o dacite, which shows no indication of having been developed in the east. Table 8.V summarizes the eastern and western successions in the central cratonic area, and Fig. 8.8 shows their approximate areal distributions. In the Gatooma and Que Que areas the bimodal pile includes part of the felsic formation of Bliss (1970; and in Bliss and Stidolph, 1969) and Harrison (1970; and in Condie and Harrison, 1976); and the calc-alkaline assemblage includes the Umniati group and Maliyami formation (Table 8.111) respectively, of these two authors. TABLE 8.V Simplified western and eastern successions of the Upper Greenstones in the central cratonic area (After Wilson, 1979) Western succession
CALC-ALKALINE SUITE of basalt, andesite, dacite flows and pyroclasts; some mafic sills (? k 4 km thick) BIMODAL VOLCANIC SUITE tholeiitic, and some high-magnesium, basalt and metabasalt pillows and massive flows, some peridotites; alternating with dacitic flows, tuffs and agglomerates. Mafic sills. Grits and conglomerates locally derived from pyroclasts. Sericitic schists. Banded iron-formation, phyllites, local limestones in places stromatolitic. (? 4 km thick)
Eastern succession
Formation name in Belingwe belt (eastern succession)
Not developed
-
THOLEIITIC, and some highmagnesium, basalt and metabasalt pillowed and massive flows, in part interbedded with underlying phyllites ( ? 1-2 km thick) PHYLLITES, banded iron-formation, local conglomerates and grits; some limestones in places stromatolitic (up to 2 km thick)
-
Cheshire
+_
As for eastern succession
As for eastern succession, where recognized
THOLEIITIC basalt and metabasalt pillowed and massive flows. Subsidiary sills. Very minor pyroclasts. (4-6 km thick) PERIDOTITIC and HIGH-MAGNESIUM basalt pillowed and massive flows (1-2 km thick) Banded iron-formation, local conglomerates and grits, some limestones in places stromatolitic ( u p to 0.25 km thick)
Zeederbergs
Reliance
Manjeri
468
Detailed evaluation of the stratigraphy is still required for the greenstone belts to the north and south. The main features of the eastern and western successions, however, can still be recognized and indicate that the western succession is confined t o a broad zone traceable from Botswana (e.g., Key et al., 1976) north-northeastwards across the whole craton (Fig. 8.11). On Wilson’s correlations the five localities of stromatolites now known in the Zimbabwean Archaean are restricted to the Upper Greenstones, and to limestone occurrences above (three) and below (two) the thick Reliancecum-Zeederbergs-type volcanic pile (Fig. 8.8 and Table 8.V). The upper lime-
Fig. 8.11. T h e main greenstone belts (Bulawayan and Shamvaian Groups) of the Rhodesian Archaean craton with the Great “Dyke” and its associated fractures. I = Umvimeela dyke; 2 = Great “Dyke”; 3 = Main swarm; 4 = Bubi and Crystal springs swarms; 5 = East Dyke; 6 = Popoteke fault. The dotted line represents the likely craton-wide division between the western and eastern successions of the Upper Greenstones as recognized in the central cratonic area (Modified after Wilson, 1979).
469
Fig. 8.12. Strornatolites near base of t h e Cheshire formation (Upper Greenstones) of t h e Belingwe greenstone belt.
stones have so far yielded the most spectacular examples, such as those near the base of the Cheshire formation (Fig. 8.12) at Belingwe (Bickle et al., 1975) and the Huntsman limestone stromatolites about 60 km north of Bulawayo (Macgregor, 1941). As described later, the correlations also suggest that economic sulphide-
470 nickel mineralization within the Basement Complex is possibly restricted t o the Upper Greenstones. The Shamvaian Group The Shamvaian subdivision, as a System, was first established by Macgregor (1947) who correlated several unconnected sedimentary sequences occurring above his Bulawayan System in various greenstone belts. His type area was the Shamva Grits northeast of Salisbury. In other areas subsequent workers also assigned certain sedimentary sequences to this major subdivision and a tendency arose to regard all sedimentary rocks overlying a thick volcanic pile as part of the Shamvaian subdivision. Bliss and Stidolph (1969) stressed the point that some of the rocks classed as Shamvaian, even involving some originally included in this unit by Macgregor, might be integral interbedded parts of the Bulawayan succession. The recent attempt at correlations across greenstone belts (Wilson, 1979) has emphasized this point further and has shown that rocks of the Cheshire-type sedimentary unit seen at Belingwe, and parts of the reworked felsic volcanic assemblages of the western bimodal suite of the Upper Greenstones have, in some areas, been assigned to the Shamvaian subdivision by previous workers. Implicit in Macgregor’s concept of the Shamvaian System, and retained in the current Group terminology of the Zimbabwe Geological Survey (Stagman, 1978) is that it consists of rocks deposited unconformably on the Bulawayan succession, and that it is separated from this succession by major folding. On this basis the distribution of the Shamvaian Group in the central cratonic area is shown in Fig. 8.8. To the north, in the Salisbury greenstone belt and in the belts west of Salisbury, major developments of Shamvaian Group rocks also occur. No detailed sedimentological studies have been attempted on the Shamvaian Group but it seems probable that it was deposited in a number of isolated basins. The main rocks are generally poorly sorted arkoses and subgreywacke types, rapidly deposited in shallow water, and derived from a terrane that was both volcanic and granitic. Many of the conglomerates are characterized by the extreme polymict nature of their phenoclasts. In the Shamva area, north of Salisbury, granitic boulders and pebbles are conspicuous, but these are not always numerous in other areas. Harrison (1970) in the Que Que area, recognized transgressions and unconformities between the four formations which here constitute the Shamvaian Group. A basal gritconglomerate formation is overlain by a pelitic formation with ferruginous slates, which is in turn overlain by two further formations dominated by fine- t o mediumgrained sub-greywackes. In the Fort Victoria greenstone belt, limestones mark the base of Shamvaian rocks in the east and west (Wilson, 1964). Thick, mostly polymict conglomerates are, however, characteristic of much of the lower part of the
471 highly deformed succession. In the upper part of the succession there is a major development of pelitic rocks which in some areas are markedly ferruginous with a prominent development of banded iron-formation. Metamorphism of these iron-formation beds has resulted in the local production of a variety of riebeckite. Some interbedded felsic volcanic rocks are present in the Shamvaian Group in the Fort Victoria area. LATE GRANITES
On the latest geological map of Zimbabwe (Stagman, 1977) the Archaean granitic rocks are separated into the Younger Granites and the Older Gneiss Complex. The latter subdivision does not consist entirely of gneisses and it
Fig. 8.13. Distribution of the late (“Younger”) granites in relation to Zimbabwean greenstone belts. (Modified after Stagman, 1 9 7 7 ) .
472 does not necessarily demarcate the pre-Upper Greenstones basement, but the Younger Granites do represent the best delineation so far of intrusions which postdate the main greenstone belts (Fig. 8.13). In the central cratonic area two main suites, termed Sesombi and Chilimanzi, respectively (Wilson et al., 1978; Wilson, 1979), can be recognized among these Younger Granites (Fig. 8.14).
Sesombi Suite This consists of a number of tonalites and some granodiorites which occur in the west in a broad zone parallel t o the north-northeast-trending line of the western succession of the Upper Greenstones. The Sesombi tonalite, which typifies the suite, intrudes the Maliyami formation of the western calc-alkaline assemblage northwest of Que Que. It has a marked metamorphic aureole up t o hornblende-hornfels facies, which fades gradually into the low regional greenschist facies of the country rocks. Steeply dipping stretch lineations in the aureole, together with marginal planar fabric in the pluton itself, are consistent with diapiric intrusion (Robertson, 1976). In a few areas the pluton cuts across the aureole fabric and is itself massive on the contact, with some inclusions. The full extent of the Sesombi Suite is not accurately known at this stage but it seems likely that the tonalites and gneissic tonalites separating the Fort Rixon-Shangani greenstone belt from the Bulawayo belt t o the west are all part of this suite (cf. Figs. 8.8 and 8.13). Added complications arise, both inside and outside the central area, with the possible presence of older tonalites such as the c. 2900 Ma Mashaba-type. Nowhere are rocks of the Shamvaian Group in contact with Sesombi Suite proper in the central area, but small dioritic intrusions equated with the Sesombi pluton (Harrison, 1970) cut the Shamvaian Group in the Que Que area. Most authors assume a post-Shamvaian age for the suite but it seems possible that the early stages of its intrusion were occurring while Shamvaian Group sediments were being deposited on surface. The Rb-Sr ages and low initial 87Sr/86Srratios for the Sesombi and Somabula intrusions, the only members of the suite so far dated, are indistinguishable from the Maliyami formation calc-alkaline rocks (Table 8.111). Also, Hawkesworth and O’Nions (1977) have pointed out chemical similarities between the Sesombi pluton and certain of the dacitic volcanics near Bulawayo. Thus the Sesombi pluton and, by analogy, the Somabula tonalite together with the whole suite, can perhaps be regarded as late plutonic expressions of the same major magmatic episode which produced the andesites and dacites of the upper part of the western succession. The low initial ratios, moreover, argue against sources for these tonalites and volcanic rocks in remelting of earlier granite-gneiss terrane (Hawkesworth et al., 1975; Hawkesworth and O’Nions, 1977).
473
Chilimanzi Suite This later suite of intrusions constitutes the last major granitic event of the Rhodesian Basement Complex. In spite of the unreliability of the K-Ar results, the Chilimanzi Suite effectively encompasses the “2600 Ma granites” of Wilson and Harrison (1973; and in Wilson, 1973b) with some additions. Petrologically the rocks are largely adamellites with some granodiorites, and most are medium- t o coarse-grained. Some intrusions contain magacrysts of potash feldspar; and these porphyritic varieties may constitute a complete intrusion or may be only of sporadic development within a massive body. Locally a gneissic margin may be developed but, in general, the intrusions are evengrained and massive up t o their sharp contacts with the country rocks. Their final emplacement, at the levels now exposed, appears to have been passive with stoping and fragmentation of country rock to form xenoliths. In places they display a thermal aureole up to hornblende-hornfels facies. Although the suite is defined in the central area (Fig. 8.14) similar intrusions occur across the craton (Fig. 8.13). East of the Great “Dyke” they make up the great majority of the extensive “Younger Granites”, although some of the larger masses are composite intrusions. The three-dimensional shape is not known but some are clearly sheet-like in form. In the south, in particular, an east-northeasterly trend is apparent in the outcrop pattern of several of the intrusions. In the extreme east the granites are deformed by tectonism associated with the Mocambique mobile belt. The limited isotopic evidence is compatible with an origin for the Chilimanzi Suite, as opposed t o the Sesombi Suite, in remelting of earlier sialic crust. Rb-Sr dating has been done on three of the major intrusions in the southern area, the Chilimanzi, Zimbabwe and Victoria Porphyritic granites respectively (Table 8.111 and Fig. 8.13). When the whole-rock data from all three are plotted as one population, the result is a good isochron, free from any indication of geological scatter, which gives an age of 2625 t 25 Ma and an initial 87Sr/86Srratio of 0.7040 (Hickman, 1976). I.D.M. Robertson (1973a) in discussing the origin of several of the Chilimanzi Suite granites in the southeast of the craton, suggested that they were late-stage derivatives of granulite metamorphism in the Limpopo mobile belt. Hickman (1976), in effect, supports this view and has explained the somewhat higher initial ratio of these southern granites by a derivation in partial melting of certain c. 2900 Maald gneisses within the mobile belt. Exactly what parts of the pre-2600 Ma crust might have been involved under the exact cratonic area elsewhere must remain speculation at this stage (cf. also Hawkesworth et al., 1979). LAYERED INTRUSIONS AND DYKE SWARMS
Within the central cratonic area are a number of ultramafic layered intrusions, related sills and certain mafic dyke swarms (Fig. 8.14).
474
M a s h a b a - Chibi d y k e s
L a t e r cover r o c k s
Chilimanzi Suite
Sesornbi
Suite
1
I
[z1 Moshaba
Ultramafic Suite
Late GrOn'tes
Upper Greenstones ( a n d Shamvaian Group1 Lower Greenstones
Fig. 8.14. Distribution of late granites, Mashaba-Chibi dykes, and Mashaba Ultramafic Suite in the central cratonic area (Simplified after Wilson, 1979, and Stagman, 1977). 1 = Sesombi tonalite; 2 = Somabula tonalite; 3 = Chilimanzi batholith; 4 = Victoria Porphyritic granite; 5 = Zimbabwe batholith; 6 = Mashaba Igneous Complex; 7 = Shabani intrusion; 8 = Gurumbatumba-Filabusi intrusion; 9 = Shabani intrusion; 10 = Ingolubi intrusion.
475 The best documented of the layered intrusions are those of Mashaba (p. 570) and Shabani (Laubscher, 1964, 1968). These and several smaller ultramafic sills intrude the crustal segment containing the major development of c. 3500 Ma rocks. The Mashaba Igneous Complex also intrudes the c. 2900 Ma Mashaba tonalite and is itself cut by adamellites of the c. 2600 Ma Chilimanzi Suite. Also within this crustal segment is a major development of uralitized and saussuritized mafic dyke swarms. Of these some constitute modified ring and radial patterns (cf. Fig. 9.16) considered by Wilson (1968a, b) t o be late phases of the Mashaba Igneous Complex, whereas a few, near Mashaba village, predate the Complex. Other dykes, t o the northwest and southwest of Mashaba, strike east-northeast. All these dykes are conspicuously absent from the Chilimanzi Suite and in places are seen t o be cut by these late adamellites. Several other layered intrusions occur in the western part of the central cratonic area. Wilson (1979) correlates these with the Mashaba-Shabani intrusions and collectively terms them the Mashaba Ultramafic Suite; the various dykes, including some southwest of Belingwe (Orpen, 1976), he groups as the Mashaba-Chibi dykes. The dykes have not been recognized in the Upper Greenstones, and the Mashaba Ultramafic Suite is largely confined t o rocks older than the Upper Greenstones. West of the Great "Dyke" most of the occurrences of ultramafic intrusions are in Lower Greenstones. West of Belingwe, however, the Gurumbatumba-Filabusi layered intrusion, a southerly dipping sheet, cuts not only the deformed Lower Greenstones of the Belingwe and Filabusi greenstone belts, but also the various tonalitic rocks intrusive into them. At Belingwe this ultramafic intrusion and various dykes cut the Hokonui formation as well as the Chingezi tonalite and c. 2900 Ma Chingezi gneisses. To the northwest of Shangani an undated tonalite-granodiorite, Sesombi Suite-type stock truncates the Shangani ultramafic intrusion (Harrison, 1969). Wilson (1979) suggests that these various ultramafic intrusions and dyke swarms are coeval with the thick, high-magnesium t o tholeiitic, volcanic pile common to the western and eastern successions of the Upper Greenstones; and that the dykes were possibly feeders t o some of the flows. The more recent unpublished work of Orpen (1978) on the Ingolubi Complex (Fig. 8.14) at Belingwe lends support t o the proposal. A large sill of this complex cuts not only the Koodoovale formation of the Lower Greenstones, but also the Reliance formation of the Upper Greenstones. The subsequent folding of the Upper Greenstones has affected the sill. STRUCTURE
With the recognition of granite-eenstone terranes of different ages and a time-span approaching a thousand million years, the structural story of the
476 Rhodesian Archaean craton is extremely complex Little is really understood of the tectonic history of the pre-Upper Greenstones basement, but the c. 3500 Ma events, as described earlier, included the development of the major recumbent fold/nappe structure of Selukwe. A variety of granitic rocks was also involved, some postdating and some predating the nappe development. Even less is currently known of the structure of the Lower Greenstones. To a very large extent, however, the configuration of Zimbabwean greenstone belts is that of the Upper Greenstones, and it was, in effect, the main elements of this pattern that Macgregor was trying to explain in his classic review paper of 1951. In this he drew attention to the synclinal structure and arcuate margins of the greenstone belts and elaborated on the idea, first suggested by Maufe (Maufe et al., 1919) of a causal connection between the greenstone belt structures and the upward rise of the intervening, apparently ovoid, granitic masses. He termed these ovoids “gregarious batholiths”. The essential control in this concept is gravity. Macgregor envisaged the downsagging of the dominantly volcanic cover into a less dense sialic substratum. As it sank, at points where “for some reason or other it was heaviest”, the granites rose, and the overlying rocks slid outwards and towards the sinking portions. Many workers have since accepted multiple intrusion of granites as responsible for the tectonic style of Archaean cratons and for most of the thermal and dynamic metamorphism of the greenstone belts. Their derivation of the granitic magmas is dependent on the greenstone belt evolutionary model they invoke, and may involve melting of sialic basement, root-zone melting of the greenstone belts or direct derivation from the mantle (cf. Anhaeusser et al., 1969; Hunter, 1974; Anhaeusser, 1975). Macgregor’s diagram depicting his gregarious batholiths is reproduced in Fig. 8.15. That it is a very idealistic representation (which he emphasized in his original (1951) paper) is now quite clear. His ‘%batholiths”are not, as his map might suggest, homogeneous masses with gneissic margins. They contain granites and gneisses of various ages, some post-dating and some predating the Upper Greenstones. The trends shown within his ovoids are largely hypothetical and many are now known t o be incorrect; indeed in Stagman’s (1977) simple two-fold subdivision of granitic rocks, even the ovoids themselves are no longer apparent (Fig. 8.13). In effect, Macgregor’s batholiths are not really batholiths at all. Nevertheless most of his batholithic names are still retained in one form or another; some for specific large intrusions and suites, e.g., Chibi, Chilimanzi and Sesombi; others for the geographical areas covered in Macgregor’s original use of the words; e.g., Rhodesdale and Shangani. In the light of present knowledge Macgregor’s explanation for the @anite-greenstone terrain pattern is a grossly oversimplified one. Diapiric intrusion has undoubtedly played a deformational role but, on the other hand,
477
Fig. 8.1 5. A.M. Macgregor’s diagram explaining Zimbabwean Archaean structures by gregariaous batholiths. (From Macgregor (1951), reproduced with permission of the Geological Society of South Africa).
the widespread late adamellites trangress most of the structures and postdate most of the tectonism. Nor are the greenstone belts everywhere synclinal; in places they are remnants of larger structures disrupted by invading late granites. Only recently, however, has any serious attention again been given t o explaining the pattern in alternative or additional ways which do not involve vertical tectonics. Snowden and Bickle (1976) consider that a fold interference pattern, with a west-northwest trend superimposed on an earlier east-northeast trend, best explains the configuration northeast of Salisbury, i.e., in the area which includes Macgregor’s Chindamora and Madziwa batholiths. They recognize
478 these deformations in the main greenstone belt rocks and in the youngest granites (of the Chin(d)amora batholith) which intrude them. Snowden (1976) further suggests that interfering cross folds might have produced the cratonic pattern of greenstone belts. The earlier work of Stowe (1971) is also important in this respect. He recognized four phases of tectonic development within the craton. His primary phase largely corresponds t o what are here called the c. 3500 Ma events; his next three phases involved the main greenstone belts. Some of his correlations and suggestions may have been superseded by the more recent evidence, but basically, in dealing with the main greenstone belts, he emphasized the widespread, as opposed t o local, nature of the deformations. He postulated regional cross folding on northeast and northwest trends; followed by a phase of vertical tectonics effected by gneissic domes produced from the reworking of older basement; and finally a period of regional eastnortheast-trending folding and a flood of post-Shamvaian granites. The studies of Coward et al. (1976a), Coward and James (1974), Coward (1976) and Key et al. (1976) also confirm that much of the deformation within greenstone belts is regional and not simply due t o local intrusion of granites. The first group of authors go further and emphasize that tectonic events in the craton cannot be divorced from such events in adjacent mobile belts. From their studies of the southern and southwestern parts of Zimbabwe and northeast Botswana they trace tectonism of Limpopo age northwards, well into the Rhodesian craton. Coward et al. (1976a) explain deformation in the craton and adjacent mobile belt in terms of large intracratonic block movements. During the main periods of Limpopo deformation they consider that the Rhodesian cratonic block moved southwestwards relative t o the mobile belt, by possibly as much as 200 km. Within the cratonic block they envisage a scheme of “billiard ball” tectonics in which smaller blocks moved relative t o one another and produced local shear zones of intense flattening. The horizontal shortening resulting from the southwesterly movement was most intense in northeastern Botswana and southwestern Zimbabwe, with up to 50% in the greenstone belts. The amount was less in the intervening granitic areas. Some of these behaved homogeneously and show a pervasive fabric throughout; others behaved as massive competent bodies which, internally, deformed heterogeneously with the local production of mylonitized zones. In the southeast of Zimbabwe the block movement produced east-northeast-trending sinistral shear zones parallel t o the Limpopo trend. The linear east-northeast trend of the greenstone belts in the southeast part of the craton (Fig. 8.13) is thus parallel t o the major shear direction, and the near northwesterly aligned greenstone belts, which in places curve into the shear trend, are approximately perpendicular t o the shortening direction. The story, however, is even more complex since regional deformation pre-
479 dating these main periods, and affecting what are apparently Upper Greenstones, can also be recognized in the southwest of the craton. Litherland (1973) and Key et al. (1976) have speculated on the possibility of early nappe structures in northeastern Botswana with northeasterly directed movement of the nappes. An extensive area of overturned volcanic rocks is found in the greenstone belts in this area; the rocks dip southwest but young to the northeast. These early structures were intruded, and to some extent further deformed, by diapiric granites (possibly the Sesombi Suite) before the main phases of deformation. Within Zimbabwe, some 120 km south of Bulawayo, Coward et al. (1976b) suggest possible thrusting of earlier gneisses over greenstone belt rocks as part of this early deformation. It is difficult to evaluate the extent and full meaning of these early structures. Earlier block movements may well explain some, but their recognition, together with the regional pattern of subsequent deformation of the Upper Greenstones, make it clear that the configuration of the greenstone belts of the Rhodesian craton can no longer be explained merely in terms of multiple granitic intrusion. METAMORPHISM
The long time-span involved in the development of the Rhodesian craton also places constraints on the evaluation of the metamorphic history, which is probably just as complex as the structural story. The scattered remnants of the Sebakwian Group, however, show regional metamorphism up t o amphibolite facies around Mashaba (Wilson, 1968a) and north of Umvuma (Bliss, 1962) and t o upper amphibolite, and possibly even granulite facies, west of Selukwe (Stowe, 1968a, 1973). Metamorphic olivine-orthopyroxene assemblages are recorded by Stowe in the Selukwe area. It is not clear t o what extent these form part of the early regional events; some, but not all, occur near later dolerite dykes and thus may represent local contact effects. East of Mashaba the contact effects of the c. 3500 Ma Mushandike granite have produced andalusite and cordierite in the adjacent Sebakwian Group sedimentary rocks. In the main greenstone belts (ignoring the complications of the Lower and Upper Greenstone subdivisions in this context) most workers have recorded distinct and rapid increases of metamorphic grade as intrusive granite contacts are approached. Saggerson and Turner (1976) however, in the only regional metamorphic appraisal so far attempted, recognized a metamorphic zoning across the craton, which is not solely dependent on granitic intrusions as a heat source. In defining zones, Saggerson and Turner (1976) use the maximum regional metamorphic grade reported from any single greenstone belt and note that rocks of lower grade may occur. Their zonation is semiconcentric and is evident both in the Bulawayan
480
and Shamvaian Groups. The pattern is obscured by t h e later cover rocks in the west, and by the (later) Moqambique mobile belt in the east. The lowest grades are found in the Midlands (Gwelo-Que Que-Hartley) area, where the paragenesis is that of zeolite t o lower greenschist facies, with the maximum of these lowest grades found in the Maliyami formation volcanic rocks. Many of these retain their original textures despite some replacement by carbonate; prehnite and zoisite are stable, and zeolite-filled amygdales are still preserved. Outwards from the Midlands area there is a progressive increase through lower and upper greenschist facies t o amphibolite facies on the cratonic margins with the Limpopo mobile belt in the south and the Zambezi belt in the northeast. The occurrence of andalusite and cordierite-anthophyllite, and the lack of garnet in the plagioclaseamphibolite assemblages, emphasize the low-pressure facies conditions of the regional metamorphism. The authors suggest that the regional zonation is related t o thermal highs along the Limpopo and possibly the Zambezi belts. They note that the late adamellites transgress the regional zonation pattern and suggest that continued thermo-tectonic activity from the Limpopo and Zambezi zones may have produced these late cratonic intrusions. In this, their thinking parallels that of I.D.M. Robertson (1973a) and Hickman (1976) who, as mentioned earlier, have also argued that certain of the Chilimanzi Suite in the south cratonic area are related t o events in the Limpopo mobile belt. MINERALIZATION
Economic mineralization covering a wide range of minerals and metals is an important feature of the Zimbabwean Basement Complex, but here it is possible t o mention only some of the more general aspects. Gold mining was the mainstay of the country’s economy for many years and it was realized at an early stage that the gold occurred in the greenstone belts, or “gold” or “schist” belts as they were locally known, and to a lesser extent within the granites and gneisses. Phaup (1964) offers the simplest subdivision of Zimbabwean gold deposits and recognizes two main types; these are quartz veins and sulphide replacement bodies. The quartz-vein deposits are higher in grade, but gold concentrations occupy only a small portion of the vein strike and the veins are notoriously lenticular. Gold mineralization in these deposits invariably shows a structural control related t o regional and local stress fields. The replacement deposits are larger and Phaup describes the commonest type as mispickel, pyrite pyrrhotite and, in places, stibnite replacement of greenstone belt volcanic rocks and sedimentary schists. He also describes pyrrhotite-pyrite gold replacement bodies in faulted and brecciated banded iron-formation. More recent authors (e.g., Anhaeusser, 1976; Fripp, 1976) have placed a different emphasis on the association of stratiform gold-
481 sulphide mineralization in such banded iron-formation rocks and ferruginous cherts. Fripp looks t o a genetic link between this association and volcanic exhalative processes, although clearly later fracturing and folding have helped to control economic concentration. Zimbabwean Archaean gold deposits, whatever might be the age(s) of actual mineralization, are not restricted to any one age of greenstone belt rocks. Gold has been won from numerous mines in the Sebakwian Group at Selukwe (Stowe, 1968a, c) and from a lesser number in the Sebakwian rocks at Mashaba. Most deposits, however, not surprisingly, occur in the widespread Upper Greenstones subdivision of the Bulawayan Group with a high number of large producers around Que Que and Gatooma. Large deposits are notably absent from the thick lower tholeiitic pile, especially in the Gwelo, Fort Rixon-Shangani, Belingwe and Fort Victoria belts, and from the western calc-alkaline assemblage. Little is known of the regional zoning but several areas are noted for their associated minerals. Galena, for example, is common in the Umtali greenstone belt in the extreme east of the craton, whereas arsenopyrite is widespread in the Gwanda belt, 70 km south of Bulawayo. Even in the same greenstone belt, however, groups of mines may be characterized by different association of sulphide minerals. Some of the complexities are well shown by Stowe (1979) who, in the area around Que Que, recognizes the introduction of gold in three main phases. The first was a widespread pyrite or pyrite-arsenopyrite phase; the second was a zinc-copper-lead sequence over a small area; and the third, over an even smaller area, was an antimony mineralization. The association of gold with pre- or post-mineralization felsic dykes is apparent in a number of mines. There is also, in places, a close spatial relationship between large and small tonalite-granodiorite intrusions and the distribution of gold deposits. The emplacement of these intrusions is thought by many t o have produced the tectonically favourable areas, as well as the thermal control, for gold concentration (Anhaeusser, 1976). In several areas regional fracture patterns can be recognized (Stowe, 1968c; 1979). Around Que Que, Stowe (1979) notes that the directions of ore-body fissures are regionally consistent and strike north, northwest and northeast, respectively, for the early arsenical mineralization, with an additional eaststriking, flatdipping set related to the final antimony phase. Many of these fissure directions he relates t o major pre-mineralization “master shears” resulting from differential movement between rocks of different competencies. He suggests that these “master shears” might have formed during the final phase of intrusion of late granites such as the Sesombi tonalite, but goes on t o say that these “master shears” may be the result of craton-wide regional stress and related to the deformation envisaged by Coward et al. (1976a) during the main periods of Limpopo deformation. Another important mineral is chrysotile asbestos. This is mined in the
482 southern part of the craton with the bulk of production coming from the two major members of the Mashaba Ultramafic Suite at Shabani (Laubscher, 1964, 1968) and Mashaba (Wilson, 1968a, b), respectively. In the Mashaba area Wilson envisaged a relationship between intrusion of the Chilimani Suite granites, deformation, and fibre development. However, as is possible with Stowe’s “master shears” around Que Que, the fibre-producing deformation at Mashaba and Shabani can be related t o cratonic events during the Limpopo deformation as envisaged by Coward et al. (1976a) and the regional control seems better explained in terms of the cratonic southwesterly slide (Wilson, 1979). The country rock of the two intrusions is the crustal segment containing the major development of c. 3500 Ma rocks. Shortening in response t o the southwesterly slide can be regarded as one of limited block movement, internal t o this stable segment, and controlled largely by the position of the less competent “lubricants” of the ultramafic sheets. The fibre deposits are developed in the serpentinised olivine-rich ultramafic rocks close t o the major shears and thrusts (see also, p. 572). Important nickeliferous sulphide deposits of two types occur in the Rhodesian Basement Complex, associated with gabbro-metagabbro and ultramafic rocks, respectively. Unlike the gold mineralization, there does seem to be an age factor involved and, at least within the central cratonic area, the deposits appear t o be confined t o the Upper Greenstones (Wilson, 1979). Most seem explicable as part of the high-magnesium Reliance-type igneous activity at the base of the Upper Greenstones, but some nickel mineralization is also evident in ultramafic rocks of the overlying bimodal suite of the western succession, and in a differentiated mafic intrusion cutting the calc-alkaline volcanic assemblage. The status of the R4ashaba Ultramafic Suite in the nickel mineralization is not clear at this stage. High-grade stratiform chrornitite deposits are mined at Selukwe (Cotterill, 1969, 1979) and to a lesser extent at Mashaba (Wilson, 1968a) and chromitite occurrences are known in several of the Sebakwian Group remnants in the c. 3500 Ma gneisses separating these two centres. Early formed chromite is a feature of ultramafic rocks of all ages in the Rhodesian Archaean, ranging from the earliest Sebakwian t o the Great “Dyke”. However, if Cotterill’s interpretation of the chromitite-bearing serpentinite at Mashaba, as discussed earlier (and see also p. 572) is correct, then economic chromitite deposits in the Basement Complex are apparently confined to the Sebakwian Group. At the other end of the time scale beryllium-lithium pegmatite mineralization is a feature of some of the Chilimanzi Suite of late granites. PROBLEMS AND MODELS
Granitic rocks of three different approximate ages are now known from the Rhodesian Archaean, and all three would appear to involve sodic and potassic varieties. Two are associated with greenstone volcano-sedimentary
483 sequences; at c. 3500 Ma with the Sebakwian Group, and at c. 2700 Ma with the Upper Greenstones (Bulawayan) and Shamvaian Group. The age of the Lower Greenstones is still unknown but field evidence, and such isotopic data as there are, indicate that they are not part of the Sebakwian Group but rather something closer in time t o the Upper Greenstones. The recently recognized c. 2900 Ma granitic event may well be associated with the Lower Greenstones, if the Mashaba-type tonalites and Chingezi gneisses represent plutonism complementary t o their felsic volcanicity (Wilson et al., 1978; Wilson, 1979). The status of the Umniati River and Gwenoro gneisses remains obscure. They give ages of around 2800 Ma with low initial 87Sr/86Srratios (Table 8.111) but statistically the results are indistinguishable from the c. 2900 Ma granites and from the c. 2700 Ma granites and greenstones. Stowe’s (1968a, b, 1973) mapping supports a pre-Upper Greenstones age for the Gwenoro gneisses, and thus they are perhaps further phases of plutonism complementary to the Lower Greenstones. Alternatively, but less likely perhaps, they may be early manifestations of the Sesombi Suite, on the negative evidence of the apparent lack of Mashaba-Chibi dykes or Mashaba Ultramafic Suite within them. These dyke swarms and ultramafic intrusions, if they are coeval with the high-magnesium-tholeiitic volcanic pile or indeed any part of the Upper Greenstones, become an important tool in the delineation of the pre-Upper Greenstones basement. Taken in conjunction with the recent stratigraphic correlations they provide evidence for an extensive basement t o a widespread Upper Greenstones cover sequence, although more work is needed, both inside and outside the central cratonic area, t o delineate accurately the full extent of both. A near full spectrum of geological environments has now been proposed for the generation of Archaean greenstone belts (see Hunter, 1974; P’indley, 1976) and in the Zimbabwean context a number of models have been put forward. With the little that remains of the Sebakwian Group, it is difficult to evaluate its origin. The concordance of strike and dip of the greenstone belt remnants is such to preclude any conclusions regarding basement-cover relationships, although it is clear from the Selukwe succession that complex sialic crust was present at an early stage. Thus most models refer to the widespread main greenstone belts. Table 8.VI summarizes some of the more recent suggestions, with the omission of those which do not recognize the presence of an early sialic basement (e.g., Anhaeusser, 1973). Obviously only the most recent take cognisance of the implications of the Lower and Upper Greenstones subdivision of the Bulawayan Group, although the concept was to some extent inherent in Stowe (1971). The early models of Stowe (1971) and Wilson (1973b) both recognize a sialic basement containing infolded greenstone belt remnants. Wilson envisaged a much larger area of basement than did Stowe and saw in the
TABLE 8.VI Summary of recent models for greenstone belt generation in Zimbabwe Stowe (1971)
Bulawayan eugeosynclines o n northwest and southeast sides o f , and in fractures across, an older central protocraton consisting of MacGregor’s Shangani and Rhodesdale “batholiths” and some of country to east
Wilson (197313)
Northwest, north-northkast and east-northeast pattern o f mantle tapping fractures in early basement
Coward e t al. (1976b)
Southwest part of craton: triple junction of major, minor and failed spreading arms directed northwest, east-northeast and s o u t h s o u t h w e s t , respectively
Condie and Harrison (1976)
Midlands area: Bulawayan lower basaltic pile from depleted lherzolite source beneath spreading centre in marginal sea basin, sialic crust (Rhodesdale “batholith”) t o east. Eclogite sources for overlying andesites and dacites in descending slab west of basin
Hawkesworth and O’Nions ( 1 9 7 7 )
Upper Greenstones: rifting within a continental block for tholeiites and calc-alkaline rocks; derivation by differential partial melting and subsequent polybaric fractionation of essentially homogeneous source; Sesombi tonalite and late dacites from eclogite source
Katz (1976)
Development of greenstone belts of Rhodesian and Kaapvaal cratons in rift-ridge systems a t high angles t o , and initiated by, transform zone of Limpopo mobile belt: reactivation produced Great “Dyke” rift
Wilson 1 9 7 9 ; Wilson e t al. ( 1 9 7 8 )
Rifting confined to Rhodesian cratonic block: initiated by transcurrent movement in Limpopo mobile belt, (and possibly subparallel belt to north), t o explain Upper Greenstones, Great “Dyke” and possibly Lower Greenstones
three dominant greenstone belt trends of east-northeast, northeast and north-northeast, a reflection of an original pattern of mantle tapping fractures, which in turn he considered t o mark the sites of the thickest and heaviest development of volcanic cover. In this he was extending the general explanation of Anhaeusser et al. (1969) and attempting to produce a control in the downsagging required to produce Macgregor’s gregarious batholiths. He identified strike lengths of 500 km but offered no explanation for their generation.
485
In a more recent, plate tectonic, scheme Coward et al. (197613) also see some possible primary significance in these three directions. They point out that a triple junction model for the southwestern part of the craton, could explain the curvature of the greenstone belt outcrop pattern in this area; introduce a control in the east-northeast-trending major shears of the Limpop0 deformation; and explain the crustal shortening, which they suggest could represent a period of crustal collision which closed the original basin. While such a model might explain some of the structural features of this region, this southwestern area cannot be considered in isolation from the rest of the craton. Nowhere in the central area, for example, is there evidence that any part of the Upper Greenstones represents ocean floor; wherever the base of the Upper Greenstones is known to be preserved it rests on older granite-greenstone terrain. Similarly this is a weakness of the Condie and Harrison (1976) model which also envisages the lower basaltic pile as ocean floor. Their various tentative proposals are based on keochemical data and are made in the light of Harrison’s (1970) interpretation of the Que Que area stratigraphy. The chemistry is given in terms of average analyses for the various rock types and these, and the model, d o not distinguish between the Lower and Upper Greenstones. Their derivation of the calc-alkaline andesites, however, invokes subduction and an eclogite source. Hawkesworth and O’Nions (1977), on the other hand, argue against subduction for the generation of the calc-alkaline association. Their samples come almost entirely from the Upper Greenstones. They recognize two broad igneous series; an Fe-enrichment tholeiitic series (which includes peridotitic komatiites as defined by Brooks and Hart, 1974) and a calcalkaline series marked by andesites. Effectively they are dealing with samples from the Reliance and Zeederbergs formations at Belingwe; the equivalents of the Zeederbergs formation at Fort Victoria and Bulawayo; and the calcalkaline Maliyami formation of Harrison (1970) at Que Que. They maintain that, in marked contrast t o modern plate margins, the trace-element characteristics of the tholeiitic and most of the calc-alkaline rocks suggest derivation from chemically similar sources. They favour rifting within a continental block for both rock suites. In considering rifting models for greenstone belts it is interesting t o note that several workers have looked to the Great “Dyke” not only as an abortive rift but some have suggested, or implied, that the “Dyke” is an abortive greenstone belt (e.g., Coward et al., 1976; Katz, 1976; Wilson et al., 1979). This is in keeping also with the parental magma of the “Dyke” which was some kind of magnesium-rich tholeiite (p. 575). Thinking on these lines, and recourse t o some kind of rifting accompanying movements of major crustal blocks, form the basis of the most recent models. Katz (1976) suggests that the Limpopo mobile belt is a transform zone which, during its initiation, developed the greenstone belts of the Rhodesian
486 and Kaapvaal cratons in a rift-ridge system at high angles t o it, and that subsequent reactivation of this zone produced the rift of the Great “Dyke”. Wilson (1979) and Wilson et al. (1978) elaborate on a somewhat similar scheme but note that the Upper Greenstones subdivision of the main greenstone belts has not been recognized south of the Limpopo belt. These authors point out also that the approximate north-northeast line, so prominent in the Great “Dyke” fracture pattern, is a trend also evident in the present distribution of the western succession of Upper Greenstones, in particular in the calc-alkaline assemblage; and that it is, moreover, one reflected through time in the Rhodesian Archaean. Their suggestions for a possible rifting model invoke major intracratonic block movements involving transcurrent displacement along what is now the Limpopo mobile belt, and possibly also along a sub-parallel belt to the north. These movements were accompanied by craton-wide fractures on an approximate northeasterly trend, but confined to the Rhodesian block. For the Upper Greenstones this block was effectively the pre-2700 Ma Rhodesian craton, the full extent of which is not known. Similar even earlier movements, they suggest, might have produced the Lower Greenstones. By 2500 Ma, however, in perhaps a last but unsuccessful attempt to produce a greenstone belt, the Rhodesian craton, finally stabilized by the late granites of the extensive Chilimanzi suite, fractured sufficiently to produce only the Great “Dyke” pattern. Their speculations imply that the greenstone belt tectonics were intracratonic but not that all Archaean tectonics were such. In summary, present knowledge does not allow conclusions t o be drawn regarding a final model for the generation of the Zimbabwean greenstone belts. On any modern analogy the tholeiitic and calc-alkaline provinces as recorded in the Upper Greenstones would correspond to the constructive and destructive phases of a plate cycle. Bearing in mind the evidence for a widespread sialic basement, perhaps the closest modern analogy for the lower formations is in rifted continental areas; but it is not clear from the Zimbabwean evidence whether a regime producing rifting in the Archaean also produced major ocean basins elsewhere, as happened in later geological time (Wilson et al., 1978). In contrast, although the full western extent of the calc-alkaline assemblage is obscured by younger rocks, on present exposures it has a linear development; and thus the present areal distribution of the younger volcanic rocks of the Upper Greenstones is reminiscent of modern destructive continental margins. Whether this necessitates an origin in a regime analogous to present-day plate tectonics involving subdunction is still debatable, and highlighted by the conflicting views expressed by the, albeit preliminary, geochemical studies. Recent thinking reflects the increasing tendency to seek an interrelationship of events in the craton and adjacent mobile belts to explain generation of greenstone belts, regional structures and metamorphism. If ideas concern-
TABLE 8.VII Summary of major events in the Rhodesian Archaean craton (After Wilson, 1979) Ma 2500-
Great “Dyke” WEST Sesombi Suite
2600-
Chilimanzi Suite granites
2700-
Sesombi Suite granites Shamvaian Group Upper Greenstones (Bulawayan Group)
2800-
Shamvaian G r o 3
NNIVNNNZNN
I
Calc-a1kaline Mafidhigh Mg‘2 felsic bimodal TholeiiticHigh Mg
EAST Shamvaian GroujNNNNlVrVNrVlV
- Mafic
(some high Mg) sediments
- _ _ - _ _ _ - ---- Tholeiitic High-Mg
Mashaba-Chibi dykes and Mashaba Ultramafic Suite
29003000-
Mashaba tonalite ? Lower Greenstones - - - - - - - - - - - - - (Bulawayan Group)
Mafic/high Mg-felsic bimodal volcanism; sediments
3100320033003400-
Mont d’Or Formation granite
3500-
Mushandike granite Selukwe nappe Sebakwian Group and granites - - - - - - - - - -
3600-
~ ~ f ~ c - v ~ ~ ~ n ~ s mafic and ultramafic v o l c a n i a ZlGranites and intrusion b w
~
~
~
e
~
~
m lb
m
4
~
~
~
488 ing major block movements are correct then such movements, whatever their fundamental cause, have had a major controlling influence on the evolution of the craton from possibly 3000 Ma-2500 Ma. Any final model for the Zimbabwean greenstone belts, however, must recognize the likely presence of granite-reenstone terrains of three different ages within the cratonic area, and, hopefully, explain all three. Table 8.VII summarizes the major events in the Rhodesian Archaean craton. C. Angola
*
INTRODUCTION
Angola has a surface area of over 770,000 km2 but, geologically, approximately 65% of the country is covered by continental Upper Cretaceous formations as well as vast expanses of Tertiary sediments - mainly Kalahari Beds. These younger formations effectively obscure all but the western third of the territory where Precambrian rocks of varying age are exposed in a n o r t h s o u t h belt over 1250 km long and averaging approximately 400 km wide. Very little of Angola has been mapped geologically and extensive areas remain to be prospected and explored. Geochronological studies undertaken by Cahen and Snelling (1966) have delimited areas of varying ages and these, combined with the subdivision of the region into minero-geographic provinces, will be used to describe the geological characteristics of the country. THE EARLY PRECAMBRIAN (-3300-1700
Ma)
The Kasai craton, which is centred in Zaire, extends across the northern third of Angola. Ages of approximately 3300 Ma have been recorded in the region north of the Angolan border (Cahen and Snelling, 1966), but in northwestern Angola available ages fall within a 2700-1800 Ma range. From north to south across northeastern Angola Cahen and Snelling (1966) have distinguished the following main geological complexes: (1)the Upper Luanyi Gneiss; (2) the Dibaya granitemigmatite complex; (3) the Kasai-Lomami noritic-gabbro and charnockite complex; (4)the Luiza metasedimentary complex; (5) the Lulua sedimentary and volcanic complex; and (6) a mainly migmatitic and granitic complex south of 8" S. The Upper Luanyi Gneiss appears to be the oldest assemblage (-3300
* By C.R. Anhaeusser.
489 Ma). This is followed by the Dibaya granite-migmatite complex (-2700 Ma) and the charnockitic cycle (2650-2500 Ma) which followed the deposition of a poorly documented succession of sediments and the intrusion of noritic gabbro, granites and dolerites (Kasai-Lomami complex). The sedimentary sequence, which has a basal unconformity, has been involved in migmatization, charnockitization, deformation, and uplift, accompanied by pegmatite invasion. The Luiza metasedimentary assemblages rest unconformably on the Kasai-Lomami rocks (2120-1910 Ma) and the Lulua sediments and volcanics overlie the Dibaya granites and migmatites but have not been dated. In northwestern Angola the granitic basement, which appears in places beneath the West Congo geosyncline, is considered t o be 2500 Ma or even older (Cahen and Snelling, 1966). No data are available from northwest Angola itself, but age relations have been extrapolated southwards from the neighbouring Congo Republic and from Gabon (Zanidian orogeny - 2500 Ma or older). In central and southwestern Angola vast areas of the country are underlain by the socalled “Metamorphic Series”, for which there are numerous age determinations in the 1690-1400 Ma range (Cahen and Snelling, 1966). Mapping and age determinations carried out by De Carvalho (1969) indicate that the Kunene anorthosite complex may be as old as 2080 Ma. This complex intrudes the “Metamorphic Series’’ which consists of schists, metaquartzites and amphibolites and is overlain by crystalline limestones and gneisso-migmatitic rocks. A wide variety of granitic and porphyritic granitoid bodies also occur in the area and these range in age from +2080 Ma t o approximately 925 Ma. The anorthosite complex of southern Angola extends southward across the Kunene River into South West Africa and is the largest known mass of its type, having an exposed outcrop area of 12,500 km2 plus an estimated additional 4000 km2 concealed beneath younger sediments (Simpson, 1970). It is exposed as a longitudinally elongated massif over a distance of 300 km, and varies in width between 30 and 100 km. The predominant rock type exposed in the southern half of the complex comprises pale, massive, leucocratic, anorthosite, made up mainly of plagioclase (An,,,-An,,) together with minor amounts of orthopyroxene and/or clinopyroxene as well as olivine. The presence of olivine and pyroxene suggests that types transitional to troctolitic anorthosites are developed in some areas. Dark troctolitic anorthosites predominate in the northern part of the complex. The rocks are dark grey, almost black in colour and are further characterized by the tabular habit of the plagioclase crystals (average grain size usually 10-15 mm, varying up to 10 cm), the common development of igneous lamination, as well as faintly discernible rhythmic layering in some of the coarsely banded rocks. Mafic and ultramafic rocks appear t o be confined t o the border zones of
490 the anorthosite complex and comprise norites, gabbros, troctolite, pyroxenite and serpentinite. Titaniferous magnetite is a constant accessory mineral in the troctolitic anorthosite but in some places it is a major constituent making up approximately 20% of the rock. Random samples of massive ore indicated the following compositional ranges: Fe, 49.4-52.3%; Ti02, 13.419.7%;VzOs, 0.4-0.796 (Simpson, 1970). THE MIDDLE PRECAMBRIAN (- 1700-850 Ma)
The Zanidian orogeny (2500 Ma and older) in northwestern Angola, is believed by Cahen and Snelling (1966) t o be the older of two pre-West Congolean orogenies; the younger being referred t o as the Mayumbian orogeny (1950-1100 Ma). This interpretation differs from that of Stanton et al. (1963) who regarded the Mayumbian as a metamorphosed equivalent of the Sansikwa series of t h e West Congolean geosyncline. In central Angola a vast unmapped granitic terrane has yielded ages that cluster into a 1690-1400 Ma-old group and a 1280-925 Ma-old group. Very little is known about the rocks in this area but they are probably related t o similar rocks in southwestern Angola discussed later. Also in central Angola is a sequence of undated sediments referred to as the Oendeolongo System (Mouta, 1954). This succession, which has a lower argillaceous group consisting of a basal conglomerate, followed by feldspathic sandstones, greywacke and schists, and an upper siliceous group, consisting of banded iron-formations, quartzites and sandstones, occurs in scattered outliers resting, in places, unconformably on the granitic basement (Hall, 1949) but also being locally intensely tectonically disturbed (often t o the vertical). Igneous rocks (intrusive or extrusive?) consisting of dolerites and rhyolites or rhyolite tuffs are also present in the area. Mouta (1954) correlated the Oendeolongo System rocks with those of the pre-West Congolean Mayumbian in northwestern Angola but this interpretation is open t o doubt. In southwestern Angola the “Metamorphic Series” mentioned earlier shows a wide range of ages from approximately 2 0 8 0 4 2 5 Ma. Leucogranites, dated at -1700 Ma flank the Kunene anorthosite complex and are the equivalents of the Fransfontein granite in South West Africa described by Clifford et al. (1969) and which represents part of the Huabian Episode of granite emplacement of batholitic dimensions. A later cycle in the area is represented by the early Kibaran igneous event which is manifested by the intrusion of two varieties of generally red porphyritic granites (1100 f 200 Ma) and which was presumably followed by the Oendeolongo (Kibaran?) sedimentation episode. T H E LATE PRECAMBRIAN (-850-500
Ma)
Rocks falling into this time period are represented in Angola by the West Congo System which is a geosynclinal sequence that has been correlated
491 TABLE 8.VIII Succession of the West Congo System - N.W. Angola (modified after Cahen and Snelling, 1966)
L a f e cycle (circa 600-500 Ma) Schisto Greseux Schisto-calcaire Tillite superieure Haut Shiloango-Louila Tillite inferieure
arenites, pelites (some red beds), disconformities, unconformities carbonate and pelite, some arenites (upper tillite) tillite, disconformity, slight unconformity pelites, carbonate rocks, arenaceous units, disconformity (lower tillite) tillite, basaltic
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . OROGENESIS 7 00-600 Ma DISCONFORMITY-UNCONFORMITY
. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Early c y c l e (circa 850-700 Ma) Sansikwa-Bembe carbonate, pelitic and arenaceous rocks
with the Damara-Katanga (Clifford, 1970). The original trough of deposition extended from South West Africa to Gabon but preservation, in Angola, of segments of the initial succession are only found in the southwest and northwest of this territory. Offshore developments of this sequence are also known t o extend for more than 600 km south of Luanda. In northwest Angola, the type locality of the West Congo System occurs at the southern end of the 1200 km by 350 km Angola-Gabon trough. Here the succession, which consists of a variety of pelitic sediments as well as carbonate rocks and tillites (Table 8.VIII), attains a thickness of approximately 1 4 km. Orogeny caused a broad belt of parallel NW-SE folds which increase in intensity towards the west, culminating in a zone of thrusting along which the pre-geosynclinal basement successions override the West Congolean formations. Eastwards the folding dies out progressively. Faulting appears prominent in southwestern Angola and seemingly the West Congo geosyncline was faulted offshore. Regional metamorphism increases westwards with the folding where largely post-tectonic granitic and other igneous rocks were intruded along the western margin of the West Congo System. MINERALIZATION
In the Early Precambrian successions, which occupy only about 7% of the exposed surface area of Angola, mineral deposits that have been exploited include occurrences of mica, barytes, lead and zinc. Minor deposits of asbestos, beryl, quartz, graphite, and uranium are known. The Middle Precambrian formations occupy approximately 24% of the land area of Angola and have numerous occurrences of iron-ore, uranium,
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manganese, gold and copper. Minor deposits include tungsten, barytes, phosphates, kaolin, beryl, lead, zinc, graphite, silver and tin. Most of the iron-ores are titaniferous and are related t o basic or ultrabasic rocks but some reputedly high-grade hematitic ores, presumably of metasedimentary origin (being associated with carbonate rocks), occur in central Angola (Hall, 1949). Manganese deposits occur in the region east of Luanda where psilomelane admixed with manganite and braunite occurs in a coarse feldspathic grit, presumably of Late Precambrian age (Hall, 1949). Oxidized ores comprising concretionary pyrolusite and earthy wad occur in places and elsewhere narrow seams of rhodonite have also been reported. Most of the manganese orebodies have resulted from the metamorphism of manganiferous sediments and grades of up to 50% Mn have been mined. Gold deposits in Angola are almost without exception of alluvial type. According t o Hall (1949) most of the deposits occur in rivers taking their source in areas occupied by Oendeolongo rocks which survive as roof pendants and as scattered and much denuded outliers overlying the granitic basement rocks in southcentral Angola. Copper mineralization, also of Late Precambrian age, appear mainly in shears and fractures in a number of localities in northwest and southern Angola. The copper deposits in northwest Angola are associated with grits, sandstones, siltstones, limestones and tillites in the Bembe area of the West Congo geosyncline. Supergene carbonate ores as well as syngenetic sulphide ores have been mined a t numerous localities. REFERENCES Part A Allsopp, H.L., 1961. Rb-Sr age measurements on total rock and separated-mineral fractions from the Old Granite of the central Transvaal. J. Geophys. Res., 66: 1499-1508. Allsopp, H.L., 1964. Rubidium-strontium ages from the western Transvaal. Nature, London, 204 (4956): 351-3
