Precambrian of the Southern Hemisphere

Developments in Precambrian Geology 2 PRECAMBRIAN OF THE SOUTHERN HEMISPHERE DEVELOPMENTS I N PRECAMBRIAN GEOLOGY Adv...

Author: Donald Raymond Hunter

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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.

.... ... . . .. . ..... ...

. .... .. .. .. .... . . .. . . .. . . . .... .. . . .. ... . . .. . . . . . .. . ... . .. .. .. .. . ..... .. . . . .. ..

THE PALAEONTOLOGICAL RECORD E. Suszczynski . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

.

.

.

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.

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(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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98 Finucane, K.J., 1953. The geological structure of northwestern and central Western Australia in relation t o mineralization. In: A.B. Edwards (Editor), Geology of Australian Ore Deposits. Australas Inst. Min. Metall., Melbourne, pp. 188-194. Fitton, M.J., Horwitz, R.C. and Sylvester, G., 1975. Stratigraphy of the Early Precambrian in the west Pilbara, Western Australia. CSIRO Min. Res. Lab. Div. Mineral. Rept. FP 11.

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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.

101 Morris, R.C., 1973. A pilot study of phosphorus distribution in parts of the Brockman iron formation, Hamersley Group, Western Australia. West. Aust. Geol. Surv. Annu. Rep., 1972: 75-81. Muhling, P.C. and De Laeter, J.R., 1971. The age of granitic rocks in the Poona-Dalgaranga area of the Yilgarn Block, Western Australia. West. Aust. Geol. Surv. Annu. Rep., 1970: 8-73. Muhling, P.C. and Low, G.H., 1973. Explanatory notes on the Yalgoo 1 : 250,000 geological sheet, Western Australia. Geol. Surv. West. Aust. Rec. 1973/6 (unpubl.). Nance, W.B. and Taylor, S.R., 1977. Rare earth element patterns and crustal evolution 11. Archaean sedimentary rocks from Kalgoorlie, Australia. Geochim. Cosmochim. Acta, 41 : 225-231. Nesbitt, R.W., 1971. Skeletal crystal forms in the ultramafic rocks of the Yilgarn Block, Western Australia: Evidence for an Archaean ultramafic liquid. Spec. Publ. Geol. Soc. Aust., 3: 331-347. Nesbitt, R.W. and Sun, S.S., 1976. Geochemistry of Archaeanspinifex-textured peridotites and magnesian and low-magnesian tholeiites. Earth Planet. Sci. Lett., 31 : 433-453. Noldart, A.J. and Wyatt, J.D., 1962. The geology of portion of the Pilbara Goldfield covering the Marble Bar and Nullagine 4-mile map sheets. Geol. Surv. West. Aust. Bull., 115. O’Beirne, W.R., 1968. The Acid Porphyries and Porphyroid Rocks of the Kalgoorlie Area. Ph.D. Thesis, University of West. Australia (unpubl.). O’Driscoll, E.S.T., 1971. The tectonic pattern of the nickel province of southwestern Australia. Soc. Mining Geol. Japan Spec. Issue, 3 : 23-33. Oversby, V.M., 1975. Lead isotope systematics and ages of Archaean acid intrusives in the Kalgoorlie-Norseman area, Western Australia. Geochim. Cosmochim. Acta, 39: 1107-1125. Oversby, V.M., 1976. Isotopic ages and geochemistry of Archaean acid igneous rocks from the Pilbara, Western Australia. Geochim. Cosmochim. Acta, 40: 817-829. Page, R.W., 1976. Geochronology of Archaean and lower Proterozoic rocks in the Rum Jungle-Alligator River area, Northern Territory, Australia. 25th Int. Geol. Congr. Sydney, 1976, Sect. 1: 18-19. Pearce, J.A., 1976. Statistical analysis of major element patterns in basalts. J. Petrol., 1 7 : 15-43. Pidgeon, R.T., 1978. 3450 m.y.-old volcanics in the Archaean layered greenstone Succession of the Pilbara Block, Western Australia. Earth Planet. Sci. Lett., 37: 421-428. Prider, R.T., 1948. Igneous activity, metamorphism and ore formation in Western Australia. J.R. SOC.West. Austr., 31: 43--84. Prider, R.T., 1952. Southwest Yilgarnia. In: Sir Douglas Mawson Anniversary Volume, University of Adelaide, pp. 142-151. Prider, R.T., 1965. Geology and mineralization of the Western Australian Shield. In: J. McAndrew (Editor), Geology of Australian Ore Deposits. Australas. Inst. Min. Metall., Melbourne, pp. 56-65. Purvis, A.C., Nesbitt, R.W. and Hallberg, J.A., 1972. The geology of part of the Carr Boyd Rocks complex and its associated nickel mineralization, Western Australia. Econ. Geol., 6 7 : 1093-1113. Richards, J.R., Berry, H. and Rhodes, J.M., 1966. Isotopic and lead alpha ages on some Australian zircons. J. Geol. Soc. Austr., 13: 69-96. Roddick, J.C., Compston, W. and Durney, D.W., 1975. The radiometric age of the Mount Keith Granodiorite, a maximum age estimate for an Archaean greenstone sequence in the Yilgarn Block, Western Australia. Precambrian Res., 3 : 55-78. Ross, J.R. and Hopkins, G.M.F., 1975. Kambalda nickel sulphide deposits. In C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5: 100-121.

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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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105

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 ~~~

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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

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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__.

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71

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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

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.

.

.

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.

.

.

.

.

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Psammopelite 8 pelite Potosi type gneiss Lode horlzon 8 ore lenses Potosi

type gneiss

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Pscirnrnopslite 8 psammite

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. . . .

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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.

Peat, C.J., Muir, M.D., Plumb, K.A., McKirdy, D.A. and Nowick, M.S., 1978. Proterozoic microfossils from the Roper Group, Northern Territory, Australia. BMR J. Aust. Geol. Geophys., 3 : 1-17. Plumb, K.A. and Derrick, G.M., 1975. Geology of the Proterozoic rocks of the Kimberley t o Mount Isa region. In: C.L. Knight (Editor), Economic Geology of Australia and Papua New Guinea. 1. Metals. Australas. Inst. Min. Metall. Monogr., 5: 217-252. Plumb, K.A. and Sweet, I.P., 1974. Regional significance of recent correlations across the Murphy Tectonic Ridge, Westmoreland area. In: Recent technical and social advances in the North Australian minerals industry. Australas. Inst. Min. Metall., NW Queensland Branch, pp. 199-206. Plumb, K.A., Shergold, J.H. and Stefanski, M.Z., 1976. Significance of Middle Cambrian trilobites from Elcho Island, Northern Territory. BMR J. Aust. Geol. Geophys., 1 : 51-55. Preiss, W.V., 1972. The systematics of South Australian Precambrian and Cambrian stromatolites. Part 1. Trans. R. Soc. South Aust., 96: 67-100, 1 5 figs. Preiss, W.V., 1976a. Intercontinental correlations. In: M.R. Walter (Editor), Stromatolites. Amsterdam, Elsevier, pp. 359-370. Preiss, W.V., 1976b. Proterozoic stromatolites from the Nabberu and Officer Basins, Western Australia and their biostratigraphic significance. Geol. Surv. South Aust. Rep. Invest., 47: 1-51. Preiss, W.V., 1977. The biostratigraphic potential of Precambrian stromatolites. Precambrian Res., 5: 207-219. Preiss, W.V. and Walter, M.R., 1975. Stromatolites of the Australian Precambrian: Use in intra- and intercontinental correlations. Int. Geol. Correlation Progr., Symp. “Correlation of the Precambrian”, Moscow, Abstr., pp. 74-75. Randal, M.A., 1966. Brunette Downs, N.T. Explan. Notes, 1 : 250,000 Geol. Ser., Bur. Miner. Resour. Geol. Geophys. Aust., SE/53-11, 16 pp. Richter, R., 1955. Die altesten Fossilien Sud-Afrikas. Senckenbergiana Lethaea, 36 : 243289. Rudwick, M.J.S., 1964. The infra-Cambrian glaciation and the origin of the Cambrian fauna. In: A.E.M. Nairn (Editor), Problems of Palaeoclimatology. Interscience, London, pp. 150--155,184-185. Schopf, J.W., 1968. Microflora of the Bitter Springs Formation, Late Precambrian, Central Australia. J. Paleontol., 42: 6 5 1 - 6 8 8 . Schopf, J.W., 1970. Precambrian micro-organisms and evolutionary events prior t o the origin of vascular plants. Biol. Rev., 45: 319-352. Schopf, J.W., 1972. Evolutionary significance of the Bitter Springs (Late Precambrian) microflora. Int. Geol. Congr. 24th Montreal, Proc. Sect., 1: 68-77. Schopf, J.W., 1975a. Biostratigraphic usefulness of Precambrian microbiotas. Int. Geol. Correlation Progr., Symp. “Correlation of the Precambrian”, Moscow, Abstr., pp. 3940. Schopf, J.W., 1975b. Precambrian palaeobiology: problems and perspectives. Annu. Rev. Earth Planet. Sci., 3 : 213-249. Schopf, J.W., 1976. Are the oldest “fossils”, fossils? Origins Life, 7: 19-36. Schopf, J.W. and Barghoorn, E.S., 1969. Microorganisms from the late Precambrian of South Australia. J. Paleontol., 43: 111-118. Schopf, J.W. and Blacic, J.W., 1971. New microorganisms from the Bitter Springs Formation (Late Precambrian) of the North-Central Amadeus Basin, Australia. J. Paleontol., 45: 925-960. Schopf, J.W. and Fairchild, T.R., 1973. Late Precambrian microfossils: a new stromatolitic microbiota from Boorthanna, South Australia. Nature, 242 : 537-538. Schopf, J.W. and Oehler, D.Z., 1976. How old are the eukaryotes? Science, 193: 47-49.

396 Schopf, J.W., Haugh, B.N., Molnar, R.E. and Satterthwait, D.F., 1973. On t h e development of metaphytes and metazoans. J. Paleontol., 4 7 : 1-9. Siegel, S.M. and Siegel, B.Z., 1968. A living organism morphologically comparable t o the Precambrian genus Kakabekia. Am. J. Bot., 5 5 : 684+87. Smith, J.W. and Croxford, N.J.W., 1 9 7 3 . Sulphur isotope ratios in t h e McArthur leadzinc-silver deposit. Nature Phys. Sci., 245: 10-12. Sprigg, R.C., 1947. Early Cambrian ( ? ) jellyfishes from t h e Flinders Ranges, South Australia. Trans. R. SOC.South Aust., 71: 212-224. Timofeev, B.V., 1 9 6 6 , 1974. Mikropaleofitologicheskoe issledovanie drevnikh svit. Moscow, Nauka, 1 4 7 pp. 89 pls. Translation published as: Micropalaeontological research into ancient strata, British Library Lending Division, 214 pp. 8 9 pls. (1974). Wade, M., 1968. Preservation of soft-bodied animals in Precambrian sandstones a t Ediacara, South Australia, Lethaia, 1: 238-267. Wade, M., 1969. Medusae f r o m uppermost Precambrian o r Cambrian sandstones, Central Australia. Palaeontology, 1 2 : 351-365. Wade, M., 1970. The stratigraphic distribution of t h e Ediacara fauna in Australia. Trans. R. SOC.South Aust.. 9 4 : 87-104. Walter, M.R., 1970. Stromatolites and t h e Biostratigraphy of t h e Australian Precambrian. Ph.D. Thesis, University of Adelaide, 466 pp. (unpubl.). Walter, M.R., 1 9 7 2 . Stromatolites and the biostratigraphy of the Australian Precambrian and Cambrian. Spec. Pap. Palaeont., Palaeontol. Assoc. London, 1 1 : 1 9 0 pp. Walter, M.R., 1976. Introduction. In: M.R. Walter (Editor), Stromatolites. Elsevier, Amsterdam, pp. 1-3. Walter, M.K., 1977. Interpreting stromatolites. Am. Sci., 6 5 : 563-571. Walter, M.R., Bauld, J . and Brock, T.D., 1 9 7 2 . Siliceous algal and bacterial stromatolites in h o t spring and geyser effluents of Yellowstone National Park. Science, 1 7 8 : 402405. Walter, M.R., Bauld, J. and Brock, T.D., 1976a. Microbiology and morphogenesis of columnar stromatolites (Conophyton, Vacerrilla) f r o m h o t springs in Yellowstone National Park. In: M.R. Walter (Editor), Stromatolites. Elsevier, Amsterdam, pp. 273310. Walter, M.R., Goode, A.D.T. and Hall, W.D.M., 1976b. Microfossils from a newly discovered Precambrian stromatolitic iron formation in Western Australia. Nature, 26 1: 22 1-223.

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,

492

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): 351-363. Allsopp, H.L., Roberts, H.R., Schreiner, G.D.L. and Hunter, D.R., 1962. Rb-Sr age measurements on various Swaziland granites. J. Geophys. Res., 67( 13): 5307-5313. Allsopp, H.L., Ulrych, T.J. and Nicolaysen, L.O., 1968. Dating some significant events in the history of the Swaziland System by the Rb-Sr isochron method. Can. J. Earth Sci., 5 : 6 0 5 4 1 9 . Allsopp, H.L., Viljoen, M.J. and Viljoen, R.P., 1973. Strontium isotopic studies of the mafic and felsic rocks of the Onverwacht Group of the Swaziland Sequence. Geol. Rundsch., 63(3): 902-917. Anhaeusser, C.R., 1973a. The evolution of the early Precambrian crust of southern Africa. Philos. Trans. R. SOC.London, Ser. A, 273: 359-388. Anhaeusser, C.R., 197313. The geology and geochemistry of the Archaean granites and gneisses of the Johannesburg-Pretoria dome. Spec. Publ. Geol. SOC.S. Afr., 3: 361385. Anhaeusser, C.R., 1975. Precambrian tectonic environments. In: F.A. Donath (Editor), Annu. Rev. Earth Planet. Sci. Annual Reviews Inc., California, 3: 31-53.

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Part B Anhaeusser, C.R., 1973. The evolution of the early Precambrian crust of southern Africa. Philos. Trans. R. SOC.London. Ser. A, 273: 359-388.

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497 Hawkesworth, C.J., Bickle, M.J., Gledhill, A.R., Wilson, J.F. and Orpen, J.L., 1977. A 2.9-b.y. event in the Rhodesian Archaean. Earth Planet. Sci. Lett., 43: 285-297. Hickman, M.H., 1974a. 3500 m.y.-old granite in southern Africa. Nature, 251: 295296. Hickman, M.H., 1974b. Geochronological study of the Rhodesian craton and the Limpop0 belt. Univ. Leeds Res. Inst. Afr. Geol. Annu. Rep., 18: 6 1 - 6 2 . Hickman, M.H., 1976. Geochronological Studies in the Limpopo Mobile Belt and Part of the Adjacent Rhodesian Craton. Ph.D. Thesis, Univ. Leeds, England (unpubl.). Hunter, D.R., 1974. Crustal Development in the Kaapvaal Craton, I. The Archaean. Precambrian Res., 1: 259-294. Jahn, B.M. and Condie, K.C., 1976. On the age of Rhodesian greenstone belts. Contrib. Mineral. Petrol., 57, 317-330. Katz, M.B., 1976. Early Precambrian granulites-greenstones, transform mobile belts and ridge-rifts in early crust?In: B.F. Windley (Editor), The Early History of the Earth. Wiley, New York, N.Y., pp. 147-155. Key, R.M., Litherland, M. and Hepworth, J.V., 1976. The evolution of the Archaean crust of northeast Botswana. Precambrian Res., 3: 375-413. Laubscher, D.H., 1964. The occurrence and origin of chrysotile asbestos and associated rocks, Shabani, Southern Rhodesia. In: S.H. Haughton (Editor), The Geology of Some Ore Deposits of Southern Africa. Geol. SOC.S. Afr., pp. 5 9 3 - 6 2 4 . Laubscher, D.H., 1968. The origin and occurrence of chrysotile asbestos in the Shabani and Mashaba areas, Rhodesia. Trans. Geol. SOC.S . Afr., 71, Annex: 195-204. Litherland, M., 1973. Uniformitarian approach to Archaean “Schist Relics”. Nature, Phys. Sci., 242: 125-127. Macgregor, A.M., 1937. The geology of the country around Hunters Road, Gwelo District. South. Rhod. Geol. Surv. Bull., 31: 78 pp. Macgregor, A.M., 1941. A Precambrian algal limestone in southern Rhodesia. Trans. Geol. SOC.S. Afr., 43: 9-15. Macgregor, A.M., 1947. An outline of the geological history of Southern Rhodesia. South. Rhod. Geol. Surv. Bull., 38: 7 3 pp. Macgregor, A.M., 1951. Some milestones in the Precambrian of Southern Rhodesia. Proc. Geol. Soc. S. Afr., 54: 27-71. Martin, A., 1978. The geology of the Belingwe-Shabani schist belt. Rhod. Geol. Surv. Bull., 83: 220 pp. Maufe, H.B., Lightfoot, B. and Zealley, A.E.V., 1919. T h e geology of the Selukwe mineral belt. South. Rhod. Geol. Surv. Bull., 3 : 96 pp. Moorbath, S., Wilson, J.F. and Cotterill, P., 1976. Early Archaean age for t h e Sebakwian Group at Selukwe, Rhodesia. Nature, 264: 536-538. Moorbath, S., Wilson, J.F., Goodwin, R. and Humm, M., 1977. Further Rb-Sr age and isotope data o n early and late Archaean rocks from the Rhodesian craton. Precambrian Res., 5 : 229-239. Nisbet, E.G., Bickle, M.J. and Martin, A,, 1977. The mafic and ultramafic lavas of the Belingwe greenstone belt, Rhodesia. J. Petrol., 18(3): 521-566. Orpen, J.L., 1976. Discussion of “Some aspects of the Archaean history of southern Rhodesia” by M.P. Coward and R.M. Shackleton. Univ. Leeds Res. Inst. Afr. Geol. Annu. Res., 20: 28-29. Orpen, J.L., 1978. The Geology of the Southwestern Part of the Belingwe Greenstone Belt, and Adjacent Areas. D. Phil. Thesis Univ. Rhodesia, Salisbury (unpubl.). Phaup, A.E., 1964. Gold Mines of Southern Rhodesia - Introduction. In: S.H. Haughton (Editor), The Geology of Some Ore Deposits of Southern Africa. Spec. Publ. Geol. SOC.S. Afr., 2: 1-7. Phaup, A.E., 1973. Chemical Analyses of Rocks, Ores and Minerals of Rhodesia. Rhod. Geol. Surv. Bull., 71: 297 pp.

498 Robertson, D.K., 1973. A model discussing the early history of the earth based on the study of lead isotope ratios from veins in some Archaean cratons in Africa. Geochim. Cosmochim. Acta, 37: 2099-2124. Robertson, I.D.M., 1973a. Potash granites of the southern edge of the Rhodesian Craton and North Marginal Zone of the Limpopo Mobile Belt. Spec. Publ. Geol. SOC.S. Afr., 3 : 265-276. Robertson, I.D.M., 1973b. The geology of the country around Mount Towla, Gevanda District. Rhod. Geol. Surv. Bull., 68: 166 pp. Robertson, I.D.M., 1976. T h e geology of t h e country around Battlefields, Gatooma District. Rhod. Geol. Surv. Bull., 76: 258 pp. Saggerson, E.P. and Turner, L.M., 1976. A review of the distribution of metamorphism in the ancient Rhodesian craton. Precambrian Res., 3: 1-53. Snowden, P.A., 1976. The Geology of the Granitic Terrain North and East of Salisbury with Particular Reference t o the Chinamora Batholith. D. Phil. Thesis Univ. Rhodesia, Salisbury (unpubl.). Snowden, P.A. and Bickle, M.J., 1976. The Chinamora batholith: diapiric intrusion or interference fold? J. Geol. SOC.Lontlon, 132: 131-137. Stagman, J.G., 1977. Provisional geological map of Rhodesia 1 : 1,000,000, 7th Ed. Rhod. Geol. Surv. Salisbury. Stagman, J.G., 1978. An outline of the geology of Rhodesia. Trans. Geol. Soc. S. Afr., 80: 126. Stowe, C.W., 1968a. T h e geology of the country south and west of Selukwe. Rhod. Geol. Surv. Bull., 59: 209 pp. Stowe, C.W., 1968b. Intersecting fold trends in the Rhodesian Basement south and west of Selukwe. Trans. Geol. Soc. S. Afr., 71, Annex: 53-78. Stowe, C.W., 1968c. The fissure pattern and mineralisation in the Mont d’Or area, Selukwe, Rhodesia. Trans. Geol. SOC.S. Afr. 7 1 (Annex): 91-102. Stowe, C.W., 1971. Summary of the tectonic development of the Rhodesian Archaean Craton. Spec. Publ. Geol. SOC.Aust., 3 : 377-387. Stowe, C.W., 1973. The older tonalite gneiss complex in the Selukwe area Rhodesia. Spec. Publ. Geol. SOC.S. Afr., 3: 85-95. Stowe, C.W., 1974. Alpine-type structures in the Rhodesian basement complex at Selukwe. J. Geol. SOC.London, 130: 411-425. Stowe, C.W., 1979. Gold and associated mineralisation in the Que Que area, Rhodesia. Spec. Publ. Geol. SOC.S. Afr., 5: 39-48. Viljoen, M.J. and Viljoen, R.P., 1969. A reappraisal of granite-greenstone terrains in shield areas based on the Barberton model. Spec. Publ. Geol. SOC.S. Afr., 2: 245-274. Wiles, J.W., 1971. Provisional geological map of Rhodesia. 1 : 1,000,000, 6th Ed. Rhod. Geol. Surv. Salisbury. Windley, B.F., 1976. New tectonic models for the evolution of Archaean continents and oceans. In: B.F. Windley (Editor), The Early History of the Earth. Wiley, New York, N.Y., pp. 105-111. Wilson, J.F., 1964. T h e geology of the country around Fort Victoria. South. Rhod. Geol. Surv. Bull., 58: 147 pp. Wilson, J.F., 1968a. The geology of the country around Mashaba. Rhod. Geol. Surv. Bull., 6 2 : 239 pp. Wilson, J.F., 1968b. The Mashaba Igneous Complex and its subsequent deformation. Trans. Geol. SOC.S. Afr., 71, Annex: 175-188. Wilson, J.F., 1973a. Granites and gneisses of the area around Mashaba, Rhodesia. Spec. Publ. Geol. SOC.S. Afr., 3: 79-84. Wilson, J.F., 1973b. The Rhodesian Archaean craton - an essay in cratonic evolution. Philos. Trans. R. SOC.London. Ser. A. 273: 389-411.

Wilson, J.F., 1979. A preliminary reappraisal of the Rhodesian Basement Complex. Spec. Publ. Geol. SOC.S. Afr., 5: 1-23. Wilson, J.F. and Harrison, N.M., 1973. Recent K-Ar age determinations o n some Rhodesiangranites. Spec. Publ. Geol. SOC.S. Afr., 3: 69-78. 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. Worst, B.G., 1956. The geology of the country between Belingwe and West Nicholson. South. Rhod. Geol. Surv. Bull., 43: 218 pp.

Part C Cahen, L. and Snelling, N.J., 1966. The Geochronology of Equatorial Africa. NorthHolland Publishing Company, Amsterdam, 195 pp. Clifford, T.N., 1970. T h e structural framework of Africa. In: T.N. Clifford and I. Gass (Editors), African Magmatism and Tectonics. Oliver and Boyd, Edinburgh, pp. 1-26. Clifford, T.N., Rooke, J.M. and Allsopp, H.L., 1969. Petrochemistry and age of the Franzfontein granitic rocks of northern South West Africa. Geochim. Cosmochim. Acta, 33: 973. De Carvalho, H., 1969. Contribution 5 la Gdochronologie d u Sud-Oeste d e L’Angola. Bol. Serv. Gdol. Minas, Angola, 19: 23-33. Hall, T.C.F., 1949. T h e mineral resources of Angola. S.R. Direccso Provincial dos Services de Geologia e Minas. Angola, 1-238: 5 4 pp. Mouta, F., 1954. Noticia explicatura d o esboco geolbgico d e Angola (1 : 2 000 000). Junta Invest. Ultramas, Lisbon. Simpson, E.S.W., 1970. T h e anorthosite of southern Angola: a review of present data. I n : T.N. Clifford and I. Gass (Editors), African Magmatism and Tectonics. Oliver and Boyd, Edinburgh, pp. 89-96. Stanton, W.I., Schermerhorn, L.J.G. and Korpershoek, H.R., 1963. The West Congo System. Bol. Serv. Geol. Minas, 8 : 69-78.

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501

CHAPTER 9

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

*

Distribution, thickness and areal extent The Pongola Group outcrops in southern Swaziland, in the southeastern corner of the Transvaal and in northern Natal, Republic of South Africa (Fig. 9.1). Outcrops are scattered over an area measuring 250 km northsouth, by approximately 130 km east-west (Fig. 9.2). A minimum depositional area of 32,500 km2 is inferred, since much of the basin has been removed by erosion, destroyed by granite intrusion or lies covered beneath Phanerozoic sediments of the Karoo Sequence. The Pongola Group thins appreciably towards the south. In the northern basin it has a thickness of close to 11,000 m (Matthews and Scharrer, 1968). In the Wit-Mfolozi inlier, near the southernmost Pongola occurrences, only 2500 m of sediments and volcanics are preserved. Over this same span, the lower (mainly volcanic) unit, known as the Insuzi Subgroup, thins from 6100 m t o 1900 m. (Fig. 9.3).

The age of the Pongola Group has been fairly closely established by radiometric dating of its basement, of volcanic rocks within it, and of igneous rocks intrusive into it. Pongola strata rest non-conformably on granitic rocks dated at 3060 Ma by the Rb/Sr whole-rock method (Hunter, 1974). They are intruded by a differentiated, sheet-like assemblage of basic and grano-

* By A. Button.

SOUTH

f

BOTSWANA

WEST \

f..’ Durban

Pon Elmbelh

-

Fig. 9.1. Locality map for the Pongola and Witwatersrand basins.

503 phyric rocks known as the Usushwana Complex, which has given a Rb/Sr isochron of 2874 t 30 Ma (Davies et al., 1969). Zircons from Insuzi Subgroup lavas, have given a U-Pb age of 3090 90 Ma (Burger and Coertze, 1973), while the same authors mention a 3150 150 Ma Rb/Sr whole rock determination on felsites from Swaziland. The Pongola Group is some 3000 Ma old, and is one of the oldest relatively undeformed platform sequences known.

*

Structure Considering its great age, the Pongola Group is relatively mildly deformed. In the principal outcrop area, immediately south of Swaziland, it is deformed in a number of gentle domes and basins, with dips generally in the range of 0 t o 30". This pattern of interference folding is thought to be due t o the pressure exerted by large intrusive granitic bodies which lie t o the east (Fig. 9.2). The intrusive contacts range from conformable t o sharply transgressive, and are marked by a distinct metamorphic aureole. In the Amsterdam region, the Pongola Group is preserved in a synclinal structure with an axis tending south-southeast and plunging in the same direction. It is in this region that the Usushwana Complex attains its maximum development. It is usually intruded along the contact of the PslZgola and the underlying granitic basement. In Swaziland, some of the Pongola rocks occur as large xenoliths within younger intrusive granitic rocks. The xenoliths are steep-dipping, structurally disturbed and metamorphosed. In the southern inliers, the overall structural pattern cannot be determined. In the Wit-Mfolozi occurrences, for example, Pongola strata strike northwest, and dip at angles of between 10 and 25 degrees t o the northeast. In the southernmost outcrops, adjacent t o the Namaqua Natal mobile belt, the Pongola strata are relatively strongly folded, metamorphosed and are involved in a later period of thrust faulting. While the Insuzi Subgroup is developed along the entire western side of the basin, it is rarely present along the eastern side. It seems probable that the depositional axis of the basin shifted in an easterly direction after Insuzi volcanism, so that the overlying Mozaan Subgroup unconformably oversteps the volcanics towards the east.

Stra t igrap hy The Pongola Group is divided into two major divisions, separated by an unconformity. The lower assemblage, which rests non-conformably on older granitoid basement, is known as the Insuzi Subgroup, and is composed mainly of basic volcanics with intercalated sediments. The overlying Mozaan Subgroup comprises alternating arenaceous and argillaceous sediments, with

~HLOEANE VRYHEID 0

P

I

J-MFOLOZI INLIER

0

10

20

30

40

50

60

70

80

90

Fig. 9.2. Geological map of the Pongola basin.

100 KM

505 some iron formation intervals. The unconformity at the base of the Mozaan truncates successively lower units within the Insuzi when traced t o the southeast. ( 1 ) The pre-Insuzi basement In general, the Insuzi Subgroup rests non-conformably on an Archaean basement, composed of granitoid rocks and some greenstone remnants. In the southern part of the basin, the basal Insuzi quartzite rests with a sharp contact on a gently undulating surface cut across fresh biotite-microcline granite. By contrast, in the principal outcrop area further north, a palaeoregolith, up t o 7.5 m thick, is developed above the granitic basement (Matthews and Scharrer, 1968). The palaeoregolith is composed of sericite, quartz and vermiculite, with some microcrystalline quartz. It grades downwards to the underlying granite. Near its top, the regolith develops a crude stratification, and the proportion of sericite diminishes, suggesting incipient re-working. The regolith is thought to have formed by in situ chemical weathering under sub-aerial conditions.

( 2 ) The Insuzi Subgroup The Insuzi Subgroup decreases in thickness towards the south, from 6100 m along the Pongola River, to 1800 m in the south, in the Wit-Mfolozi inlier. The decrease in thickness is probably due t o the combination of a lower rate of subsidence during deposition, and post-depositional uplift and erosion in the south. Little detailed work has been published on the northern basin. The Insuzi commences with 850 m of arenaceous sediment. The basal 120 m comprise a rapidly varying assemblage of argillaceous and arkosic arenites and grits, and some conglomerates and breccias. They are succeeded by 730 m of quartz arenite with some granule-conglomerates (Matthews and Scharrer, 1968). The basaltic Insuzi lavas, which incorporate two or three major intervals of quartz arenite and some shale (Fig. 9.3), attain a total thickness of close to 5200 m (Humphrey and Krige, 1932; Visser et al., 1947). The lavas are usually amygdaloidal. Smaller volumes of acid extrusives are present in places. In the Wit-Mfolozi inlier, the Insuzi attains a maximum thickness of close to 1900 m (Matthews, 1967). Five major stratigraphic subdivisions have been recognized (Table 9.1). The basal quartzite is a mature quartz arenite, and was deposited on a planed surface of basement granitic rocks. The arenite wedges out locally against small palaeotopographic basement highs. A thin conglomerate and an arkosic interval are developed at the base in places. Matthews regarded the basal quartzite as a “near-littoral” shelf sand. The lower and upper volcanic zones comprise amygdaloidal basalts, pillowed in places. Amygdales are of chlorite, quartz, calcite, epidote and

506 VRYHEID-PIET RETlEF

AREA

2

+w: + i

1

-+

Ushushwana Complex

+++

t+

Iron-rich

L O

+++ ++++ ++++

t++

++

Sediments(where thin. Indicated b y + )

Argillaceous Sediments

L+++++

Arenaceous Sediments

Modified after Beukes (1973) and Matthews (1967)

-

Unconformity

E l

Volcanic Rocks

Granite Basement

Fig. 9.3. Stratigraphic columns for t h e Pongola Group.

pyrite. According to Matthews, basalts are of two types, an aphantic, darkgreen to black variety, and a greyish-green type. In both, the primary minerals have been deuterically altered to an assemblage of turbid plagioclase, chlorite, epidote, calcite, leucoxene and opaque minerals. Unpublished

507 TABLE 9.1 Stratigraphy of the Pongola Group in the Wit-Mfolozi inlier (after Matthews, 1 9 6 7 ) Approximate thickness (m) 3. Shale-quartzite zone > 305 2. Quartzite-ironstone zone 183 1 . Basal quartzite-conglomerate zone 0-2 1 3 - - - - - - - - - Unconformity - - - - - - - - - - - - - - - - - - - - - - 5. Upper volcanic zone 700 4 . Quartzite-dolomite zone 762 3. Banded shale zone 61-243 Insuzi Subgroup 2. Lower volcanic zone 0-122 1. Basal quartzite 0- 61 - - - - - - - - - Non.con formity - - - - - - - - - - - - - - - - - - - - - Crystalline basement

Mozaan Subgroup

chemical data (P.E. Matthews, pers. commun., 1977) ind6ate that both tholeiitic basalts and high-magnesia basalts are developed in the Insuzi Subgroup. The banded shale zone comprises graded-bedded and small-scale rippledrift bedded siltstone units in a shale background. Lenticular, fining-up channel-fill deposits with basal conglomeratic lags were described by Matthews, who considered that the zone as a whole could represent a tidal flat deposit. The quartzite-dolomite zone comprises an assemblage of quartz arenites, with lenticular developments of carbonate-cemented quartz arenites, carbonates, shales, tuffaceous sandstones and breccias. The quartz arenites are cross-bedded and contain lenses of granule conglomerates. Shale clasts are developed in some of the arenites. Matthews (1967) described a variety of carbonates, including aphanitic dolomites and pisolitic and oolitic dolomites. The carbonates are frequently partly replaced by chert. Matthews suggested an algal origin for some of the pisolitic and oolitic structures. Von Brunn and Mason (1977) and Mason and Von Brunn (1977) have re-examined parts of the quartzite-dolomite zone. They have described lowrelief, linked domical stromatolites, usually partially silicified. Fenestral textures, filled by coarse sparry dolomite, and intraclastic dolomites (some of them structured by herringbone cross-bedding) confirm that deposition was in the intertidal zone. Fining-up cycles have been documented for Insuzi sediments in the WitMfolozi (Von Brunn and Mason, 1977) and the Mpongosa inliers (Von Brunn, 1974; Von Brunn and Hobday, 1976; Hobday and Von Brunn, 1976). The cycles, interpreted as having formed by progradation of muddy tidal flats across shallow-shelf sands, are very similar t o cycles in the Mozaan Sub-

508 group, and will be described in a subsequent section. The southernmost exposures of the Insuzi Subgroup lie adjacent to the Tugela Fault, which marks the southern boundary of the Kaapvaal Block (Fig. 9.2). Here, the Insuzi succession is deformed and metamorphosed. It comprises a succession of basic schists, kyanite-bearing phyllites, quartzsericite schists, sheared amygdaloidal lavas and clean quartz arenites (Du Toit, 1931; Matthews, 1959). Sheets of epidiorite (originally gabbroic sills) are present, and are folded along with the Insuzi rocks. Locally, along the northwestern margin of these outcrops, a granite, intrusive into the Insuzi, is developed. The Insuzi rocks were involved in a later period of intense thrust faulting, directed t o the north. The thrust faulting postdates a Late Precambrian-Early Cambrian succession known as the Ntingwe Group (Matthews, 1959, 1972). Prior to Mozaan deposition, the Insuzi Subgroup was tilted and eroded. Locally, faulting of pre-Mozaan age has been recorded (Matthews, 1967). ( 3 ) The Mozaan Subgroup The Mozaan Subgroup comprises an alternation of shales and quartz arenites, with some iron formation and conglomerate (Fig. 9.3). In the northern outcrops, it attains a thickness of over 4600 m (Matthews and Scharrer, 1968), while in the southern outcrops, only 700 m are preserved (Matthews, 1967). A sequence of 14 quartziteshale cycles is known from the northern basin, where the quartz arenite/shale ratio is somewhat less than 1. In Swaziland, the overall ratio is 0.8 (Hunter, 1963), but near the base, quartzites exceed shales by a factor of over 2, and a number of pyritic conglomerates are developed. Many of the shales and phyllites are ferruginous, particularly the l s t , 6th, 7th and l o t h , in which iron-formation is developed. In the 6th shale, for example, the iron-formation comprises magnetite with tremoliteactinolite, quartz, chlorite and some spessartite (Hunter, 1963). The 7th shale contains two iron-formations. Iron-formations frequently show contorted bedding. Some of the Mozaan iron-formations contain ooliths comprising chert nucleii, surrounded by siderite (Beukes, 1973). The lower shale units are aluminous, containing andalusite and pyrophyllite. A crystal tuff has been described in the 12th shale by Hunter (1963). In Swaziland, an amygdaloidal basalt is the uppermost stratigraphic unit in the Mozaan. Similar rocks occur near the top of the succession southeast of Piet Retief, and are presently thought to be amygdale-bearing diabase sills (P.E. Matthews, pers. commun., 1977). Both thermal and dynamic metamorphism affect Mozaan sediments in Swaziland. A low-grade regional thermal event is suggested by the recrystallization of quartzite, the development of andalusite and the amphibolegarnet-hlorite assemblage in the iron formations. Adjacent to intrusive granite contacts, sillimanite-bearing quartzites are developed. Hypersthene, staurolite, magnetite and garnet-bearing granulites are developed locally

509 (Hunter, 1963). In addition, up t o three deformational events have been impressed onto Mozaan sediments in parts of Swaziland. Detailed palaeoenvironmental work has not been carried out on the northern parts of the basin. The alternation of thick shales and some iron-formation, with quartz arenites suggests deposition on a shallow marine shelf. According t o Hunter (1963), shales become thinner and more sandy when followed to the north, suggesting that the basin-margin lay in that direction. In the Wit-Mfolozi inlier, the preserved Mozaan sediments are only 700 m thick. Here, the unconformity with the Insuzi Subgroup is indicated by discordant strikes and dips, as well as by sedimentary overlap. Within the inlier, some 1200 m of Insuzi strata are unconformably truncated towards the southeast. The unconformity is overlain by up to 213 m of quartzite with some conglomerate lenses. The basal conglomerate is auriferous and uraniferous in the vicinity of Denny Dalton (Fig. 9.2). It consists of clasts of vein quartz, quartzite and some chert, lava and shale fragments in a pyritic matrix, and is up to about 1m thick. The conglomerate-bearing assemblage is limited to the southeasternmost outcrops. The overlying quartzites, which fine upwards through shale t o iron formation, become thicker towards the southeast, from 30 t o 180 m. The iron-formationshale assemblage is overlain by a quartzite with a 10 cm basal, chamositic or glauconitic conglomerate. This quartzite is followed by over 300 m of shale, including two or three 1 5 m quartzitic sandstone beds (Table 9.1). Von Brunn and Hobday (1976) have documented a pattern of cyclic sedimentation in the basal part of the Mozaan Subgroup. An arenite lithofacies grades up to an alternating arenite-argillite lithofacies, which is covered by an argillaceous lithofacies. The first assemblage has been interpreted as a shallow-marine deposit, formed along an east-west-trending shoreline. The arenites show trough cross-bedding (directed to the southeast) and troughlike scours (with long axes oriented east-northeast). Pebble long axes and ripple crests have a preferred east-west orientation. The uppermost few metres are structured by herringbone cross-statification, double-crested and flat-topped ripples, suggesting deposition on a lower tidal sandflat. The overlying lithofacies commences with flaser-bedded arenite, giving way to wavy- and lenticular-bedded arenite-argillite alternations. Desiccation cracks are commonly developed, as are ripple marks and mudclasts. Deposition was probably on the middle part of a tidal flat. The overlying argdlite lithofacies comprises mudstone with graded siltstone- -mudstone laminae, with some layers of mudclast microbreccia, and shows evidence of desiccation. It is probably a high tidal-flat deposit. Associated jaspilitic ironformation is inferred to have been deposited in small, periodically flooded depressions. The lithofacies are arranged in fining upward cycles, ranging from 1 2 to

510 25 m in thickness. The thickness of such cycles, can, subject t o certain qualifications, be taken as a measure of tidal range (Klein, 1971). If these figures are accurate estimates of tidal amplitude in the Pongola basin, they probably indicate exceptionally large shelf-widths, which is in keeping with the large size of the Gondwanan supercontinent. Tidal sedimentation appears to dominate the Pongola Group; similar fining-up cycles have been described from the Insuzi Subgroup (Von Brunn, 1974; Hobday and Von Brunn, 1976; Von Brunn and Mason, 1977). Post-depositional history Shortly after deposition of the Pongola Group, the succession was intruded by a large volume of gabbroic magma, in the form of differentiated sheets and sills. The thickest of these is the Usushwana Complex, emplaced mainly along the basal unconformity of the Pongola succession. As many as 10 diabasic sills were intruded into the Mozaan Subgroup, most of them along quartzite-shale contacts (Humphrey and Krige, 1932). The Pongola was then intruded, along its eastern margin by a suite of granitic rocks. Granitic intrusion resulted in metamorphism and some of the deformation impressed on the strata. Granites were also intruded into Pongola rocks along the sothern margin of the basin, near the Tugela fault (Fig. 9.2). Along the southern margin of the Kaapvaal Block, the Pongola sediments were involved in northward-directed thrust faulting, probably during Early Palaeozoic times (Matthews, 1972). A final event t o influence the Pongola was the intrusion of sills and dykes of Karoo dolerite.

Mineral deposits At present, the Pongola Group is not a significant source of economically exploitable mineral deposits (Hunter, 1962). In Swaziland, small tonnages of diaspore and pyrophyllite have been mined from a zone of aluminous phyllites in the Insuzi Subgroup. Andalusite and sillimanite are present in large quantities in the Mozaan sediments, but have not been worked. Magnetic banded iron-formations of the Mozaan Subgroup represent a potential source of iron, presently unexploited. Small, sub-economic secondary manganese deposits are found associated with some Mozaan sediments. In the Wit-Mfolozi inlier, the basal Mozaan conglomerate has been mined for gold on a small scale in the past. The Denny Dalton Mine is being actively prospected at present for both gold and uranium. Du Toit (1931) has recorded a number of gold prospects in pyritic conglomerates of the Insuzi Subgroup in the southernmost outcrops, adjacent t o the Tugela fault.

511 THE WITWATERSRAND SUPERGROUP *

Distribution, thickness and areal extent The known development of the Witwatersrand Supergroup is confined to the southern portion of the Transvaal and the northern sector of the Orange Free State (Fig. 9.1). The easternmost development of Witwatersrand rocks lies 120 km to the west of the westernmost outcrops of the Pongola Supergroup. Sediments and volcanics of the sequence have been proved to be present over an ovoidal area, trending southwest--northeast and measuring 350 km in this direction and 200 km in a northwesterly direction. The basin is open at its northeastern and southwestern extremities, so that the original areal extent of the depository remains to be determined, under an extensive

Witwatersrand strata

,

Landskroon Dome

Fig. 9:4. The outcrop pattern of Witwatersrand strata and basement granite domes, showing the restriction of outcrops to the flanks of the domes, along the sides of these structural culminations which plunge towards the depositional/tectonic axis of the basin. The position of the outcrop and sub-outcrop of the base of the Witwatersrand Superp o u p has been determined by surface mapping, by magnetometric and gravimetric geophysical surveys, and by core-drilling (from Pretorius, 1976; reproduced by permission Elsevier Scientific Publishing Co.).

* By D.A. Pretorius.

512 cover of Precambrian and Phanerozoic strata. The comparatively limited degree to which the Witwatersrand rocks outcrop can be seen in Fig. 9.4. It is only as a result of the substantial amount of mining of, and exploration for, gold and uranium mineralization that the sub-cover extent of the basin is as well known as has been delineated in Fig. 9.4. The original depositional area of the Witwatersrand assemblage must have covered at least 80,000--100,000 km2. The progressive increase in the size of the older Precambrian basins on the Kaapvaal craton can be seen in the fact that the Pongola depository was possibly of the order of 30,00040,000 km2, while the two basins succeeding the Witwatersrand are conkm2, in the case of the sidered t o have been about 200,000-300,000 Ventersdorp Supergroup, and about 400,000--500,000 km2, in the case of the Transvaal sequence of rocks. There is a suggestion of a possible tioubling in size between successive basins. The composite thickness of the strata in the Witwatersrand Supergroup has been estimated at 11,000 m. There is no one area within the basin, in which the whole succession has been encountered. This thickness has been

Fig. 9.5. The relationship between goldfields, fluvial fans, and major fold axes in the Witwatersrand Basin. The arcuate shape of the basin is the result of two superimposed directions of folding and has been formed about the northwesterly-trending Koppies anticline. The arrows indicate the direction of transportation of sediment from the source-area situated to the northwest of the basin (from Pretorius, 1976; reproduced by permission Elsevier Scientific Publishing Co.).

513 determined for the accumulation of strata over the depositional axis of the basin, which arcs round the Vredefort dome (Fig. 9.4), between this structure and the Landskroon, Westerdam, Johannesburg, and Devon domes. In the six goldfields shown in Fig. 9.5, the stratigraphic thicknesses of the Witwatersrand Supergroup are as follows: Welkom - 3200 m; Klerksdorp 7500 m; Carletonville - 2800 m; West Rand - 7400 m; East Rand - 3400 m; and Evander - 1500 m. The maximum recorded thickness on the northwestern edge of the basin is 7500 m, while that on the southeastern edge is 2500 m. The above information has been taken from Haughton (1964), Pretorius (1975), Pretorius (1976), and Whiteside et al. (1976). These same four publications form the basis for all other information contained in the descriptions of the Witwatersrand Supergroup, which follow. They are essentially review articles and thus contain summaries of the very considerable volume of literature which has been published on the Witwatersrand Supergroup and its contained gold and uranium mineralization.

The only rocks with which the Witwatersrand strata have been observed to be in contact belong to the Archaean granite-eenstone terrane that rims the basin on all sides. N o developments of the Pongola Supergroup have been recorded in the vicinity of the Witwatersrand depository. A mean age, from both U-Pb and Rb-Sr measurements, of the granites on which Witwatersrand rocks rest is 3000 Ma. The same methods of dating have placed an age of 2800 Ma on the lavas in the basal members of the Witwatersrand sequence. N o other formations in the supergroup have been dated. An age of 2300 Ma has been obtained on lavas in the Ventersdorp Supergroup, stratigraphically well up in the overall succession. Detrital uraninite and monazite in the clastic sediments of the Witwatersrand assemblage have yielded mean ages of 3100 Ma. There is evidence of metamorphic overprints at 2300 Ma (Ventersdorp age) and 2000 Ma (Bushveld Complex age). From the wide spread between the upper and lower age-limites mentioned above, it is estimated that the Witwatersrand Supergroup was laid down in the period 2500-2800 Ma. Its place in the hierarchy of older Precambrian assemblages on the Kaapvaal craton fits between the Pongola Supergroup (2800-3100 Ma) and the Ventersdorp (2300-2600 Ma) and the Transvaal (2100-2400 Ma) sequences.

S true tu re The Witwatersrand basin takes the general form of an asymmetrical synclinorium, the axis of which trends southwest-northeast, with the depositional axis being closer to the northwestern edge of the depository than to

514 the southeastern rim. Dips of beds on both limbs decrease stratigraphically upwards, from very-steep-to-vertical at the base t o less than 20" for the uppermost members. The two long-sides of the basin have been fixed with a fair degree of reliability, but the northeastern and southwestern limits are structural, and not depositional, so that the original extent of the basin remains unknown in these two directions. The present geometry of the depository is dominated by a pattern of interference folding, the two principal components of which trend between northeast and east-northeast and between northwest and north-northwest. The result of the interaction of these two trends has been the production of a series of structural depressions and culminations, represented, respectively, by basins containing the maximum preserved thickness of Witwatersrand strata and of overlying Ventersdorp, Tjansvaal, and Karroo cover and by domes of Archaean granites and greensteons, either exposed or concealed beneath a thinner cover of Precambrian and Phanerozoic rocks. The more important of the domes are shown in Fig. 9.4. The effects of upwarping and downwarping have been magnified by the generation of substantial faulting parallel to the traces of the fold axes. The overall result has been the distribution of Witwatersrand rocks according to an array of alternating synclines-grabens and anticlines-horsts. The outcrops of such strata are restricted t o the flanks of the basement domes. Without repeated upward movement of these structures, it is unlikely that the Witwatersrand Supergroup would have been exposed, but would have remained concealed beneath the younger cover. The area of the Kaapvaal craton which contains Witwatersrand rocks is one which has some of the best-developed and most-frequently-occurring mantled basement domes. The main fold structures in the basin and surrounding country are shown in Fig. 9.5. The northwest trend remains relatively straight, but the northeast fold axes appear to have been bent about the former, presenting a welldeveloped concavity t o the southeast, with a regional plunge to the northwest. Five northeasterly-trending anticlinal axes have been indicated, and these have progressively tighter hinge-zones from northwest to southeast. This increase in the degree of constriction of the cores of the anticlines culminates in the structural highlight of the Vredefort dome, with overturned strata on its northwestern rim, high-grade metamorphism, and shockwave minor structures. The folds within the interference pattern, which give rise to the basins and domes demarcated in Fig. 9.4, have wave-lengths of 30-40 km for both the northwest and the northeast sets. The most important of the many dislocations which affect the Witwatersrand strata are the strike-faults which trend parallel t o the depositional axis of the basin and t o the northwestern and southeastern long-sides. These originated as normal faults, with downthrows towards the depositional axis. They are accompanied by parallel, antithetic, normal faults which dip towards the planes of the major longitudinal faults, thus resulting in elon-

51 5 gated fault-blocks of varying size, which are wedge-shaped in vertical section. Many of the strike-faults show three phases of activation and re-activation. They first acted as normal faults, contributing to a deepening of the basin under an increasing load of sediments; then they assumed the role of highangled reverse faults; and, finally, they took the form of wrench faults. The second and third movements were responses t o space problems caused by the infilling of the depository and by the continued rising of the granite domes. The negative crustal feature in which the Witwatersrand rocks were laid down originally took the form of a yoked basin, or half-graben, with a long axis striking between east-northeast and northeast. The fault-bounded, and thus more unstable, side of the basin lay to the northwest of the depositional axis. The southeastern side was subjected t o more gentle downwarping. The continuous re-activation of the bounding faults on the northwestern side led to repeated elevation and re-elevation of the source-area, further t o the northwest, and the consequent prevalence of generally higher-energy depositional conditions than on the southeastern edge of the basin.

Stratigraphy ( 1 ) Pre-Witwatersrand basement The unconformable contact between the Archaean basement and the Witwatersrand formations represents an hiatus of 300-500 Ma. The contact shows no sharp elevation differences over short distances, leading to the conclusion that the Archaean terrane had been peneplaned t o a surface reflecting only very broad and relatively gentle undulations. That greater uplift and erosion occurred on the northwestern side of the Witwatersrand depository prior to, and during, its development can be seen in the nature of the preserved basement on each of the two long sides of the basin. On the northwestern flank, only the lowermost members of the Onverwacht Group of the Swaziland Supergroup have been preserved, whereas there are members of the Onverwacht, Fig Tree, and Moodies present on the southeastern side. Most of the Archaean greenstone stratigraphy has been eroded away on the more active side of the Witwatersrand depository, supporting the contention that tectonic activation was much more marked on the fault-bounded side of the basin than on the downwarped edge. ( 2 ) General stratigraphy The three groups and seven sub-groups into which the Witwatersrand stratigraphy has been divided are depicted in Table 9.11. Differences of opinion exist as t o what formations should be included in the Witwatersrand Supergroup. Some investigators subscribe t o the possibility that the Dominion Group might actually represent, in its own right, a pre-Witwatersrand, post-Pongola Supergroup, while others maintain that the Klipriviersberg Group of the overlying Ventersdorp Supergroup more cor-

516 TABLE 9.11 Composite stratigraphical thickness of the Witwatersrand Supergroup, compiled from maximum thicknesses of groups and sub-groups as present in type-areas of development Central Rand Group 3200 m

’

Turffontein Sub-Group 1700 m Johannesburg Sub-Group 1500 m

,Jeppestown

Witwatersrand Supergroup 11,000 m

West Rand Group 5100 m

Sub-Group 1400 m Government Sub-Group 2000 m Hospital Hill Sub-Group ‘1700 m

’

Dominion Group 2700 m

Syferfontein Sub-Group 2100 m Rhenosterhoek Sub-Group 600 m

rectly belongs in the uppermost part of the Witwatersrand succession. If the Klipriviersberg is a terminal volcanic phase in the evolution of the Witwatersrand depository, then a well-developed symmetry would be present in the stratigraphic-lithologic response to processes of basin-fill: (5) terminal, very-high-energy, volcanic phase (Klipriviersberg Group); (4) upper, high-energy, sedimentary phase (Turffontein and Johannesburg Sub-Groups); (3) central, pivotal, low-energy, sedimentary phase (Jeppestown Sub-Group); (2) lower, high-energy, sedimentary phase (Government and Hospital Hill Sub-Groups); (1)initial, very-high-energy, volcanic phase (Dominion Group).

( 3 ) Dominion Group The Rhenosterhoek Sub-Group consists of 600 m of andesites, tuffs, tuffaceous breccias, and quartz-feldspar porphyries. At the base are 40 m of conglomerates, sericitic quartzites, and shaly quartzites. The Syferfontein Sub-Group is devoid of sediments and is composed of rhyolites, which predominate, and lesser amounts of andesite, tuff, and volcanic ash. Of the total thickness of 2700 m, volcanics constitute 2630 m, coarse clastics 60 m, and fine clastics 1 0 m. The volcanics: sediments ratio is 37.9 : 1 and the coarse-clastics : fine-clastics ratio 6.0 : 1. The size of the Dominion depository is estimated to have been of the order of 15,000 km’. ( 4 ) West Rand Group The gross thickness of 5100 m of the middle of the three Witwatersrand

517 groups is made up of 2300 m of coarse clastics, 2300 m of fine clastics, and 500 m of volcanics, the last-mentioned of which are restricted t o one episode in the Jeppestown Sub-Group. The volcanics : sediments ratio is 0.1 : 1and the coarse-clastics : fine-clatics ratio 1.0 : 1. Except for certain exposures in the Klerksdorp area, the sediments of the West Rand Group, which have been preserved, represent a more-distal lowerenergy facies than characterize the Central Rand Group. Conglomerates are only sparsely developed in the former group, whereas they make up a substantial percentage of the total range of lithologies in the latter group. Of the coarse clastics, those in the Hospital Hill Sub-Group are supermature orthoquartzites, while those belonging t o the Government and Jeppestown subgroups are sub-greywackes, composed of quartz and chlorite, with small quantities of muscovite and, in some horizons only, fine-grained, disseminated magnetite. The fine clastics are made up of varying quantities of quartz, kaolinite, sericite, chlorite, and chloritoid, with the ratios of phyllosilicates varying from horizon to horizon. A characteristic feature of the shales is that, in general, the content of sodium is higher than that of potassium. Banded iron-formations and ferruginous shales are present in the Hospital Hill Sub-Group, indicating the prevalence, at times, of chemical sedimentation, a condition which has not been observed to have existed, to any conspicuous extent, in the overlying components of the Witwatersrand Supergroup. The evidence of tidal re-working in the orthoquartzites, the presence of structures typical of tidal inlets and intertidal sandflats, and the development of chemical sedimentation point t o possible distal-shelf facies being present in the lowermost part of the West Rand Group and to the existence, at that time, of a more open basin. Thereafter, the basin appears to have closed, and fluvial and lacustrine sedimentation became dominant. It is conjectured that the size of the West Rand depository was about 100,000 km2, with a width of some 250 km. The northwestern edge of the basin lay approximately 100 km t o the northwest of the depositional axis, while the southeastern rim was about 150 km to the southeast of the axis, illustrating the asymmetry of the basin. A regression of the basin-edge took place on the northwestern side during the formation of the West Rand Group, with the depositional limits of the Hospital Hill, Government, and Jeppestown sub-groups advancing progressively into the basin. To the south, southeast, and east, stratigraphically higher strata onlapped, and generally transgressive conditions prevailed towards the southwestern, southeastern, and northeastern edges of the basin.

( 5 ) Central Rand Group This group has a maximum thickness of 3200 m, and, in the Johannesburg and Turffontein sub-groups combined, there are 2700 m of coarse clastics, 200 m of fine clastics, and 300 m of volcanics. Only one period of volcanic activity has been observed, towards the top of the Johannesburg Sub-Group.

518 The ratio volcanics : sediments is 0.1 : 1, the same as for the West Rand Group, and the ratio coarse clastics : fine clastics is 12.6 : 1, substantially higher than for the underlying group, indicating a much more proximal environment, in general, for the deposition of the Central Rand Group. In addition to the frequent presence of conglomerates, the coarse clastics comprise hydrothermally altered, feldspathic quartzites, in which the quartz content varies between 70 and 9076,the balance being made up of muscovite, pyrophyllite, chlorite, chloritoid, and chert. The quartzites of the Johannesburg Sub-Group, in which are found the greater proportion of economically exploitable, gold-bearing horizons, are 10-15% higher in silica than are the other quartzites in the group. It is also only in the Johannesburg Group that recrystallized quartz is present. The composition of the fine clastics is similar to that of the members of the West Rand Group, except that magnesia is generally higher. The size of the Central Rand depository is thought t o have been of the order of 80,000 km2, pointing to the possibility that there was a progressive shrinking of the Witwatersrand basin, with time. The width of the basin was 200 km, some 50 km narrower than that which is thought to have prevailed during West Rand times. The northwestern side of the Central Rand basin was about 70 km t o the northwest of the depositional axis and the southeastern side about 130 km t o the southeast, evidence that, despite the shrinking of the depository from West Rand through Central Rand times, the asymmetry was maintained. Between the beginning of West Rand sedimentation and the commencement of Central Rand deposition, the northwestern edge of the basin advanced some 60 km towards the depositional axis. The general pattern of progressive regression which developed from Hospital Hill, through Government, and into Jeppestown times, continued into Johannesburg and Turffontein times. The resulting off-lap of equivalent depositional facies contributed t o the general coarsening-upwards of lithologies, which placed the highest-energy clastics in the upper portion of the Turffontein Sub-Group.

Post-depositional history Sedimentation in the Witwatersrand basin was succeeded by the Ventersdrop period of major volcanic activity, possibly the most extensive in the post-Archaean, pre-Phanerozoic history of the Kaapvaal craton. Crustal adjustments placed the depositional axis of the Ventersdorp basin t o the northwest of that of the Witwatersrand depository, while the axis of the Transvaal basin was developed still further northwestwards. A very considerable movement of fault-blocks was concurrent with the outpouring of the Ventersdorp lavas, and the most intense structural deformation of the Witwatersrand strata was completed before the beginning of Transvaal basin-

519 development. The rising of the granite domes continued through Ventersdorp into Transvaal times, so that the depositional surface of the Transvaal depository represented a combination of elevated domes of basement granite and uplifted fault-blocks of Witwatersrand and Ventersdorp rocks. There was far less tectonic deformation of the Witwatersrand strata in post-Ventersdorp times than during the period of volcanic activity. The granite domes continued t o rise until, at least, Karroo times (Upper Paleozoic- -Lower Mesozoic), bringing the Witwatersrand rocks through the Precambrian and Phanerozoic cover. Two main periods of metamorphism affected the rocks of the Witwatersrand Supergroup - one associated with the major period of volcanic activity during Ventersdorp times and the other with the emplacement of the Bushveld Complex. At least five periods of dyke-and-sill intrusion can be seen in the Witwatersrand strata. The oldest of these is a product of a Ventersdorp, pre-Transvaal episode of intrusion, and then Bushveld Complex-age and Pilanes-berg-age dykes and sills were emplaced. A long time-interval intervened before the Karroo dolerites were intruded. The youngest dykes are kimberlitic in composition and might be of Cretaceous age. By volume, the dykes and sills occupy 5-10% of the space taken up by members of the Witwatersrand Supergroup.

Mineral deposits The Witwatersrand basin contains the greatest goldfields yet discovered in the World. The sediments are also an important source of uranium. Goldfields developed within a fluvial fan which formed at the interface of a river, flowing southeastwards from a source-area on the northwestern side of the basin, and a shallow-water lake. Six such fluvial fans have been delineated on the northwestern side of the depositional axis (Fig. 9.5). No fans of comparable dimensions and economic importance have been located on the southeastern side of the depository, where more gentle downwarping and lower-energy sedimentation took place. Concentrations of detrital gold and uranium took place on unconformities which developed between pulses of sedimentation. In many instances, the beginning of a new cycle of sediment-accumulation was marked by the laying down of gravel, so that conglomerates are often present immediately above the unconformities. Most of the gold was deposited in such conglomerates or in algal mats which grew during breaks in sedimentation associated with the development of the unconformities. Only a very limited amount of gold has been won from the strata in the West Rand Group - a more-distal, lower-energy phase of deposition. By far the greatest concentration of gold and uranium is present in a number of horizons in the Johannesburg SubGroup. Important deposits of gold and uranium are also located in the Turffontein Sub-Group, more so in the lower members.

520 There would appear to be a broad correlation between depositional energies and optimum conditions for mineralization. The generally moredistal facies of the West Rand Group did not host favourable conditions for accumulations of detrital heavy minerals, while the more-proximal environment of the Turffontein Sub-Group supported higher-energy conditions that were not particularly conducive to the settling and re-working of small grains of gold and uraninite. Intermediate facies with intermediate depositional energies in Johannesburg times favoured processes and responses that were best suited to concentrating heavy minerals in braided river channels in a fluvial fan. THE VENTERSDORP SUPERGROUP *

Distribution, thickness and areal extent The Ventersdorp Supergroup outcrops mainly in the central and western Transvaal, in the northwestern Cape Province, and in part of the Orange Free State and Botswana. A very substantial part of the Ventersdorp is covered by younger units (the Transvaal and Karoo Supergroups) in the Orange Free State, the northwestern Cape and the central Transvaal (Fig. 9.6). The overall distribution is in a northeast-trending belt, 750 km long by 350 km wide, comprising an area of some 260,000 km’. The Ventersdorp attains a thickness in excess of 3678 m in the Bothaville area of the Orange Free State. It is characteristically between 2000 and 3000 m thick. Thickest developments are usually within the confines of the Witwatersrand structural basin.

Considering its widespread occurrence, the Ventersdorp is very inadequately dated. A date of 2300 Ma was given for a single U/Pb determination on zircon in a quartz porphyry near Klerksdorp (Van Niekerk and Burger, 1964). The Zoetlief acid lavas, now considered as part of the Ventersdorp succession, were dated in the range 2500-2700 Ma (Van Niekerk and Burger, 1968). Grobler and Emslie (1976) quote an isotopic date of 2310 Ma for a quartz porphyry in the T’Kuip Hills, in the northern Cape, in a succession probably equivalent t o the middle division of the Ventersdorp. A determination of 2620 f 50 Ma has been made (U/Pb on zircons) on 15 samples from the Makwassie Quartz Porphyry Formation (C.B. van Niekerk, pers. commun., 1977). In Botswana, Rb-Sr whole-rock determinations give ages of around 2200 Ma, with one isochron at 2695 It 125 Ma

* By A. Button.

Fig. 9.6. Locality map for the Ventersdorp basin.

522 (Key, 1977). The Pb-Pb and U-Pb methods (three determinations) give ages of around 2630-2640 Ma. The radiometric dating of the Ventersdorp is confusing. A 2200--2700 Ma spread is unrealistically large. It seems probable that the true age is around 2600 Ma, and that younger dates represent subsequent thermal events, such as burial by the Transvaal Supergroup or Bushveld intrusion. Struc tu re Over much of its outcrop area, the Ventersdorp is gently deformed, and is sometimes horizontally disposed. Within the Potchefstroom synclinorium, it has the same structural pattern as the Witwatersrand. Around parts of the Vredefort Dome, the succession is vertical or even overturned. Gravity faulting, which has produced numerous horsts and grabens in the Witwatersrand Basin, has affected the overlying Ventersdorp strata. Most of these faults were active during Ventersdorp times, and had important effects on stratigraphy and sedimentation. Stratigraphy The Ventersdorp and associated units (Zoetlief, Pniel, Kuip and Klipriviersberg successions) have a history of complicated terminology and correlation. Largely as a result of the stratigraphic analysis of the Ventersdorp and associated units (Winter, 1965, 1976), much of the confusion is settled. The history of nomenclature and correlation will not be repeated, being described in depth in Winter’s papers. Surface exposures of the Ventersdorp lavas tend t o be poor, the dips often low and the lithology variable along strike. As a consequence, the type sections of most of the Ventersdorp formations were selected from borehole intersections in the Bothaville area of the Orange Free State. Here, a very large number of boreholes has been sunk in search of auriferous and uraniferous Witwatersrand conglomerates, and has provided material for a study of vertical and lateral relationships in the Ventersdorp. The Ventersdorp Supergroup rests on a large variety of formations, including elements of the Archaean basement, the Dominion Reef succession, and parts of the upper and lower divisions of the Witwatersrand Supergroup * The Ventersdorp has been subdivided into three major units, from the base to the top, the Klipriviersberg, the Platberg and the Pniel Groups (Winter, 1976). ( 1 ) The Klipriviersberg Group The Klipriviersberg Group has a thickness, in its type borehole, of 1533 m, but is up t o 2130 m thick in places (Winter, 1976; Whiteside, 1970). It has

523

Fig. 9.7. Schematic representation of Ventersdorp stratigraphy and stratigraphic relations.

been divided into six formations (Fig. 9.7 and Table 9.111). The group consists largely of basic lavas with layers of agglomerate and ashflow tuff. It usually rests on the Upper Witwatersrand, where the relationship varies from nearly conformable t o moderately unconformable. The Klipriviersberg is largely restricted t o the limits of the Upper Witwatersrand Basin and wedges out rapidly beyond these limits (Fig. 9.7). The lowest units within the group tend t o fill topographic depressions on the pre-Ventersdorp surface, so that younger units lap on t o the basement on basement-highs. There are a number of markers in the Klipriviersberg, some of which have a wide distribution. The Alberton Porphyry, for example, is found in all outcrops northeast of Klerksdorp, as far as the Evander Goldfield (Whiteside, 1970). Purple-topped ash-flow tuffs are characteristic of the Orkney Formation in the Klerksdorp area and t o the south, while the Jeanette Agglomerate is a useful marker in the Orange Free State. Over much of its extent, the Klipriviersberg Group rests on a conglomerate known as the Ventersdorp Contact Reef. The conglomerate is usually only a few metres thick, and can be shown to truncate unconformably various stratigraphic units of the Upper Witwatersrand. Zones of payable gold mineralization within the conglomerate can usually be related to suboutcropping Witwatersrand conglomerates. Whiteside (1970) recorded, in the Klerksdorp area, a 305 m-thick assemblage below the Klipriviersberg. This unit, referred t o as the Vaal Bend For-

524 TABLE 9.111 Stratigraphic subdivision of the Ventersdorp Supergroup (after Whiteside, 1 9 7 0 ; Winter, 1976; and Wyatt, 1976). Group

Formation

Member

Thickness (m)

Pniel

Allanridge Bothaville

745 390

Edenville

568

Lorraine

2 17

Lithology

Amygdaloidal lava Conglomerate, arenite, shale - - - - - - - - - - Unconformity or disconformity - - - - - - - - - - - - - Rietgat Garfield 930 Amygdaloidal lava and sediments Platberg Makwassie Quartz 3 64 Quartz porphyry, minor Porphyry sediment Kameeldoorns 555 Coarse- and fine-grained sediment - _ - - - - - - - - Unconformity - - - - - - - - - - - - - - - - - - - - - -

Klipriviersberg

Jeanette Agglomerate Orkney Alberton Porphyry Westonaria

Vaal Bend

97 480 171 Meredale

150

305

Lava with green chalcedonic and milky quartz amygdales Variolitic and spherulitic lava Agglomerate Purple-topped ash-flow tuffs Porphyritic lavas Talcose tuffs, high-Mg basalt in Meredale Member Basic and acid lava with sediments

mation, and comprised of acid and basic lavas with intercalated sediments, can probably be considered as an additional, locally developed formation in the Klipriviersberg Group. Winter (pers. commun., 1977) is of the opinion that the Vaal Bend Formation is a composite of pre- and post-Klipriviersberg rocks in a region of stratigraphic attenuation. The lavas of the Ventersdorp Supergroup have long been referred to as andesites. Wyatt (1976), in his study of the Klipriviersberg in its type-area, showed that this was not the case, and that the volcanics are usually tholeiitic basalts, albeit with calc-alkaline affinities. Wyatt recognized a chemically distinct unit in the Westonaria Formation, which he called the Meredale Member. It is comprised of light- and dark-coloured volcanic layers on a centimetre- to metre-scale. The dark layers frequently contain spheres of the lighter material, which are thought t o represent an insoluble liquid

525

phase. The dark fraction has a characteristic high-Mg chemistry, unlike the balance of the Klipriviersberg. The Klipriviersberg Group appears t o have been extruded on land, and probably represents an analogue of Phanerozoic flood basalts. Evidence for sub-aerial extrusion includes some pahoehoe flows, repeated layers of flowtop breccia (Wyatt, 1976), and the paucity of pillow structures, which have only been described from a few locations (Whiteside, 1970). The flows probably emanated from fissures, which, in the Johannesburg area, lay to the north of the outcrop belt, and were oriented east-west (Wyatt, 1976). The Klipriviersberg Group is usually truncated by an unconformity, overlain by sediments of the Kameeldoorns Formation. The Klipriviersberg was block-faulted after extrusion had largely ceased. Erosion of horst blocks of Klipriviersberg material floored by older formations supplied much of the coarse clastic detritus in the overlying unit. (2) The Platberg Group The Platberg Group includes the Kameeldoorns (base), Makwassie Quartz Porphyry and Rietgat (top) formations (see Fig. 9.7 and Table 9.111). The Kameeldoorns Formation (555 m thick in its type borehole) comprises a highly variable assemblage of coarse- and fine-grained sedimentary material. Some objections have been raised t o the recognition of the Kameeldoorns as a formation, since, where the grabens were filled by volcanic flows, the formation is not developed (Whiteside, 1970). The Kameeldooms is made up of coarse clastic sediment (often boulder conglomerate and talus breccia) near t o the scarp faces of the horst blocks. The composition of the clastics reflects the geology of the horst block, which in addition t o Klipriviersberg material, may also include Witwatersrand and basement formations. Away from the horst blocks, the conglomerates interfinger with immature arenaceous, argillaceous and carbonate sediments. The Kameeldoorns probably encompasses the basal clastic material in the Ventersdorp equivalents previously referred to as the Zoetlief Formation. The middle, and most distinctive, unit of the Platberg Group is the Makwassie Quartz Porphyry Formation, 364 m thick in its type borehole, but increasing t o 2000 m in places. It is developed throughout the Orange Free State and the southwestern Transvaal. The Makwassie comprises quartzporphyry flows with quartz-free porphyritic and non-porphyritic lavas, ashflows and some sediments. Quartz porphyry flows up to 600 m thick have been recorded near volcanic centres (Winter, 1976). Away from such centres, the proportion of quartz porphyry in the formation diminishes. Flow structures, agglomeratic units and a gradation of quartz-porphyry to bedded tuffs has been taken as evidence for the extrusive origin of these very thick flows. The Makwassie usually follows on Kameeldoorns sediments. Where these are absent, it rests on any one of a number of older units. Moving away from the axis of the upper Witwatersrand Basin, it covers, in sequence, the Klip-

526 riviersberg, the upper and lower divisions of the Witwatersrand Supergroup, finally lapping onto the Archaean granitic basement. In a number of areas in the Southwestern Transvaal and the northwestern Cape, a sequence of quartz porphyries and associated sediments has been mapped, resting on the basement and covered by basic Ventersdorp lavas. These were referred to as the Zoetlief Formation. On the basis of Winter’s stratigraphic work, the quartz porphyries of the Zoetlief may be assigned t o the Makwassie. The Rietgat Formation (930 m thick in the type borehole) follows conformably on the Makwassie, and could be contemporaneous with parts of that formation. It is a mixed volcanic and sedimentary unit characterized by an upward decrease in the proportion of volcanic material. The lavas are intermediate in composition and include both porphyritic and amygdaloidal types, the latter sometimes containing jasper amygdales. Flow tops show signs of weathering and erosion. Beds of chert and of limestone are characteristically developed in the Rietgat sediments. Winter (1963) described stromatolites in Rietgat limestones, and was of the opinion that they were formed in small ephemeral lakes. Walter (1972) identified the Ventersdorp stromatolites as Gruneria f. nou. Grobler and Emslie (1976) described stromatolitic limestones and cherts in the T’Kuip hills in the northern Cape. The stromatolites are interbedded with arkose and intermediate volcanics in an assemblage underlain by acid lava and arkose, resting on the Archaean granitic basement. The stromatolitic assemblage is probably a Rietgat equivalent. Winter (1976) distinguished the Garfield Member (360 m thick) as a local unit near the base of the Rietgat. It may be equivalent in age t o parts of the Makwassie. It comprises mainly porphyritic lava of intermediate composition, with minor intercalated sediment and some non-porphyritic lava.

( 3 ) The Pniel Group The Pniel Group comprises the two uppermost units in the Ventersdorp Supergroup. It rests with an erosional contact on underlying units of the Ventersdorp, and is overlain, apparently conformably, by the Transvaal Supergroup. The basal unit of the Pniel is the Bothaville Formation, 390 m thick in its type borehole. It consists of clastic sediments, characteristically in a cycle which commences with conglomerate, grading up through impure arenites to shales, and coarsening upwards to conglomerate once again. The arenites are somewhat more mature than those of the Rietgat Formation. The conglomerates are made up of clasts of jasper, banded chert and quartz porphyry, the latter being derived from the Makwassie. The Allanridge Formation is the uppermost unit of the Ventersdorp Supergroup. It is 745 m thick in its type borehole and is composed of finegrained, intermediate composition, amygdaloidal and slightly porphyritic

527 lava. It usually follows conformably on the Bothaville Formation, but is known to lap onto the basement near the margins of the Ventersdorp Basin. The Allanridge Formation is overlain by the Black Reef Quartzite (the basal unit of the Transvaal Supergroup) with little evidence of any significant period of erosion. Winter (1965) has expressed the view that the Pniel succession could represent the initial phase of the Transvaal Supergroup in the southwestern Transvaal, Orange Free State and northwestern Cape. It is thus possible that the Pniel is a time equivalent of other pre-Black Reef units in the Transvaal Supergroup, including the Wolkberg Group (Button, 1973a) and the Buffalo Springs Group (Tyler, 1976).

Post-depositional history Subsequent t o Ventersdorp volcanism, parts of the basin were subjected to intensive erosion. For example, in the Johannesburg area, the upper units of the Ventersdorp, presuming they were formed here, were erosively removed before Black Reef deposition. Here, a sericitic weathering profile is developed on the basic volcanics (Wyatt, 1976). Ventersdorp volcanics suffered a common structural history to the Transvaal Supergroup, and were strongly deformed around the Vredefort Dome, and were also basined around the Bushveld Complex.

Mineral deposits The only important mineral deposits known to be associated with the Ventersdorp are those of the Ventersdorp Contact Reef, below the Klipriviersberg Group (De Kock, 1964; Knowles, 1966). Conglomerates in this unit contribute importantly t o the gold production of the West Wits Line and Klerksdorp regions. THE TRANSVAAL SUPERGROUP

*

Distribution, thickness and areal extent The Transvaal Supergroup is restricted to two structural basins, one in the northwestern Cape and southern Botswana, the second in the Transvaal, extending into eastern Botswana (Fig. 9.8). The basins have a preserved area of about 250,000 km’, but must have originally covered an area twice that size (Button, 1973d). The thickness of the Transvaal succession is up t o 15 km, if the 3 km-thick Rooiberg Felsite, at the top of the pile in the central Transvaal is included. The succession in the northwestern Cape is usually between 5 and 6 km thick.

* By A. Button.

Fig. 9.8. Locality map for the Transvaal basin.

529

The age of the Transvaal Supergroup is fixed in relation to geochronological determinations within the succession, and within overlying formations. A lava in the Pretoria Group (upper unit of the Transvaal Superg;oup) has been dated at 2224 f 21 Ma by the Rb/Sr whole-rock method (D. Crampton, pers. commun., 1972). Associated shales have yielded a date of 2263 f 85 Ma by the same method (Hamilton, 1976). The mafic phase of the Bushveld Complex intrudes Transvaal strata, and has been dated at 2095 f 24 Ma (Hamilton, 1976). In the northwestern Cape, lavas in the Olifantshoek Group, overlying the Transvaal sequence, have been dated at 2070 f 90 Ma (Rb/Sr whole-rock, Crampton, 1974). Transvaal sedimentation probably commenced at, or slightly before 2300 Ma ago, and ceased around 2100 Ma ago.

Structure The Transvaal Supergroup outcrops in a number of structural domains. In the Transvaal, the Bushveld d o m a i n is characterized by gentle-to-moderate dips inwards towards the complex (Fig. 9.9). Within the complex, Transvaal rocks are exposed in the Marble Hall and Crocodile River floor-highs, where they are metamorphosed and plastically deformed. Along the Mhlapitsi fold belt, the sediments are strongly folded by rejuvenation of the east-northeasttrending Murchison tectonic zone (Fig. 9.9). The strata of the Transvaal Basin are deformed in a major rim syncline (the Potchefstroom synclinorium) around the Vredefort Dome. The basins in the Transvaal and in the Cape and Botswana are separated by a major, north-northwest-trending arch, where pre-Transvaal rocks (mainly Ventersdorp lavas) outcrop. In the Cape-Botswana sub-basin, the strata dip at 1-2" west in the Ghaapplato domain. The degree of deformation increases gradually towards the Doomberg lineament, adjacent to which the Transvaal rocks are strongly folded along north to north-northwesttrending axes. The Transvaal Basin is little affected by faults. In the eastern Transvaal, gravity-faulting, on a north-northeast trend, has resulted in a series of horsts and grabens in the Sabie-Pilgrim's Rest area, and in the Penge region. Around Thabazimbi, reverse strike faulting has resulted in the repetition of stratigraphic sequence. In the Hartebeespoort Dam region, north of Johannesburg, a northwest-trending graben has dropped down Transvaal and Bushveld rocks. Near Zeerust, a series of north-trending gravity faults cuts Transvaal strata.

Fig. 9.9. Map showing geology of the Transvaal basin.

531 Stratigraphy and conditions of deposition The Transvaal strata are conveniently subdivided into three major units, separated by unconformities. These are the basal volcanic and clastic unit, the chemical sedimentary unit, and the upper clastic unit (Button, 1976a). ( 1 ) Basal volcanic and clastic unit Representatives of this unit outcrop along the northern and northeastern margins of the basin, where they are known, respectively, as the Buffalo Springs Group (Tyler, 1976) and the Wolkberg Group. Stratigraphic units within these groups are strongly influenced by palaeotopography, wedging out against basement-highs (usually granitic tracts), and attaining maximum thicknesses in ancient valleys, frequently situated over greenstone belts (Button, 1973a). In both regions, the basal groups thin markedly away from the basin axis, and are eventually overstepped by the Black Reef Quartzite, the basal arenite of the overlying chemical sedimentary unit. The Wolkberg Group, which is up t o 2000 m thick, comprises six formations (Fig. 9.10). A basaltic lava, pillowed in places and with stromatolitic carbonates between flows (Button, 1973b), is sandwiched between two coarse-grained, sometimes conglomeratic, feldspathic subgreywacke formations. The latter probably represent the products of braided fluvial sedimentation. The remaining formations comprise alternating feldspathic quartzite and siltstoneshale units, frequently arranged in prograding, coarsening-up cycles. Some cycles are demonstrably due to deltaic progradation into water-depths of 30-50 m. The Buffalo Springs Group (up to 1500 m thick) commences with a substantial thickness of immature arenite and some conglomerate, again the products of braided fluvial sedimentation (Tyler, 1976). The clastics are overlain by basaltic, and then by rhyolitic volcanics. Minor developments of sediment, including some stromatolitic chert, are found between some of the rhyolitic flows. ( 2 ) Chemical sedimentary unit This unit is developed throughout the Transvaal Basin. It varies in thickness up t o a maximum of about 3500 m. In the eastern basin, it is known as the Chuniespoort Group, while in the northern Cape it comprises the Campbell and Griquatown groups (Fig. 9.10). In both basins, the stratigraphy is similar, commencing with a basal quartzite, passing up through a transition zone t o carbonates and cherts, and terminating with banded iron formation. In the Transvaal, the basal (Black Reef) quartzite follows conformably and gradationally on the basal volcanic and clastic unit in the axial parts of the basin. Towards the basin margins, it oversteps this unit, and rests unconformably on older formations, including Archaean granites and greenstone

532 , I 8

o r

.n

-

ROOIBERGFELSITE

INTRUSION PLANE OF BUSHVELD COMPLEX WLLSTRoOMLAVA HOUTENBEK FORMATION STEENKAMPSBERG WARlLITE

NEOERHORST FORMATION L A K H V A L E I OUARTZITE VERMONT FORMATION MAGALIESBERG QUARTZITE

LYDENBURG SHALE MEMBER MACHADOOORP VOLCANIC MEMBER PRUISSEN DOLOMITE MEMBER BOVENSHALE MEMBER OASWORT OUARTZITE OWAALHEUVEL FORMATION HEKWORT BASALT BOSHDEK FORMATION NOOITGEDACHT QUARTZITE MEMBER OUITSCHLANO FORMATION PENGE FORMATION MALMANI DOLOMITE BLACK REEF QUARTZITE

I

SILVERTON SHALE

Y

w

I 0

1T M__

~

~

~

~

~

l

~

~

"nCnn'olm~ty

Banded ,ran iOrmsliOn

VOELWATER JASPER

Dolomite

/ MABINFORMATION

-

SELATI FORMATION SCHELEM FORMATION

ONGELUKLAVA

Folded Snad

MAKGANYENE OlAMlCTlTE S"dIP

Dlarnlrt8lL

GAMAGARA SHALE KOEGAS FORMATION ASBESHEUWELS FORMATION

nikosic A r e w e

GHAAPPLATO DOLOMITE

- ABELERASMUSVOLCANICS SEKORORO FORMATION

SCHMIDTSDRIF FORMATION

A

Acid Volcanic

0

E l i 8 i VakanolRlloved

Yl

c i m x Basemm,

Greensfone BE!!

Fig. 9.10. Stratigraphic columns for the Transvaal basin. A. Bushveld sub-basin. B. Cape-Botswana sub-basin.

belts, the Witwatersrand and Ventersdorp successions. The quartzite is up to 500 m thick along the Selati trough (Fig. 9.9) in the northeastern Transvaal, where it incorporates 100 m of basalt (Button, 1973a). It is usually less than 30 m thick in the rest of the basin. The Black Reef commences with a basal

533 conglomerate and is overlain by quartzite, usually compositionally and texturally mature. It is thought to have been deposited in the subtidal domain of a very broad, gently shelving sea which lapped across a large proportion of the Kaapvaal craton. The Black Reef passes up into carbonate sediments through a transition zone of dolomite interlayered with quartzite and carbonaceous shale. The shales are sometimes mud-cracked, which suggests intertidal conditions of deposition. The carbonate pile (Malmani Dolomite in the Transvaal, Campbell Group in the northern Cape) can be subdivided into a number of units, based on parameters such as the proportion of chert, shale and limestone and the colour, texture and chemistry of the dolomite. In the Transvaal, five units are consistently developed (Button, 1973c; Eriksson et al., 1976), the transition zone (previous paragraph), the lower dolomite and chert unit (lightcoloured, chert-rich dolomite), the chert-poor unit (dark, chert-poor dolomite with shale and quartzite in places), the upper dolomite and chert unit and the mixed zone (chert-poor dolomite, some limestone and thin shales and iron-formation lenses). In the northwestern Cape, there is, on average, less chert and more limestone in the section. There, a basal quartzite and shalebearing transition zone is recognized (the Schmidtsdrif Formation, Visser and Grobler, 1972), overlain by a thin but very persistent iron-formation, passing up t o recrystallized dolomite and chert, overlain by a dolomite and limestone assemblage (Beukes, 1977). Beukes recognized a total of 11 forrnations within the carbonate succession. When traced to the southwest, towards the Doornberg Lineament, the platform carbonates interfinger with a thinner assemblage comprising ferruginous carbonates, iron-formation and carbonaceous shale. Beukes has interpreted this facies change as due to a transition from a shallow carbonate platform t o a basinal environment. Mechanical and organo-sedimentary structures are developed throughout the carbonates and cherts, and have formed the basis of local stratigraphic subdivisions (Eriksson and Truswell, 1974). Stromatolites vary from algal lamination, through domed and columnar forms, to very large elongate mounds (up t o 100 m long, 2 m high and 10 m wide, Eriksson and Truswell, 1973). Although it is abundantly stromatolitic, preservation of algal remains is poor, and is presently limited to one algal microfossil (Nagy, 1974) and to thread-like traces on scanning-electron-microscope images of etched carbonate surfaces (McGregor et al., 1974). Mechanical structures include ripple marks, ripple cross-lamination, oolites (in inverse-graded units) and intraclastic breccias. Signs of evaporitic sedimentation are not abundant, but are, nevertheless, present. These include length-slow silica (Truswell and Eriksson, 1975), tepee structures (Eriksson, 1977) and gypsum casts (Bertrand-Safarti, 1976). The carbonates were deposited in a spectrum of sedimentary environments on a very wide, flat, gently shelving marine platform, backed by a lowlying continental hinterland. Lagoonal, as well as subtidal, intertidal and

534 supratidal facies have been recognized (Eriksson et al., 1976). Carbonisotopic patterns in the carbonates are essentially identical t o those of modern marine carbonates, supporting these conclusions (Eichmann and Schidlowski, 1975). In chemical terms, the carbonates are either true dolomites or limestones. In the dolomites, divalent iron and manganese were introduced along with, and proxy for, Mg in the crystal lattice (Button, 1975; Eriksson et al., 1975). There is cyclical increase in iron content upwards through the section, which culminates in the overlying banded iron-formation (Button, 197613). The iron-formation in the chemical sedimentary unit is known as the Penge Formation in the Transvaal, and the Griquatown Group, composed of the Asbesheuwels and Koegas formations, in the northwestern Cape. The ironformation attains a maximum thickness of about 600 m in the Transvaal and 2000 m in the Cape. It comprises macro-, meso- and micro-banded sediment with very persistent layering. Quartz, magnetite, hematite, stilpnomelane, riebeckite, minnesotaite, grunerite and carbonates (siderite, ankerite, ferrodolomite, dolomite and calcite), either in mixed or monomineralic layers, comprise the iron-formation (Beukes, 1973). Grunerite is restricted to the Transvaal Province, where the metamorphic grade is higher due to the intrusive Bushveld Complex. Riebeckite and grunerite are frequently present in their fibrous forms, and are mined as crocidolite and amosite asbestos, respectively. Some of the stilpnomelane-rich layers contain altered volcanic shards (La Berge, 1966). Spherical structures (5--40 microns in diameter) have been noted, and suggest a biologic contribution t o iron formation deposition (La Berge, 1973). Sediments commonly interlayered with iron formation are carbonaceous shales and iron formation intraclastic breccias. The latter are valuable stratigraphic markers in the Cape Province (Engelbrecht, 1962; Malherbe, 1970). In the Koegas Formation, a major facies change has been demonstrated, from jaspers in the north to minnesotaite-bearing slates in the south (N.J. Beukes, pers. commun., 1976). Continuity of sedimentation units, combined with lack of clastic sediments and general absence of mechanical sedimentary structures, suggest deposition in restricted, saline basins on a tectonically very stable platform. The cyclical repetition of thin iron formation lenses and intraclastic breccia and oolite at the top of the Malmani Dolomite has been used to suggest that at least some of the iron formation was deposited in basins barred from the open shelf sea by wave-built banks of carbonate detritus (Button, 197613). In the northeastern Transvaal, an assemblage of dolomite and limestone with some quartzite and shale follows on the iron formation, and is known as the Duitschland Formation (Button, 1 9 7 3 ~ ) Martini . (1977) has shown that this assemblage rests unconformably on the Penge Formation, and has suggested that it be incorporated in the Pretoria Group. However, the Duitschland is itself unconformably covered by the Pretoria Group (Button, 1973d), and was subjected to the basin-wide period of chemical weathering which preceded Pretoria Group deposition.

535 ( 3 ) Upper clastic unit Following on the chemical sedimentation described above, the sea retreated off the Kaapvaal block, exposing the strata t o the effects of both mechanical and chemical weathering. Maximum uplift and denudation (accompanied by gentle folding in places) occurred along the southeastern margin of the basin, where, in places, the entire Transvaal succession was removed. A karsted surface was very extensively developed over the outcropping carbonate formations. Following on the erosive period, a transgressive sea moved back across the craton, reworking the chemically resistant chert and iron-formation clasts left on the erosion surface, to produce a basal coarse-clastic unit to the upper clastic assemblage. The upper clastic unit is best developed in the Transvaal, where it is up t o 7 km thick (10 km if the Rooiberg Felsite is included), and is known as the Pretoria Group. The equivalent in the northern Cape, the Postmasburg Group, is thinner (up t o 2000 m), comprises a larger proportion of volcanics and chemical sediments, and is considered to be the distal shelf equivalent of the Pretoria Group. The Pretoria Group comprises a marginal-marine cyclical assemblage of quartzite and shale, with three carbonate and three volcanic units, and, in the north, a number of wedges of fluvial arkose (Fig. 9-10). A variety of depositional environments are represented, including the marine shelf (muds and sands), the intertidal (lower intertidal sands, to mixed middle tidal flat, to upper, muddy tidal flat). Carbonates, formed during periods of, or in local areas of, reduced clastic input, are stromatolitic, and were deposited both inter- and sub-tidally (Button and Vos, 1977). Glacio-marine and fluvioglacial sediments are developed in the Timeball Hill Formation and equivalents (Visser, 1971). The volcanics, which are basaltic or intermediate in composition, are usually associated with some pyroclasts (agglomerate and tuff). They were extruded both on land (Hekpoort and Dullstroom units) and in a submarine setting (Machadodorp Volcanic Member, Button, 1974). The thick (3 km) Rooiberg Felsites were extruded during the final stages of Transvaal sedimentation, but were subsequently physically separated from the rest of the Pretoria Group by the intrusion of the Bushveld Complex. Walraven (1977) has shown that the Rooiberg felsites grade up t o a red-bed sedimentary succession, known as the Loskop Formation. This formation is the last pre-Bushveld unit of the Transvaal Basin, and it, the Rooiberg Felsite and the Bushveld intrusives are unconfonnably overlain by the 1800 Ma-old Waterberg basin. In the Northern Cape, the pre-Postmasburg erosive surface had considerable relief. In particular, large basins were developed in breached anticlines in the iron-formation. The lowest stratigraphic unit of the Postmasburg Group (the Gamagara Shale) is restricted t o such basins, and laps out against the encircling palaeo-scarp. Much of the coarse-grained material in the Gamagara was shed directly off the scarps into the breached anticlinal basins.

536 The Gamagara commences with a chert breccia (where it rests on older dolomite), followed by a thin iron-formation, manganiferous, ferruginous, aluminous and carbonaceous shales with wedges of iron-formation conglomerate. This assemblage grades up to the orthoquartzitic Gamagara quartzite, which is capped by a shale. A stratigraphic hiatus was followed by deposition of the Makganyene Diamictite (an assemblage of tillitic sediments, shales and some limestone), followed by extensive volcanism on a submarine shelf to produce the Ongeluk Lava. The lavas are prominently pillowed in places, the interpillow voids being jasper-filled. Associated hyaloclastite breccias (Grobler and Botha, 1976) and massive lavas are developed. The volcanics are covered by the Voelwater Jasper, an assemblage of jasper, ironformation, some limestone and sediments very rich in manganese. The Voelwater assemblage is the youngest preserved unit in the Transvaal Supergroup in the northern Cape, and is unconformably overlain by coarse clastics of the Olifantshoek Group, which incorporates a 2070 Ma-old volcanic unit (Crampton, 1974).

Post-depositional history The major event affecting the Transvaal Basin was the emplacement of the Bushveld Complex. The mafic phase of the complex was intruded in a number of arcuate lobes, which largely determine the present outcrop shape of the basin. The Rooiberg Felsite, representing a low density cap to the Transvaal succession, was floated off above the denser Bushveld magma. The underlying Pretoria Group is heavily intruded by mafic sills related t o the complex, and has suffered pervasive contact metamorphism.

Mineral deposits An exceptionally wide variety of minerals have been produced from the Transvaal Basin. At present, most mining is in the chemical sedimentary unit, where limestone, dolomite, iron (enriched iron-formation), fluorite (replacements in dolomite) and crocidolite and amosite asbestos are produced. The Voelwater and Gamagara units of the northern Cape contain, respectively, some of the largest manganese and high-grade hematite deposits in the world. Very large tonnages of andalusite are present in Pretoria Group hornfelses, and are being mined.

B. The younger sedimentary-volcanic basins THE WATERBERG AND SOUTPANSBERG GROUPS

*

The younger basins on the Transvaal craton are represented by the Waterberg basins and the Soutpansberg trough. The intracratonic Waterberg basins

* By H. Jansen.

TABLE 9.IV Stratigraphy of the Waterberg Group

Subgroup

Southern, western and central portions of late-Waterberg basin

Northern, eastern and central portions of late-Waterberg basin

Formation

Lithology

Formation

Vaalwater

feldspathic sandstone, arkose, siltstone, shale

Vaalwater

Blouberg block-fault zone Formation

Lithology

Blouberg

sandstone, grit, conglomerate, feldspathic sandstone, arkosic conglomerate and breccia, mudstone, shale trachyte, trachyandesite, andesite

Lithology

see column on the left

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Kransberg

Cleremont

sandstone

Cleremont

see column on the left

Sandriviersberg

sandstone

Mogalakwena

sandstone, grit, conglomerate

Aasvoelk0P

siltstone, mudstone, shale, laharite, sandstone, feldspathic sandstone, conglomerate

Makgabeng

sandstone

sandstone, grit, conglomerate

Setlaole

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

Matlabas

Schildpadk0P

~

Central and western portions of Alma trough

sandstone, grit, conglomerate, feldspathic sandstone, arkose, mudstone, tuff ~~

~~

~

Eastern portion of Alma trough

Subgroup

Formation

Lithology

Formation

Lithology

Sterkrivier

sandstone, grit, conglomerate, shale, mudstone, trachyte

Nylstroom

feldspathic sandstone, grit, greywacke and Alma conglomerate, sandstone, grit, mudstone, shale - - - _ - - -- - - - - - - -sandstone, grit, congloSwaersmerate, siltstone, shale, hoek tuff ace ous grey Wac ke , trachyte, quartz porphyry

Swaershoekberge Formation

Lithology

Glentig

quartz porphyry, conglomerate, sandstone, siltstone, lava

Cullinan-Middelburg basin Formation

Lithology

Wilgerivier

sandstone, grit, conglomerate, feldspathic sandstone, arkose, volcanic rocks

P

- -- - - - - - - -- - - - - -

Tl

cn

cu CD I

cn

+ 0

541

and the pericratonic Matsap trough are characterized by a predominant arenaceous fill and a comparatively low degree of volcanicity. The Soutpansberg trough, though largely intracratonic, is considered t o be a special type of depository, namely an aulacogen (Jansen, 1976a,b) with a fill of predominant arenaceous rocks and large accumulations of lava flows. The Waterberg basins in the Transvaal are the main Waterberg basin and the Cullinan-Middelburg basin. The small Ootse basin occurs on the TranSvaal-Botswana border. Large areas in eastern Botswana are occupied by a comparatively thin platform cover of Waterberg beds and scattered erosion remnants in the Springbok Flats (Transvaal) also indicate an original thin platform cover in this area and a link between the main Waterberg and the Cullinan-Middelburg basins. In the pericratonic Matsap trough in the northern Cape Province a portion of a very thick sequence may be correlated with the Waterberg succession in the Transvaal, but its lower and upper boundaries have not yet been definitely established nor the extension of the basin in Waterberg times. The main Waterberg basin consists of two overlapping basins, namely the early Waterberg basin or Alma trough and the late Waterberg basin. The new stratigraphy of the Waterberg Group is given in Table 9.IV and Fig. 9.11. The revision of the Soutpansberg stratigraphy has not yet been completed, but for general descriptions of the lithological and structural features a provisional subdivision has been submitted (Jansen, 1976b).

Stratigraphy of the Waterberg Group The Waterberg sequence represents one major sedimentary cycle which is predominantly continental and arenaceous. Even where the two basins overlap, the relationships between the formations are largely conformable and only locally unconformable. Local unconformities are commonly developed and occur on all stratigraphic niveaux. Hence, there is no reason to subdivide the entire sequence into two major units as has been done in the past (“Loskop and Waterberg systems”). The bulk of the deposits consists of fluvial and shallow water sediments, the material being supplied by erosion of predominantly granitic and metamorphic belts. Deposition of coarse clastic sediments (conglomerate, arkose, etc.) is attributed to uplifts in source areas including basin rims, which also led t o the formation of local unconformities. The lower subdivisions, which are confined t o the Alma trough, are the Swaershoek, Glentig, Alma, Sterkrivier and Wilgerivier formations. These formations are predominantly arenaceous, but interbedded trachytic lava flows (Swaershoek formation), quartz porphyry flows (Swaershoek and Glentig formations), pyroclasts (Wilgerivier formation), feldspathic sediments (Alma formation) and argillaceous sediments (all formations) ar!e also developed. Contemporaneous block-faulting has controlled deposition of coarse clastic

542 sediments, which were laid down along or in the vicinity of old fault scarps, for instance in the Gatkop area, south of Loubad and the northern Swaershoekberge. The basal portion of the Swaershoek formation is only developed in the Nylstroom syncline which more or less coincides with a protobasin. Another protobasin is occupied by the Glentig formation which is unconformably overlain by the upper portion of the Swaershoek formation. Transgression t o the north took place during deposition of the upper portion. The maximum thickness of the Swaershoek formation in the Nylstroom syncline is 2500 m. Source rocks were predominantly Rooiberg felsite and sedimentary rocks of the Transvaal Supergroup. Large-scale erosion of the Bushveld granite started with deposition of the Alma formation which is feldspathic in the western portion of the Alma trough. To the east the Alma beds become less feldspathic and lithologically more similar t o the Swaershoek formation and finally the two formations merge into the Sterkrivier formation. The maximum thickness of the Alma formation near Alma is approximately 3000 m. The Wilgerivier formation in the Cullinan-Middelburg basin is lithologically rather similar to the Swaershoek formation. Its maximum thickness is 2700 m. The formation of the late Waterberg basin was partly caused by subsidence of a large area north of the Alma trough which was followed by a large transgression and deposition of the coarse arenaceous beds of the Schilpadkop and Setlaole formations. The southward extension of the basin where it overlaps the Alma trough, is indicated by two outliers of the Schilpadkop formation, i.e. west of Loubad and east of Nylstroom (Kranskop). The above-mentioned formations are overlain by the partly argillaceous Aasvoelkop formation in the southwestern portion of the basin and by the arenaceous Makgabeng formation in the northeastern portion. The Makgabeng formation is considered to be of aeolian origin (Meinster and Tickell, 1976). Intensive block faulting on the northeastern rim of the basin caused deposition of the coarse feldspathic members of the Blouberg formation, partly in small isolated basins or troughs. Of the upper Waterberg succession (Kransberg Subgroup) the Sandriviersberg and Mogalakwena formations are predominantly coarse-arenaceous and probable laid down by braided streams draining uplifted portions of the Limpopo belt (Tickell, 1976). Coarse conglomerates form three main zones in the Mogalakwena formation. The uppermost subdivision of the Waterberg Group, i.e. the Vaalwater formation, consists of fine-grained arenaceous, feldspathic and argillaceous sediments, which were most likely deposited in an inland basin. Near Blouberg block-faulting continued and the predominantly nonfeldspathic members of the Blouberg formation were laid down in small troughs. The largest trough, the Lebu trough, measuring 28 X 3 km was filled by trachytic and andesitic lavas. The Blouberg formation is uncon-

543 formably overlain by beds of the Soutpansberg Group, wich are correlated with the middle portion of the group in the Soutpansberg (see below). In contrast with the other formations of the Waterberg sequence, the Blouberg formation does not represent a complete succession nor a lithological unit because its members are largely confined to small depositories formed on negative fault-blocks.

Stratigraphy of the Soutpansberg Group The Soutpansberg sequence may be subdivided into a lower predominantly volcanic portion, a middle arenaceous portion which also becomes partly volcanic t o the east and an upper portion which is arenaceous, argillaceous and partly volcanic. Although several problems concerning the detailed stratigraphy are not yet solved, the available data point to a progressive development of the trough from east to west by almost continuous downwarping and downfaulting (Jansen, 1976b). The lower volcanic portion decreases in thickness from east to west, partly by contemporaneous faulting with the result that in the western Soutpansberg the arenaceous middle portion transgresses to the west and north over the Limpopo belt. However, further to the west, around Blouberg the Lebu trough and smaller troughs were already formed prior to deposition of the middle portion (see above). The Soutpansberg sequence displays a more cyclic pattern of sedimentation and volcanicity than the Waterberg one. Other differences are the large accumulations of lava flows, the absence of coarse immature (arkosic) sediments and the poor development of coarse conglomerates in the Soutpansberg sequence. Regional unconformities are absent and local intraformational unconformities are less frequent than in the Waterberg basins. The sediments in the Soutpansberg trough are predominantly of fluvial origin and partly also laid down in lakes or small inland basins, in particular the argillaceous sediments. The maximum thickness (7 km or even more) of the sequence is considerable in comparison with that of the Alma trough (6 km) and of the late Waterberg basin (3 km) and all evidence indicates that the Soutpansberg trough and the Waterberg basins were isolated depositories.

The structural pattern of the basins The structural pattern of the basins came into existence by downwarping and locally intensive contemporaneous block-faulting along the basin rims and, in the case of the Alma trough, also by a moderate to intensive postWaterberg compressional deformation. The complex structural fault-pattern of the Soutpansberg trough originated during at least three phases, i.e. contemporaneous, post-Soutpansberg and post-Karoo. In the southern portion of the Alma trough several prominent structures

544

of post-Waterberg age such as the Nylstroom syncline, the Loubad and Zwartkloof anticlines and the Swaershoekberge anticlinorium, already developed in their embryonic stage during late Bushveld and early Waterberg times. The embryonic structures are attributed to magmatic updoming and subsidence (Du Plessis, 1972). As mentioned previously the Nylstroom syncline more or less coincides with a protobasin in which the lower portion of the Swaershoek formation was laid down. Similar structures may have developed in the Cullinan-Middelburg basin, The quartz porphyry flows in the Glentig formation and basal Swaershoek beds are considered to represent late-Bushveld magmatic activity when Waterberg sedimentation had already started. After solidification of the acid rest magma the above-mentioned structures were locally subjected t o block-faulting which caused deposition of coarse conglomerates and probably also extrusion of trachytic lavas. Block-faulting continued during deposition of the Alma formation in the Gatkop area. Feldspathic graywacke, boulder conglomerate and wedge arkose were formed by rapid erosion of granite hills along fault-scarps. The early Waterberg phase of deformation was terminated by a slight compressional phase during which gentle folds were superimposed on the already existing pattern. More intensive post-Waterberg deformation led t o reactivation of the preexisting structures and the development of new ones. The belt of deformation extends from Thabazimbi to Potgietersrust over a distance of 160 km but there is no evidence of its extension far beyond the Alma trough nor of a link with the Matsap folded belt in the northern Cape Province. Its linearity is locally disrupted by the older, non-linear, partly magmatic pattern. In the Gatkop area fault-blocks of early Waterberg age were reactivated by thrusting and tear-faulting and finally overridden by the Gatkop overthrust mass. Thrusting also occurred along t;?e southern limb of the Zwartkloof anticline, but the area between the Zwartkloof and Loubad anticlines was hardly affected by post-Waterberg deformation, which in this area was evidently confined t o the margins of the older structures. The peculiar pear-shape of the Nylstroom syncline is attributed to its embryonic magmatic stage. During the post-Waterberg phase the limbs were tilted up to the vertical and locally flexured or folded with the result that the western portion of the syncline resembles a brachysyncline and the eastern portion a huge box-fold. The Swaershoekberge anticlinorium developed from an embryonic anticline or upwarp and on its northern limb faultblocks of early Waterberg age were steeply tilted. The beds of the Swaershoek formation were locally over-turned on the northern limb and thrown into small infolds in the felsite core of the anticlinorium. The compressional post-Waterberg phase hardly affected the succession in the late-Waterberg basin. Only north of the Swaershoekberge the beds of the Schilpadkop formation are tilted up t o the vertical and the beds of the Sandriviersberg formation up to 30". Finally, the entire Waterberg basin and

545 surrounding areas were subjected t o tensional stresses which gave rise t o normal faults, fractures and the intrusion of numerous diabase dykes and sills. This was the only type of deformation which affected the flat lying successions in the late Waterberg basin. However, on its northeastern rim the arkosic members of the Blouberg formation were locally tilted up t o the vertical or even overturned and subsequently unconformably overlain by the Sesalong boulder conglomerate (Mogalakwena formation). Block-faulting continued and the boulder conglomerate beds were locally flexured or tilted up to 45". During or after this phase the Blouberg block-fault zone widened and the small troughs or grabens north of Blouberg were formed. These troughs and the Lebu trough may be considered as proto'troughs of the Soutpansberg trough. The older faults were in part again reactivated during and after deposition of the Soutpansberg beds, also in post-Karoo times. The complex structural pattern of the Soutpansberg trough mainly consists of intricate fault-systems which originated during at least three phases. The older the faults are, the more they are obscured by reactivation during following phases. Contemporaneous faulting, though not conspicuous in the field, has played an important role concomitantly with downwarping and extrusion of lavas (Jansen, 1976b). The contemporaneous faults are most likely normal faults, but the development of a braided fault-pattern during a later phase may also be attributed, at least partly, t o transcurrent faulting. After deposition of the Soutpansberg beds tensional faulting and intrusion of diabase dykes and sheets were followed by regional tilting of the entire succession to the north and probably more faulting. Most of the faults were reactivated in post-Karoo times during which further tilting t o the north took place. The structural pattern of the Soutpansberg trough is, therefore, dominated by faulting and regional tilting. Compressional folds are largely confined t o argillaceous beds in the upper portion of the succession and are irregular in distribution and orientation. Their origin is not attributed t o compression on a regional scale but t o sagging of fault-blocks (Jansen, 197613). From the previous text it is clear that the Soutpansberg trough does not only differ from the Waterberg basins by its lithological and volcanic features but also by its structural pattern. The Waterberg basins developed by subsidence of the underlying crust, the centres of subsidence shifting t o the north, whereas the Soutpansberg trough developed on an almost permanent zone of weakness by rifting and downwarping which spread from east t o west.

Volcanicity in the basins In the Waterberg basins volcanic activity was of a comparatively low degree in contrast with the Soutpansberg trough, Detailed petrological and chemical data are scarce because of the high degree of alteration of the lavas. In the Soutpansberg trough the volume ratio of basaltic, andesitic and trachytic

546 lava is difficult t o estimate, but it seems that basaltic lavas predominate. However, in the Alma and Lebu troughs basalts have not been found so far. The volcanic rocks in the Cullinan--Middelburg basin are confined t o three occurrences and composed of pyroclasts and andesite. With the exception of the Cullinan-Middelburg basin pyroclasts are developed on a minor scale in the Waterberg basins. The trachytic lavas in the Swaershoek formation are locally agglomeratic and interbedded with sediments containing volcanic detritus, namely tuffaceous graywacke and a volcanic conglomerate with cavities which may be a laharite. In the late Waterberg basin cavity bearing sediments (laharites) in the Aasvoelkop formation and a thin layer of tuff and ignimbrite at the base of the Setlaole formation are not associated on outcrop with lavas, but the association of lavas (basalt, soda trachyte), ignimbrite and cavity bearing sediments has been demonstrated at two localities in the Soutpansberg. This relationship is in favour of a lahar origin of the cavity bearing sediments although the latter were most likely laid down under less torrential conditions than typical laharites. The basalts, andesites and trachytes of the Waterberg and Soutpansberg basins seem t o represent one large alkaline province intermediate in time and partly in composition, e.g., alkalinity between the Bushveld Complex and the Pilanesberg alkaline province. The basalts in the Soutpansberg are continental tholeiites, but the differentiation trend is towards slightly alkaline rocks poor in or devoid of quartz. The composition of the trachytes in the Swaershoek formation (Jansen, 1970) is abnormal because of the high FeO and TiOz contents. The distribution of trachytic lavas in the Alma trough points t o extrusion from fissures with a dominant N-S trend. The lava flows rapidly decrease in thickness and frequency east and west of the focal area near Loubad and t o a much lesser extent t o the north and south. In the Cullinan-Middelburg basin volcanic activity was of a more explosive character and the three main occurrences also display a N-S trend. The eruption centres of the occurrences in the late Waterberg basins are not known. In the Soutpansberg the lavas probably extruded from fissures along or parallel with contemporaneous faults, thus during rifting with a dominant ENE trend. The lavas in the basins most likely issued from fractures in the underlying crust which extended into the lower crust or upper mantle. In the structural control ENE and NNE trends played a dominant role. The Lebu and Alma troughs fall in the NNE-striking abyssal fracture zone which extends from the Great Dyke t o Trompsburg. The Lebu trough is situated on the intersection of this zone and the ENE-striking zone of weakness along the boundary of the central and southern marginal zones of the Limpopo belt. The Alma trough lies on the intersection of the abyssal fracture zone and the westward extension of the greenstone belt near Pietersburg which is also a zone of weakness.

547

Pulaeogeogruphy o f the basins In Fig. 9.12 palaeogeographical reconstructions of the basins are based on the distribution and lithology of the individual formations, their mode of deposition, transport directions and age determinations on the underlying basement. The original distribution of certain subdivisions almost coincides with their present one, for instance of the Alma formation, but other formations may have extended far beyond their present limits, for instance the Sandriviersberg formation south of the Waterberg plateau. Sedimentary structures abound in some of the formations. Cross-bedding and ripple marks are most common, trough type cross-bedding more than the planar type. Folded or deformed cross-bedding is developed in the Soutpansberg beds and in some subdivisions of the Waterberg sequence. Desiccation cracks and imprints of rain drops are less common, mainly in argillaceous beds. Unusual structures are displayed in the Vaalwater formation, namely convolute laminations and convolute lumps or dome-shaped structures which are attributed to thixotropic behaviour of silty sediments (De Vries, 1973, pp. 136- 137). It is evident that almost the entire Waterberg and Soutpansberg sequences were laid down under fluvial and shallow water (littoral or inland basin) conditions and only locally under aeolian conditions (Makgabeng formation). In general subsidence of the crust under the basins must have kept pace with the accumulation of detritus. The persistent transport directions from north, northeast and east in all subdivisions of the Waterberg Group indicate that the source areas were mainly the southern and central zones of the Limpopo belt, thus mainly granitic and metamorphic terrains, which were also source areas of the Transvaal basin. However block-faulting gave rise to the formation of small positive areas south of the main Waterberg basin, for instance the granite ridge(s) in the Gatkop area during deposition of the Alma formation. Another important feature is the fact that the northern boundaries of the Alma trough and Transvaal basin almost coincide with the southern boundary of the southern marginal zone of the Limpopo belt. The southern zone was already incorporated into the Transvaal craton in Transvaal times and subsequently intruded by the Bushveld Complex north of the Transvaal basin. The large transgression of middle Waterberg times initiated the formation of the late Waterberg basin and extended north of the Alma trough as far as the central zone of the Limpopo belt where late mobilizations (2100 1800 Ma) took place. Hence, the age of the middle Waterberg succession is estimated at approximately 1700 Ma and the lower age limit of the lower Waterberg succession at 2000 Ma because of the overlap between the Bushveld and Waterberg periods. The upper age limit of the Waterberg group may be approximately 1300 Ma, the average age of the Pilanesberg alkaline province.

548 During their long period of evolution (2000-1300 Ma) the Waterberg basins changed considerably in shape and area (see Fig. 9.12). The centres of subsidence shifted t o the north from Nylstroom (Swaershoek times) via Alma (Alma times) t o Vaalwater (late Waterberg times). The formation of small protobasins i s related to late Bushveld magmatic activity, for instance the protobasin coinciding with the Nylstroom syncline. Subsequently the Nylstroom protobasin extended t o the north during deposition of the upper portion of the Swaershoek formation and during deposition of the Alma formation the basin became more similar t o an elongated deep graben. Compared with the late Waterberg basin the Alma trough is confined t o a relatively small area but the maximum cumulative thickness of the deposits (6 km) exceeds that of the late Waterberg basin (3 km). During the large transgression of middle Waterberg times the late Waterberg basin extended to the west into Botswana and t o the north as far as the southern limit of the central zone of the Limpopo belt. The cent,ral portion of the broad rise north of the Alma trough and Transvaal basin subsided, but the western and eastern portions, respectively the Makoppa and Letaba rises, remained positive areas. Intensive uplifts in the central zone of the Limpopo belt (“Limpopo rise”) caused deposition of coarse immature sediments (Blouberg formation) along the northeastern basin rim and during a later stage large rivers laid down the coarse conglomerates of the Mogalakwena formation. The Limpopo and Letaba rises evolved into a broad crustal arch or upwarp. To the south the Vaalwater beds were laid down in the shrinking basin and the evolution of the Waterberg basins was terminated. On the epeirogenetic uplift in the northeastern Transvaal the Soutpansberg trough was formed by rifting and downwarping which spread from east t o west and t o a lesser extent from south t o north. This process may already have started in early Waterberg times in the eastern Soutpansberg and during its course the succession transgressed on the central zone of the Limpopo belt where late mobilizations were followed by uplifts, rapid erosion and peneplanation, most likely in middle t o late Waterberg times. According t o the aulacogen hypothesis (Jansen, 1976b) the Soutpansberg trough developed from the edge of the craton into the interior. Similar t o most aulacogens, the trough dies out in the interior of the craton concomitantly with a decrease in rifting and volcanic activity. In these areas (Botswana) the successions of the Chwapong Hills, Lotsani River and probably also the Sushong beds were laid down, which are, therefore, correlated with the Soutpansberg group. The Soutpansberg trough was probably linked with a pericratonic miogeosyncline or geosyncline, the miogeosyncline deposits being represented by the Umkondo sequence along the eastern boundary of the Rhodesian craton. Deposition of the Waterberg, Soutpansberg and Umkondo successions was in part contemporaneous: the Waterberg in intracratonic basins, the Soutpansberg in an aulacogen and the Umkondo in a miogeosyncline. The

555

Umkondo beds were partly metamorphosed during the formation of the Moqambique and Zambezi belts. Evidence of low-grade metamorphism in the Soutpansberg beds and locally even in Waterberg beds is mainly attributed to burial. The time interval of deposition of the entire Soutpansberg sequence is not entirely established but according to the aulacogen hypothesis the trough was a long lived aulacogen and thus the duration of the Soutpansberg period is at least of the same order as the Waterberg era or even more. Geodynamics of the basins The centres of subsidence of the Witwatersrand, Transvaal and Waterberg basins have shifted from the south-central Transvaal to the north and west. This centrifugal tendency is brought into relation with crustal undulations in the craton and accretion of the latter (Jansen, 1976a). The incorporation of the central zone of the Limpopo belt into the craton was followed by the formation of the late Waterberg basin and the Soutpansberg trough. The crustal undulations are attributed to lateral subcrustal flow away from the centres of subsidence, but more complex subcrustal events, for instance mantle upwelling and contraction, may have taken place under the graben-type basins, namely the Soutpansberg, Lebu and Alma troughs. The same may also apply t o the domal structures in the southern Transvaal, the origin of which is also attributed in part to lateral subcrustal flow. The migration of the basins followed the northward shifting of the loci of granite emplacement (Hunter, 1974, p. 318). The granitic belts, for instance the southern and central zones of the Limpopo belt were subjected to isostatic uplifts and erosion and the detritus supplied t o the basins. This interpretation is another expression of accretion of the craton and may also be linked with the hypothesis of crustal undulations by lateral subcrustal flow because isostatic uplifts in the granitic belts may have promoted subcrustal flow. That basining is related t o subcrustal events is also demonstrated by the type of volcanic rocks in the basins. Even some of the granites in the granitic belts, viz. the source areas of the basins, may have been generated in the mantle (Hunter, 1974, p. 318). During the formation of crustal undulations the downwarps, viz. the centres of subsidence of the basins, were also controlled by pre-existing zones of weakness in the crust (Jansen, 1976a). These zones are represented by the greenstone belts with the Murchison direction (ENE) and the NNE-striking abyssal fracture zone. The discontinuous pattern of deformation within the basins and basin rims was also in part controlled by these zones of weakness and by drag exerted on the crust by subcrustal flow, the stresses being confined t o the zones of weakness. As mentioned previously volcanic activity was also controlled by the zones of crustal weakness. The ubiquitous diabase dykes in the Waterberg basins are not confined to the latter. In the dyke pattern ENE to NE and NNE trends are also promi-

556 nent, but WNW to NW trends become more dominant in the western portion of the main Waterberg basin. NW t o NNW trends west of the Waterberg basins are displayed by the Pilanesberg dykes and major faults but very poorly developed east of the Waterberg basins. The post-Waterberg dykes indicate widespread tensional stresses in the crust which may be attributed to a general decrease in crustal temperature and in lateral subcrustal flow. This interpretation agrees with the fact that large mobilizations, uplifts and downwarps did not occur during this period. The Waterberg and Soutpansberg basins represent an important stage in the evolution of the Transvaal craton, but very few data are available on the history of the latter during the long post-Waterberg, pre-Karoo period between 1300 Ma and 250 Ma. THE UMKONDO GROUP

*

Since they were first described by Brackenbury in 1906, the Umkondo rocks have been referred to by many different names amongst which are Frontier System (1920), Sabi System (1920), Spungabera Series (1924), Umkondo System (1924), Gairezi Series (1957) and Manica Series (1963). The name Umkondo System, agreed on as a composite term to include all the various elements in 1964, was subsequently changed t o Umkondo Group to comply with recommended Precambrian lithostratigraphic terminology. The Umkondo Group is a thick sedimentary succession unconformably deposited upon granites and gneisses of the Archaean Basement Complex and on the rocks of the Limpopo Mobile Belt. The sediments form the Eastern Highlands and Chimanimani Mountains along the eastern border of Zimbabwe and the Gairezi-Barue Highlands in Western Moqambique. Vail (1964) traced the sediments farther north into the Rukori Range and the Mount Darwin area of Zimbabwe. However,recent mapping by the Zimbabwe Geological Survey has shown that the sediments in the Mount Darwin area belong to the Rushinga Group, of considerably lesser age. Here, persistent horizons of pelite, dolomite and pink meta-arkose are of identical lithology to the Sediments making up the 800 t o 850 Ma Makuti Group in the Kariba area which is known t o be the extension of the Katanga Group of southern Zambia (Broderick, 1976). The unconformity between the Umkondo and Rushinga Groups probably lies just north of Monte Senga-Senga where Moqambique geologists have distinguished a new sequence of sediments (Moqambique 1 : 250,000 geological map Sheet SUL-E-36/J 1968). Although considered by them to be part of the Umkondo Group, the extensive development of feldspathic gneisses with lesser crystalline limestone is much more typical of the Rushinga lithology.

* By V.Stocklrnayer.

557 The Umkondo Group can be sub-divided into two distinct suites or facies: a western succession of virtually flat-lying, shallow-water, dominantly argillaceous and arenaceous sediments and a sequence of deformed and metamorphosed, deeper-water quartzites and pelites lying t o the east. The former, almost wholly restricted t o Zimbabwe, is best developed in the ChipingaLower Sabi area in the south with smaller occurrences at Inyanga, Ruangwe Range and Chirwa Dome in the north. The deformed sediments, which occur mainly in Mocambique, can be traced northwards from the Statonga Ranges and Chimanimani Mountains in the south, through a series of outliers in the Bandula region into the Gairezi-Barue Highlands.

x

x

x

Scale of kilometres 0

20 L

x

20

L

L

L

40 l

80

60 l

I

i

100

I

I

t

120 I

x

3

EXPLANA TION

Aiiurium andyounger formations

x

x

w3

Rushinga Group

Rhodesia Facies

1

3

Older formations

Fauh

Do/erites and /owas

v vv

Fig. 9.13. Distribution of the Umkondo Group.

Umkondo

558 There has been great controversy about the relationships between the two facies of the Umkondo Group. Tyndale-Biscoe (1957) concluded that the deformed Gairezi Series in the Inyanga area was older than the flat-lying sequence t o the west which had been termed the Umkondo System. Their proximity was attributed to thrusting of the Gairezi Series. Slater (1965, 1967), working in the Inyanga-Ruangwe Range area, suggested that the Umkondo System and Gairezi Series were deposited contemporaneously but differed in appearance due to a combination of facies variation and degree of metamorphism. He referred t o the two components as the Inyanga Facies and the Gairezi Facies. In the south, Watson (1964, 1969) considered that the Frontier Series of the Chimanimani Mountains, which had been equated with the Gairezi Series, was the lowest member of the Umkondo System. Vail (1964) disagreed and attempted to show that the deformed Frontier Series was the metamorphosed equivalent of flat-lying Umkondo sediments immediately t o the west. Mapping in the Inyanga area (Stocklmayer, 1975) in part confirms Slater’s findings. In his area of investigation, Stocklmayer retains the terms Inyanga and Gairezi Facies of the Umkondo Group. The former contains five formations of generally shallow-water, predominantly lagoonal sediments and the latter, three formations of deeper-water, deformed sediments t o the east. The boundary between the two facies shifted laterally throughout sedimentation, resulting in an overlapping and interfingering of the Inyanga and Gairezi Facies. He further suggests that only the Basal Formation of the Gairezi Facies

TABLE 9.V The Umkondo Group Zimbabwe Facies South Volcanics Upper Argillaceous Formation Arenaceous Formation Lower Argillaceous Formation Calcareous Formation

Moqambique Facies North (Inyanga Facies) Upper Argillaceous Formation Arenaceous Formation Lower Argillaceous Formation Calcareous Formation Basal Formation

North (Gairezi Facies)

1

South

Upper Pelitic Schist Formation Lower Pelitic Schist Formation

Basal Formation

Basal Formation (Frontier Series)

559 can be correlated with the Frontier Series in the south and that the other Gairezi formations are not represented in the Chimanimani Mountains. As an expansion of the nomenclature t o embrace the entire extent of these rocks the term Zimbabwe Facies of the Umkondo Group is proposed for the western flat-lying sediments and the Moqambique Facies for the deformed sediments to the east. The Zimbabwe Facies

The succession attains a maximum thickness exceeding 3650 m in the south compared with 200 m at Inyanga. The basal sediments in the north are restricted to thin horizons of dominantly fine-grained arenite, but no Basal Formation as such has been recorded from the Chipinga-Lower Sabi area in the south. In the Middle Sabi Valley, the fine-grained, cherty, calc-hornfelses and minor intercalated limestones, typical of the Calcareous Formation throughout, are underlain by a series of arkoses, dolomites, siltstones and quartzites. However, these were considered by Swift (1962) to be part of the Calcareous Formation. Eastwards from here the calcareous rocks decrease in thickness and interfinger with shales, phyllites and ferruginous arenites which comprise the Lower Argillaceous Formation. In the vicinity of the Chimanimani Mountains these sediments may rest directly on the Basal Formation of the Moqambique Facies, in a similar way to that observed in the Gairezi Highlands. The Arenaceous Formation, 760 to 900 m thick in the south, thins to less than 75 m at Inyanga. This monotonous sequence of fine-grained, grey, feldspathic quartzites contains abundant shale intercalations in the north, but few in the south. These arenites pass upwards into phyllites and argillites comprising the Upper Argillaceous Formation in which interbeds of pale, feldspathic sandstone reach thicknesses of 60 t o 90 m. In the Middle Sabi Valley at least 200 m of basic andesite, agglomerate and tuff cap the succession. The Moqambique Facies

The sugary, white orthoquartzites and intercalated chlorite schists, typical of the Basal Formation of the Moqambique Facies, attain their maximum thickness of 2040 m south of Musapa Gap in the Chimanimani Mountains. The chlorite schists wedge out northwards and orthoquartzite, greatly reduced in thickness, is the major constituent of the basal sediments in the Bandula region. Impersistent lenses of orthoquartzite along the southwest margin of the Gairezi-Barue Highlands become interbedded with grey siltstone and quartz-mica schists t o the north. These siltstones represent the most easterly extent of the Inyanga Facies.

560 Although two distinct pelitic schist assemblages, representing variable conditions of deposition, have been recorded from the Gairezi Facies, no such distinction has been made by Moqambique geologists. These schists, comprising the bulk of the sediments of the Gairezi-Barue Highlands and Bandula area, are absent from the Chimanimani Mountains. The Lower Pelitic Schist Formation, composed dominantly of quartz-muscovite schist, quartzite and phyllite, is characterized by numerous lenses and thin horizons of epidote- and tremolite-bearing calc-hornfels. The schists of the Upper Pelitic Schist Formation are more ferruginous and contain abundant chloritoid. Calc-hornfelses are rare whereas tremolitic limestones and ferruginous quartzites are common.

Deposition The fine-grained sediments of the Calcareous Formation, a thin but extremely persistent horizon, were deposited on a flat peneplained surface in a restricted lagoonal environment. Eastwards these calcareous sediments thin and lens out in part against a beach sand barrier marking the western limits of a deeper water, higher energy depository. The lagoon, with its minor algal colonies persisted throughout the deposition of the muddy Lower Argillaceous Formation. The Arenaceous Formation marks the onset of a braided deltaic system prograding into the lagoonal basin (Button, 1976). As progradation continued, a meandering stream system built out across the braided delta, with shallow streams cutting into a flat, alluvial plain. In the south, drainage was essentially t o the northeast, but at Inyanga it was more easterly. In the east, depositional conditions were more chaotic, as shown by the great variety of sedimentary types of the two Pelitic Schist Formations of the Gairezi Facies. Numerous, impersistent lenses of well-bedded calchornfels, some containing possible stromatolites, and a decreased volume of pure quartzites indicate that there were periods of deposition into shallow water.

Structure and metamorphism The c. 2000 Ma uplift of the rocks of the Limpopo belt, which provided a source of sediments for the Waterberg Group, probably also supplied much material for the Umkondo Group. Thick quartz dolerite sills are numerous in the Calcareous and Lower Argillaceous Formations, with fewer and thinner sheets intrusive into the other formations of both the Zimbabwe and Moqambique Facies. Finegrained, chilled dolerites are ssociated with the Upper Argillaceous Formation in the Mount Selinda area. The sills were intruded after sedimentation was complete but before the onset of the first deformation.

561 Palaeomagnetically, the Precambrian dolerites of eastern Rhodesia form two distinct groups, indicating two distinct magmatic episodes (McElhinny and Opdyke, 1964). One group, intruding the Precambrian Shield, is termed the Mashonaland Dolerites and the other, associated with Umkondo sediments is known as the Umkondo Dolerites. The poles of the latter correspond well with those obtained from post-Waterberg diabases and from a 1700 Ma diabase sill cutting the Premier Mine kimberlite (Jones and McElhinny, 1966). This age is close t o a maximum K-Ar age of 1785 f 80 Ma for a hornfels from the Chipinga area and represents the minimum age of the Umkondo Group. The maximum age is more difficult to establish. At Inyanga, Umkondo sediments apparently unconformablg overlie a thin sill of porphyritic dolerite, typical of the Mashonaland Dolerite group, samples of which give a Rb-Sr age of 1850 k 20 Ma. Much of the Umkondo Group has been affected by the 500 t o 600 Ma tectonothermal event, termed the Moqambique Metamorphic Event. The western limit of this more or less coincides with the boundary between the two facies. The eastern parts of the Zimbabwe Facies are only slightly affected, but the Manica Belt (Moqambique Facies) shows well-defined zones of easterly increasing metamorphic grade (Vail, 1966). Chlorite, garnet, staurolite, kyanite and sillimanite isograds were recognized over a strike distance of 500 km. The western part of the Zimbabwe Facies in the south is comparatively undeformed except for very broad low-amplitude folds which plunge to the south. The succession is block faulted with large downthrows to the south. The deformation of the adjacent Moqambique Facies is characterized by a series of almost north-trending, asymmetrical folds in which the long limbs dip gently t o the east and the shorter, attenuated western limbs are steeply inclined and often overturned. In the Chimanimani Mountains the deformation is intense, with recumbent folds and tectonic slices thrust westwards over the comparatively undeformed sediments.

Mineralization The only known mineralization of economic interest in the Umkondo Group occurs in sediments of the Zimbabwe Facies of the Middle Sabi Valley. At the Umkondo Mine copper-ore bodies are located within a sequence of alternating shales and quartzites lying towards the top of an horizon stratigraphically equivalent to part of the Upper Argillaceous Formation. Here, in the shales and siltstones, chalcopyrite and bornite occur as nodules and in thin horizons parallel t o the bedding or as cross-cutting veins and minute stringers. The sulphides in the quartzites occur either as fine-grained interstitial disseminations or as nodules. Horizontal mineral zoning, bornite + chalcopyrite-chalcopyrite-pyrite, occurs over short distances, whereas vertically the intermediate chalcopyrite zone is absent. The copper was intro-

562 duced hydrothermally , replacing stratiform pyrite mineralization. Similar, smaller occurrences have been described from the Devuli prospect and from the Marie Louise group of prospects. Chalcocite, with minor amounts of bornite, chalcopyrite, covellite and native copper, is often concentrated in the vesicular portion of the basic andesites as well as in the agglomerates and the tuffaceous siltstones. Disseminated chalcopyrite, not uncommon in the altered dolerites, also occurs sporadically in thermally metamorphosed sediments adjacent t o dolerite intrusions.

C. Plutonic events in the cratonic areas GRANITIC EVENTS

*

Following the termination of Archaean volcanism and associated tonalitic plutonism at -3200 Ma, granitic emplacement occurred in the Kaapvaal craton at: (1)-3000 Ma; (2) -2700 Ma; (3) -2250 Ma; (4) -2000 Ma, associated with the Bushveld Complex (which is described on p. 599); and (5) post-dating Waterberg sedimentation in Botswana. Granites associated with the two older events are best exposed and known in the eastern Transvaal and Swaziland (Fig. 9.14), where the deeply incised topography has aided the establishment of field relationships. Over the remaining areas of the craton granites are generally poorly exposed on the highveld erosion surface (-1500 m) above sea level. The granitic rocks emplaced at -3000 Ma have been described and their distribution discussed (p. 451). In the southeastern Kaapvaal craton, they, together with various tonalitic and other gneisses, form the basement on which the Pongola Group was deposited. -2700 Ma granites

Granites of this age in the southeastern Transvaal and Swaziland intrude the Pongola Group occurring like the -3000 Ma-old granites, both as sheetlike and sharply discordant intrusives. In contrast to the -3000 Ma-old Lochiel granite, the roof of the 2700 Ma-old, sheeted intrusive is recognizable in southern Swaziland where gently dipping rocks of the Mozaan Subgroup overlie and are intruded by the granitic sheet. Petrologically and chemically this granite is similar t o the Lochiel granite, except for higher contents of K 2 0 and Rb. The discordant plutons of this age in the eastern Transvaal and Swaziland can be subdivided on the basis of their field relationships and petrology into

* By D.R. Hunter.

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0Non-Granitic rocks 0Post-Waterberg (?) Granites 0Bushveld Granite t 1,95 b.y. Gaborone and Palala Granites f2,3b.y. ? Mooihoek Granite Mpageni Granite ‘2.65 b.y. [I3 Kwetta Granite Pongola Granite and 2,6-2,7 b.y.Granites Dalmein Granite 2,9 b.y. Lochiel Granite ! 3 , 0 b.y. Nelspruit Mipmatites Granodiorite Suite and 3,2 b.y.Granites @ Tonalitic Diapirs @ Tonaiitic Gneisses Granites (undifferentiated) Granite Plutona (undifferentiated) Greenatone belts

0

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Fig. 9.14. Distribution of the granitic rocks in the Kaapvaal craton. (Reproduced by permission from Precambrian Research, Elsevier Scientific Publishing C O . ) .

three groups (Hunter, 1973a). The most common, Mpageni-type plutons are coarse-grained, porphyritic and distinguished by the presence of opalescent blue quartz grains and disc-shaped mafic inclusions up t o 2 m in diameter that are studded with large microcline porphyroblasts. The Kwetta-type pluton, of which only one has so far been recognized, is characterized by large microcline insets up to 2.5 cm in length that are sometimes mantled by rims of albite. Two leucocratic, coarse-grained plutons of alaskitic composi-

564 tion constitute the Mooihoek-type plutons which intrude the Kwetta pluton. The relative ages of these last two types to the Mpageni-type plutons cannot be established in the field. Granites cropping out in the Schweizer Reneke and Ventersdorp “windows” have yielded ages of -2700 Ma (Allsopp, 1964) which sets an upper limit for the age of the Witwatersrand Supergroup, the basal beds of which rest with a sedimentary contact on these granites, which are petrologically similar t o the sheeted intrusives in the eastern Kaapvaal craton. The Mashishimala pluton has gross characteristics similar t o the Mpageni-type plutons but has not been dated isotopically. Radiometric ages close to 2700 Ma are reported from pegmatites in the vicinity of the Bandolierkop pluton (Burger and Coertze, 1973), granitic rocks 80 km east of Gaborone (Crampton, 1972) and granites in the Limpopo mobile belt (Van Breemen and Dodson, 1972).

-

225 0-2600 M a granites

The coarse-grained, porphyritic and, locally, rapakivitic Gaborone granite crops out in southeastern Botswana in close spatial association with felsites (the Kanye Volcanic Group). Some authors regard the felsites and granite as co-genetic but others suggest that a considerable interval of time separates the formation of the two rock types (for discussion see Harding et al., 1974). The Kanye Volcanic Group is overlain by siliceous tuffisites, ignimbrites and agglomerates with which are interbedded tuffaceous shales, siltstones, and flows of rhyolite and biotite andesite of Ventersdorp age (Crockett, 1971a). The boundary between the Ventersdorp and Kanye rocks is difficult to identify due, according to Crockett (1971a), to the presence of a wide transition zone in which the Kanye felsites are affected by Ventersdorp gas action. While preliminary palaeomagnetic data (Evans, 1967; Jones et al., 1967) suggest that the Gaborone granite is of approximately Ventersdorp age, the Rb/ Sr isotopic data are difficult to interpret, although there is evidence that isotopic rehomogenization occurred at -2300 Ma. Harding et al. (1974) give an age of 2592 2 75 Ma as the best estimate of the age of coarse-grained granites constituting the Central Assemblage of the Gaborone Granite Complex, but the wide scatter of plots of the Kanye Volcanic Group precludes an estimation of their age. Within Botswana the Kanye Volcanic Group surrounds a central pluton of Gaborone granite, an arrangement suggestive of an original felsite roof t o the granite in a relationship similar t o that seen in the younger Bushveld Complex. The Gaborone granite is subdivided into three assemblages. The rocks of the Outer Marginal Assemblage are regarded as coarsely re-crystallized felsites, the limit of this assemblage being taken where the re-crystallization of the Kanye felsites to a coarse granitoid is visible to the naked eye (Harding et al., 1974). A medium- to coarse-grained, pink or grey, leucocratic granite

565 that is texturally homogeneous constitutes the Inner Marginal Assemblage. Aplites, pegmatites and mafic xenoliths which are characteristic of the Central Assemblage are absent. Medium- to very coarse-grained granite distinguished by the presence of reversely zoned feldspar phenocrysts is the dominant rock type of the Central Assemblage. Aplogranite or microgranite, usually more leucocratic than the porphyritic granites, occurs as sharply defined, dyke-like bodies or with gradational contacts with the porphyritic granite. In the northern Kaapvaal craton a coarse-grained porphyritic granite (the Palala granite), having lithological and chemical characteristics similar to the Gaborone and Bushveld granites, is poorly exposed. Its age emplacement is uncertain.

Post- Waterberg granites Crockett (1971b) considered that the Mmathethe and Moshaneng granites in southwestern Botswana were post-Waterberg in age. Petrologic details of these intrusives are scarce and no geochronological data are available. The Mmathethe granite is hornblende-bearing, while the Moshaneng intrusion consists of leucocratic granite and quartz syenite. The Segwagwa intrusion, located between the Moshaneng and Mmathethe bodies, is represented by scattered outcrops of leucocratic granite and syenite but boreholes have intersected hornblende diorite, pyroxene-mica diorite, gabbro and norite. It is interpreted as a small layered complex, not dissimilar to the Bushveld Complex.

Geochemistry and genesis of the granites Rare earth and other trace-element data are available for the granites in the eastern Transvaal and Swaziland. In general enrichment in Rb is the most noticeable feature in all granites, resulting in low values of the K/Rb ratio (Hunter, 197313). The 3000 Ma-old granites have steep REE patterns with small negative Eu anomalies except in the Dalmein pluton. The sheet-like Lochiel granite has a slightly flatter slope for the heavy REE’s as compared to the plutons. The 2700 Ma-old granites can be subdivided into two groups on the basis of their REE patterns. The Sicunusa, Sinceni and Mhlosheni plutons together with the sheeted Pongola granites have very large, negative Eu anomalies. Whereas their Rb content is similar t o the other plutons, Sr contents are distinctly lower in this group. The Kwetta, Mpageni, Ngwempisi and Mbabane plutons have similar REE patterns with smaller negative Eu anomalies. Although the REE patterns of the 2700 Ma-old granites are similar to the 3000 Ma-old granites, enrichment relative t o mean chondrite is about three times greater in the former (Condie and Hunter, 1976). The model that provides the best fit t o all the trace-element data involves genera-

566 tion of the 3000 Ma-old granitic liquids by partial melting of siliceous, garnet-bearing granulite. Derivation from a crustal parent is supported by the initial 87Sr/86Srratios of the Lochiel granite, if the liquids were not selectively contaminated with Sr (Condie and Hunter, 1976). This model is further supported by the relatively high 6 0 l 8 values (Taylor and Mordeckai, 1975). The origin of the 2700 Ma-old granites is best accommodated in a model involving fractional crystallization of a granodioritic parent magma similar in composition to the average -3000 Ma-old type (Condie and Hunter, 1976). The larger negative Eu anomaly and lower Ba and Sr contents in the Sicunusa, Sinceni and Mhlosheni granites is accounted for by a combination of lower oxygen fugacity, the presence of liquidus K-feldspar and the crystallization of more plagioclase. It is concluded that the major period of crustal thickening in the southeastern Kaapvaal craton terminated at -3000 Ma, creating a tectonically more stable environment within which the Pongola Group accumulated. As a consequence of this change only minor amounts of melting occurred, generating magmas that were undisturbed by tectonic activity and underwent fractional crystallization t o give rise t o the high-K20magmas intruded at -2700 Ma in the southeast Kaapvaal craton (Condie and Hunter, 1976). Data are insufficient to propose models for other granitic rocks in the Kaapvaal craton but similarities in geochemistry suggest that granites dated between -3000 Ma and -2700 Ma were probably derived by one or other of the processes proposed for the eastern Transvaal and Swaziland granites. Uncertainties with respect t o the age relationships between the Gaborone granite, Kanye Volcanic Group and the Ventersdorp lavas place constraints on genetic models. Wright (1961) concluded that all were cogenetic. In his model extrusion of lavas and pyroclastics is followed by the crystallization under this effusive cover of a hypabyssal lacolithic sheet, represented by the Kanye Volcanic Group, beneath which the Outer and Inner Marginal Assemblages formed consequent on recrystallization and differentiation. The final influx of magma crystallized as the coarse-grained Central Assemblage. The development of a thick hypabyssal phase of the composition of the Kanye Volcanic Group would require special conditions as it is reasonable t o expect that the low melting temperatures of granitic magma would yield a coarsegrained rock beneath a relatively thin cover. If Wright’s hypothesis is correct, it might be concluded that the widespread extrusion of Ventersdorp basaltic lava within the Kaapvaal craton could have provided the heat source t o initiate partial melting of sialic crust. Such a model would be in keeping with the palaeomagnetic data that suggest (Evans, 1967; Jones et al., 1967) that the Gaborone granite is of approximately Ventersdorp age. Wright’s (1961) model does not satisfy the geochronologic data that suggest a distinct interval of time between the crystallization of the Kanye Volcanic Group and the Central Assemblage of the Gaborone Granite. The possi-

567 ble existence of an important thermal event at -2250 Ma involving the Inner Marginal Assemblage is also ignored in Wright’s model. An alternative hypothesis that regards the Kanye Volcanic Group as the hypabyssal phase of the Gaborone granite, pre-dating the extrusion of the Ventersdorp lavas, presumes the total removal by erosion of an effusive cover. Harding et al. (1974) very tentatively suggest that a thin planetary crust with a steep geothermal gradient existed when the Kanye Volcanic Group crystallized in order to account for the development of a great thickness of fine-grained siliceous rocks under hypabyssal conditions. Proximity t o the planetary surface and rapid de-gassing of a volatile-rich magma would provide little opportunity for the formation of coarse-grained phases which would ultimately develop at slightly greater depths in zones protected by already crystallized felsites. This model also ignores the apparent time interval, suggested by geochronologic studies, between the Kanye Volcanic Group and the Gaborone granite Central Assemblage.

Synthesis The fragmentary record of radiometric ages suggests that the -3000 Maold granitic event is largely confined t o the southern portion of the Kaapvaal craton, the main locus of emplacement occupying an east-northeasterlytrending zone extending across the craton, whereas the -2700 Ma-old granites are developed (according t o available data) most extensively in the northern portion of the craton and particularly in Zimbabwe (Hunter, 1974b). If this is correct, it would imply that thickening and stabilization of the crust occurred laterally as well as vertically. The depositional axes of the post-3000 Ma-old ensialic basins migrated from southeast to northwest, apparently preceded by a similar migration of the main locus of granitic emplacement. Each basin is developed on the southeastern flank of successive loci and each consistently reflect sediments transport directions from the northern margins. Hunter (1974b) has suggested that this relationship, if valid, reflects a complex interaction between subcrustal heat accumulation and dissipation that involves magmatic transfer to shallower crustal levels and the development of sedimentary basins at the planetary surface. This concept would be in accord with the view that the tectonic setting of the granites emplaced at -3000 Ma and -2700 Ma differs from Phanerozoic environments (Hunter, 1973a) reflecting processes of magma emplacement unique to the Early Precambrian). The gross chemical and lithological similarities between the Gaborone, Palala and Bushveld granites invite comparison. They differ from the earlier Precambrian granites in respect to their enrichment in Fe relative to Mg and Ti, in their lack of association with the development of subsequent sedimentary basins, and their temporal association with felsic volcanism. The Bushveld granites are regarded as the products of reaction melting between

568 rising mafic magma and thick sialic crust (Hunter, 1975). Despite the uncertainty regarding the temporal relationship between the Gaborone granite and the Ventersdorp lavas, the possibility cannot be excluded that the addition of voluminous amounts of basaltic magma t o the crust at this time could have initiated melting of the crust to give rise t o the magma from which the Gaborone granite crystallized. MAFIC LAYERED COMPLEXES

Usush wana Complex

*

This 2870 Ma-old complex crops out in southeastern Transvaal and Swaziland (Fig. 9.15), and consists of two northwesterly trending dykes, 35 km apart and linked to each other by a sill-like intrusion at the base of the Pongola Group in Swaziland. The western dyke is exposed over a distance of about 100 km and has a maximum outcrop width of 20 km in the immediate vicinity of Amsterdam where a raft of Pongola Group rocks occupies the central position of the dyke. The eastern dyke is exposed over a length of 30 km and maximum width of 5 km, with a small raft of Insuzi Subgroup volcanics and sediments preserved within the southeastern portion of the dyke. Gravity data obtained from traverses across the eastern dyke are interpreted as being consistent with a steep-sided dyke structure, but such data from the sill-like intrusive in Swaziland are not sufficient to determine its configuration. Layering in the complex is only imperfectly preserved due t o the intrusion of late quartz gabbro that has swamped the earlier layered phase. Layering is best preserved on the northeastern flank of the eastern dyke in Swaziland. The following succession of layers can be recognized from the base upwards: (1)Medium-grained gabbro - 15 m. (2) Medium- to coarse-grained quartz gabbro - 105 m. (3) Zone of medium-grained to pegmatitic gabbro with disseminated chalcopyrite and pyrrhotite - 1 0 m. (4) Medium- to coarse-grained quartz gabbro - 450 m. (5) Fine- to medium-grained gabbro underlain by discontinuous lenses of serpentinite - 75 m. (6) Coarse-grained hypersthene gabbro with disseminated pyrrhotite and calcopyrite - 1 2 m. (7) Spotted plagioclase rock - 3 m. (8) Coarse-grained hypersthene gabbro with layers of medium-grained hypersthene gabbro - thickness not known. These layers dip southwestwards towards the centre of the dyke at 10".

* By D.R. Hunter.

569

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80

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Fig. 9.15. Distribution of the Usushwana Complex. (Note: Pongola Supergroup now renamed Pongola Group).

On the opposite flank of the dyke a lens of olivine pyroxenite occupies the base and dips northeastwards at 50". It is intruded by quartz gabbro that occupies much of the central portion of the dyke. The southeastern portion of the eastern dyke consists of quartz gabbro, containing bands of titaniferous magnetite up t o 2 m thick, that passes upwards into a magnetite-rich, hybrid gabbro which is gradational into granophyre, immediately underlying the roof pendant of Insuzi volcanic rocks. Quartz gabbro intrudes all the preceding rock types. The western dyke and the connecting sill-like body are composed of similar sequences to that found in the southeastern portion of the eastern dyke. Layering in all these areas is poorly developed being defined mainly by the lenticular bands of titaniferous magnetite. No deposits of economic significance have yet been located within the complex. The grade of Cu-Ni mineralization in the eastern dyke is erratic with mean Cu and Ni contents of 0.36% and 0.05%, respectively. The titaniferous magnetite bands are lenticular with Vz05contents of less than 0.2%.

570 The complex intrudes pre-Pongola granites and the Pongola Group but is intruded by two small plutons of coarse-grained porphyritic granite, dated at -2600 Ma, which disrupt the continuity of the sill linking the eastern and western dykes. The emplacement of the complex has been interpreted (Hunter, 1970) as related to intrusion along two parallel graben or rift-like structures in a similar manner to that proposed by Worst (1958) for the Great “Dyke” of Zimbabwe. The presence of pendants of the Pongola Group in northwesterly aligned structures beyond the present limits of the group is seen as preservation as a result of block-faulting. Mafic magma intruded the resulting graben; differentiated in situ to give the layered sequence, and was followed by a later intrusion of quartz gabbro. Reactivation of the bounding faults and sagging of the graben resulted in the preservation of the dyke-like portion of the intrusive, subsequent erosion destroying the sill-like, layered sequence beyond the limits of the graben. Evidence exists for the presence of major northwesterly aligned faults in the granitic terrane adjacent to the eastern dyke. Evidence for in situ differentiation is not preserved in the western dyke (E.C.I. Hammerbeck, pers. commun., 1974), leuco-gabbro, quartz gabbro and ferro-gabbro being regarded there as separate intrusions.

Mashaba Igneous Complex * This lies about 250 km south of Salisbury, Zimbabwe, at the western end of the Fort Victoria greenstone belt, and, as now exposed, is largely the remains of an ultramafic layered intrusion. Wilson (1968a,b) also describes certain adjacent mafic dyke swarms as late phases of the complex. The country rock is mostly c. 3500 Ma granite-gneiss terrane with infolded Sebakwian Group remnants. The most extensive part of the complex is the main still. This makes up the northwest arm, which dips northeast at 10-30 degrees, and most of the central sector (Fig. 9.16). The main rocks are dunite, harzburgite, both at least partly serpentinized,’and some pyroxenite. In parts of the northwest arm two macro-units can be distinguished (units 2 and 3, Fig. 9.16) with the lower unit capped by a discontinuous layer of olivine pyroxenite. Except for a few small sporadic pockets, the roof rocks of the main sill are entirely ultramafic normally consisting of a thin but persistent zone of partly serpentinized poikilitic harzburgite. High-grade contact metamorphic effects are seen in the overlying country rocks in several localities. Within the central sector, east of Mashaba, are the remains of a further layered sequence which occupies an area of about 10 km2 and in part underlies the main sill (unit 4, Fig. 9.16). This sequence shows a much fuller

* By J.F. Wilson.

571

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Fig. 9.16. Sketch map of the Mashaba Igneous Complex with ring-radial dyke pattern. A = northwest arm; B = central sector; C = northeast arm; D = eastern arm. Mineral compositions in simplified vertical columns from unpublished electron probe analyses by A.H. Wilson. Vertical thicknesses are approximate and maximum (After Wilson, 1968a).

differentiation history. The rocks range from basal serpentinized dunite, through a composite pyroxenite layer into uralitized gabbro, which is quartzbearing at the top. In the southeast of the central sector, a chromititebearing serpentinite, with well-developed fine-scale layering of chromite and olivine, occurs above the main sill (unit 1,Fig. 9.16). Wilson's (1968a,b) earlier explanation for the emplacement of these four layered units by a complicated mechanism of separate injection requires modification. His unit 4 is now best interpreted as the remains of a discrete and separate intrusion related to the same magmatic episode as the main sill;

572 and, as discussed earlier (p. 455, this volume), it now seems more likely that some, or all, of his unit 1 is part of the much older (c. 3500 Ma) Sebakwian Group (Cotterill, 1976,1979). The northeastern and eastern arms are more difficult t o evaluate. The rock types are serpentinized dunites and other peridotites; the eastern arm, has been subjected to intense shearing for most of its length and also contains a major development of talc-carbonate rocks. These two arms, together with the prominent northerly offshoot from the eastern arm west of Fort Victoria, form a crude ring (Fig. 9.16) which is matched on either side of the northeast arm by curved, uralitized and saussuritized, mafic dykes. A modified radial pattern of dykes is also evident, with swarms in the north and south that curve to converge on the ring area. The whole ring-radial pattern is cut in the east by the c. 2600 Ma adamellitic Chilimanzi batholith. Wilson (1968a,b) described the ultramafic arms of the ring pattern as steeply dipping dykes. Alternatively they could be the remains of an eastern extension of the main sill down-faulted between arcuate fractures, with the eastern arm modified by later intense shearing. The complex has been folded first on a northwest-trending axis, then on a northeast-trending axis. The first deformation accompanied major, but limited, southwesterly block sliding, with the base of what is now the northwest arm acting as a lubricating zone. Thus the lower margin of the two sections of this arm is now in thrust contact with the underlying gneisses. The second deformation produced cross warping of the northwest arm and a dome and basin structure in the central sector. In the east of the central sector this was followed by major faulting and further limited block movement southwestwards. Major economic deposits of chrysotile asbestos are developed close to the major thrusts and faults in several parts of the complex (Laubscher, 1968; Wilson, 1968a). Wilson (1968a,b) attributed the deformation of the complex t o the intrusion of the late adamellites, and suggested southwesterly sliding of rocks off the rising dome of the Chilimanzi batholith t o the northeast, with final passive injection into the deformed structures. It now seems better explained as part of the much larger regional pattern envisaged by Coward et al. (1976), affecting the craton during the main periods of Limpopo deformation (Wilson, 1979). Regional considerations and correlations concerning the complex are discussed in Chapter 8, Part B of this volume. The Great “Dyke”

*

The Great “Dyke” of Zimbabwe is a linear mass of mafic and ultramafic rocks which cuts across the Rhodesian Archaean Shield in a direction slightly

* By A.H. Wilson and J.F. Wilson.

573 east of north for about 530 km; its maximum width is 11km. In the present plane of erosion it is not a true dyke, but the remains of four elongate layered complexes arranged in a near straight line (Worst, 1960). The uniformity of the various layers over many kilometres emphasizes the structure which, both in cross-section and longitudinal section, is synclinal. The layers dip inwards from the margin towards the central axis, and, except where modified by later faulting, plunge gently southwards and northwards from the respective northern and southern limits of each complex to produce a boat-like form. Preliminary gravity work across various parts of the Great “Dyke” indicates a range from narrow to broad, cross-sectional V-shapes with the possibility of deep, central, true dyke-like feeders (Podmore, 1970 and pers. commun., 1976). The Great “Dyke” fracture system, however, also involves a number of near northerly trending true dykes and major faults (Fig. 8.11). To the east and west are the quartz gabbros of the East and Umvimeela dykes; farther east the Popoteke and related sinistral wrench faults are in places filled with similar quartz gabbro (Wilson, 1973); and extending south of the Great “Dyke” for some 85 km, to cut the metamorphic rocks of the northern marginal zone of the Limpopo mobile belt, are the various Southern Satellite Dykes. The Southern Satellites include extensions of the East and Umvimeela dykes together with the inner satellites which comprise the Main Swarm and the north-northwest-trending Bubi and Crystal Springs swarms (Robertson and Van Breemen, 1970). The largest of the inner satellites, the Main Swarm, consists of 13 dykes ranging from 100 m to 26 km long and may represent a feeder t o a Great “Dyke”-type of layered intrusion now removed by erosion. The rocks of the Main Swarm are varieties of feldspathic olivine websterites and melanorites whereas those of the inner satellites are largely gabbroic types. Several workers have determined the age of the Great “Dyke” by the RbSr whole rock isochron method. The most precise, on four samples from the Hartley and one from the Wedza Complex, is 2514 k 1 6 Ma (2 sigma) with an initial 87Sr/86Srratio of 0.70261 f 0.0004 (Hamilton, 1977). This is statistically indistinguishable from the two earlier results of 2530 f 30 Ma (Allsopp, 1965) and 2532 f 89 Ma (Davies et al., 1970). Robertson and Van Breemen (1970) determined an age of 2600 f 120 Ma from samples of the dykes of the Main Swarm. From north to south the four complexes constituting the Great “Dyke” are the Musengezi, Hartley, Selukwe and Wedza complexes, which are 43, 312, 96 and 80 km long, respectively. The successions differ in detail in each complex, but all are broadly similar in component rock types and show marked igneous layering on various scales. It is possible to divide each complex into an ultramafic sequence overlain by a mafic sequence. The lowermost part of each ultramafic sequence is not exposed, nor are the tops of the

574 mafic sequences preserved, but the approximate maximum preserved thicknesses are as follows :

Musengezi Hartley Selukwe Wedza

Mafic sequence

Ultramafic sequence

300 m 1150 m 270 m 390 m

3350 m 2200 m 1600 m 1500 m

The cumulus phases in the major part of each uZtrurnuf[c sequence are chromite, olivine and orthopyroxene. The near-monomineralic rock layers formed from these phases are respectively thin, 10-30 cm, chromitites (the “chromite seams” of Worst, 1960, 1964) and thicker dunites and bronzitites. The dunits and bronzitites contain from 1-576 interstitial cumulus chromite, and the base of each bronzitite layer is marked by olivine-bronzite cumulate as a zone transitional between bronzitite and the underlying dunite. The reaction relationship in which cumulus olivine has been replaced by postcumulus orthopyroxenes, is apparent in the upper half of each ultramafic sequence, with the development of poikilitic harzburgite as distinct layers within the dunites. Small amounts of postcumulus clinopyroxene and plagioclase are also found in this association. Figure 9.17 illustrates the succession for the Hartley Complex. Although serpentinization decreases with depth, most of the olivine cumulates are highly serpentinized in surface exposures whereas the bronzitites are largely fresh and unaltered. In consequence, the bronzitites support much more vegetation than the serpentinized dunites and their layers are clearly discernible on air photographs. Jackson (1970), using the earlier work of Worst, recognized major cyclic units, similar to those of the Stillwater Complex of Montana, in all four ultramafic sequences. In the Hartley Complex, in particular, he suggested 1 4 such units. His idealized unit is one in which a chromitite layer occurs at, or near, the base, followed successively upwards by olivine cumulate, olivinebronzite cumulate and finally bronzite cumulate. Many of his lower cyclic units appear to be incomplete, lacking the bronzitite capping. Recent unpublished work on the Hartley Complex (Wilson, 1976), however, has shown that, in many places, narrow poorly developed layers of bronzite cumulate (usually less than 5 cm thick) do cap these lower units. In addition, higher in the sequence, above some of the bronzitite layers where chromitite horizons had not been recognized by Worst, sub-economic chromitite concentrations have been found. Fig. 9 . 1 7 . Simplified vertical succession of the Hartley Complex of the Great “Dyke”, with distribution and composition of the mineral phases. C1 t o C11 indicates position of Worst’s “chromite seams”, Cr/Fe ratios from Worst (1958). Plagioclase compositions determined optically (Wilson, 1976). All other mineral compositions based on chemical analyses.

I TI

576 Wilson’s (1976) work on crystal size measurements on the cumulus olivines and orthopyroxenes of the dunites and bronzitites, respectively, reveals some interesting features. Results from the olivines suggest smaller rhythmic units with the major cyclic units, and these in turn appear to be related to concentrations of chromite. The size variations in the orthopyroxenes from the various bronzitite layers show a striking similarity t o the orthopyroxene size patterns described by Jackson (1961) from the Mountain View area of the Stillwater Complex, with the largest crystals occurring in the middle of the sequence. The symmetrical variation of the mean Cr/Fe ratios of Worst’s “chromite seams” reflects this mean crystal size pattern of the orthopyroxenes and also their composition variations. The topmost bronzitite (Bronzitite No. 1. in Fig. 9.17) itself contains welldeveloped layers continuous over long strikes. These are approximately 20 m thick and Wilson (1976) relates their rhythmic repetition to relative proportions of postcumulus minerals (notably augite and plagioclase), crystal size, and crystal orientation. In all four complexes the upper part of the ultramafic sequence is a feldspathic bronzitite, with augite making its first appearance as a cumulus phase immediately below the mafic sequence boundary. These bronziteaugite cumulates are from 10-40 m thick, but thicknesses differ between complexes, and even within a single complex. The mafic sequence overlies the bronziteaugite cumulate in each complex, and is marked by the incoming of plagioclase as a cumulus phase. It consists of various “Gabbroic rocks” and occurs in the central area of each complex forming the upper part of the boat-shaped structure. Detailed study of this sequence has been done only in the Hartley Complex, but preliminary work suggests similar mineral composition patterns and textural relationships in the other complexes. The Hartley mafic sequence is divisible into three major units, designated Lower, Middle and Upper respectively (Fig. 9.17). This subdivision is based primarily on textural form and the field relationships of mappable units (Wilson, 1976). In a very simplified approach the subdivision is based on the following cumulate distinctions : (1)Upper Mufic Unit - Cumulus pigeonite (now inverted) ferroaugite and plagioclase, with the occurrence of cumulus and postcumulus opaque oxides. (2) Middle Mafic Unit - Cumulus orthopyroxene and augite; postcumulus plagioclase. ( 3 ) Lower Mafic Unit - Cumulus and postcumulus primary orthopyroxene, augite and plagioclase. Further subdivision within each of these units can be made on cumulus textures, structures and mineral compositions. Near the base of the Lower Mafic Unit olivine is found as a cumulus phase with augite and plagioclase, and some cumulus olivine also occurs in a few of the rocks of the Middle

577 Mafic Unit. Higher in the Lower Mafic Unit, the only minerals present are primary orthopyroxene, augite and plagioclase, but cyclic variation of cumulus and postcumulus orthopyroxene allows the recognition of subunits. A feature of some layering is the occurrence of settled orthopyroxene which poikilitically encloses plagioclase. The ranges and trends of the mineral compositions for the mafic and ultramafic sequences of the Hartley complex are summarized in Fig. 9.17. The “chromite seams” of the “Dyke” are a major reserve and rich source of high-grade metallurgical ore (Worst, 1960, 1964). There is also considerable potential as a source of platinum, copper and nickel which occur in a sulphide horizon near the base of the bronziteaugite cumulate in each complex. Chrysotile asbestos, on the other hand, is of limited development and has been mined from only one locality, near the northern end of the Hartley Complex. Schemes for the emplacement of the Great “Dyke” have involved single and multiple intrusion. Worst (1958, 1960) envisaged the production of the large-scale layering in each complex by a process of injection of successive pulses of ultrabasic magma, the composition of which became gradually more basic with final injection producing the mafic rocks. He considered that each pulse had solidified, but was still hot, before the next was emplaced above it. He explained the boatshaped structure as a result of overloading of the floor of each complex, which initiated downsagging and the formation of a graben, and proposed that the lateral extensions of the mafic and ultramafic sequences were removed by this graben faulting. The fractures filled by the East and Umvimeela dykes he suggested might represent marginal hinge zones. Bichan (1968), working on the Hartley Complex, also supported such a mechanism of separate injection involving near complete solidification of each pulse of magma, but proposed a pyroxenite liquid based on the composition of the layers of the upper part of the ultramafic sequence. On various grounds these earlier suggestions on multiple intrusion (and magma compositions) have been superseded by ideas involving crystallisation from a large volume of liquid for each complex, with or without further addition of original liquid t o the crystallizing system (Jackson, 1970; Hughes, 1970; Wilson, 1976). Wilson (1976) from his detailed work on the Hartley Complex, considers that a scheme without further additions of magma, with a physicochemical mechanism similar t o that proposed for the Stillwater Complex by Jackson (1961), is the most acceptable way of explaining the large- and small-scale rhythmic structures. Estimation of the composition of the initial liquid of the “Dyke” has always raised difficulties. Tholeiitic affinities are clearly indicated from mineral compositions but the preponderance of ultramafic rocks Over mafic rocks in each complex poses a problem if a tholeiite basalt parent is assumed. The thickness disparity, on this basis, has been explained by erosion of either

578 a much thicker, or much wider, development of gabbroic rocks. Wilson (1976), however, from the clinopyroxene chemistry of the Hartley Complex estimates that only a very small vertical thickness (perhaps less than 150 m) has been removed by erosion. The possibility of lateral extension is more difficult t o evaluate. The idea of graben faulting to remove lateral extensions is equally applicable t o considerations involving a tholeiite parent, but it is questionable whether the scale of faulting was as great as Worst suggests. Marginal faulting has undoubtedly occurred in many places, but in some areas even this is doubtful (Hughes, 1970). No chilled margin has been recognized in the layered complexes. Hem’s (1950) analyses of fine-grained norites from the centre of the Hartley complex did not involve rocks representing a chill phase - the rocks are clearly part of the layered sequence. The Peregwe dyke of the Bubi Swarm of the inner satellites has, however, yielded a chilled margin. Analysis of this shows 19.96% MgO and is compatible with a magnesium-rich tholeiitic basalt magma. Recent palaeomagnetic work on the Southern Satellite Dykes (Jones et al., 1976) argues against a Great “Dyke” age for the Bubi and Crystal Springs swarms which constitute these inner satellites, but an initial liquid composition of this type for the Great “Dyke” would support the volume disparity of the mafic and ultramafic sequences, and would more easily explain some of the mineral chemistry (Wilson, 1976; Hughes, 1976). Some general aspects of the Great “Dyke” and its fracture system are discussed in Chapter 8, part B of this volume.

Kunene Anorthosite Complex

*

The major portion (88%)of this huge anorthosite mass is situated in the southwestern extremity of Angola where it straddles the Kunene River and extends into South West Africa/Nambia. The mountainous remoteness of the Kunene area, poor access and outcrop paucity within the flat-lying desert to semi-desert environment to the north, have precluded detailed mapping or systematic geological-petrological studies of the entire complex although small segments have been afforded closer attention (Geologists of the Servicos de Geologica 6 Minas and Companhia Mineira do Cunene S.A.R.L. in Angola; Kostlin, 1967; Silva, 1972; G.F.W. Menge, pers. commun., 1975). Other work has been of a broad reconnaissance nature so that high confidence cannot be placed on geological conclusions regarding the complex as a whole. The complex underlies an area of 17,378 km2 of which at least 13%is covered by thick alluvium and G.F.W. Menge (pers. commun., 1975) suggests

* By C.F. Vermaak.

579

that the stratigraphic thickness of the pile is at least 14.6 km in the south. The bulk of the country rocks are assigned to the so-called “Metamorphic Series” of southwest Angola and consist of heterogeneous varieties of massive to banded quartzo-feldspathic rocks, gneisses, granulites and leptites, exhibiting a range of radiometric ages (see below). They enclose scattered rafts and remnants of stable-shelf sediments (pure t o magnetite-bearing quartzites, marbles to calc-silicate rocks, banded jaspilites, phyllitic slates and pelitic schists) and meta-volcanics (predominantly felsic lavas and acid pyroclastics with lesser basic amygdaloidal types altered to amphibolites) whose mutual relationships and connection with the complex are extremely difficult t o establish. The supracrustals clearly represent temporally repeated sequences of comparable sedimentary-volcanic cycles currently exhibiting a disparity of metamorphic and cataclastic overprints. The ancient KasaiAngola cratonic nucleus occupies the northern and central parts of Angola so that increasing mobility would be expected t o occur southwards and westwards towards the cratonic margins where sequential disruption and metamorphism, followed by retrograde periods, will have had a profound effect on the surrounding framework and the Kunene Complex. Metamorphic overprints affecting the complex and its framework thus increase in intensity but decrease in age in a southerly direction. The preponderant outcropping rocks of the complex consist of melanocratic to leucoanorthosites (72.2%) with a more ultramafic marginal facies (3.2%) and granitic rocks (24.6%) thought t o be related to the complex. The marginal facies. Three distinct types of contact relationship are manifest: faulted contacts, interleaving contacts (where there is a coincidence of structural grain between the complex and its surroundings) and crosscutting contacts. Tectonic contacts exhibit basic selvages and widespread epidotization. Interleaving contacts produce hornblende-, garnet- and tremoliteactinolite-bearing foliated anorthositic gneisses, widespread recrystallization, assimilation and epidotization with intense shearing and rheomorphic intrusive apophyses of reddish leucoaplites, granites and syenites. At crosscutting contacts, the anorthosites assume a fine-grained gabbroic or noritic character (garnetiferous to epidote-bearing), being quartz-rich either due to countryrock assimilation or in the vicinity of rheomorphic acid dykes. False basic margins commonly originate from partial assimilation or rheomorphism of basic lavas contained in the country rocks. Rocks attributable to a true melanocratic marginal facies, possibly resulting from a differentiated basal accumulation, appear to be confined to crosscutting contacts. They are c. kilometre-wide margins consisting of peridotites (serpentinites), lherzolites, clino- and orthopyroxenites, gabbros, gabbro-norites and norites (commonly biotite-rich and olivine-bearing). Hornblendites, olivine diabases and hypersthene diabases are probably metamorphic hybrids. Satellite bodies of basic (gabbro or ultrabasic/serpentinite, dunite,

TABLE 9. VI The basic t o ultrabasic rocks of the Kunene Complex (modal analyses and chemical analyses -water after Kostlin, 1967) Chemical analyses 2

7

SiOz

44.72

TiOz A1203

0.73 14.95 0.17 14.08 0.19

FeZ03

FeO MnO MgO CaO Na20 p2os

14.92 7.61 2.09 0.42 0.12

Total

100.00

KZO

Mineral species

feldspars (Ab + An + Or + Ne) 0.21 orthopyroxene 7.18 clinopyroxene 7.27 olivine 5.70 ore (I1 + M t) 0.19 alteration (Biot + Amph etc.) 29.67 An% 4.27 Mg% total femics 0.04 mafic index felsic index

45.47

100,OO

modified diff. index

*

Reference: 1 = Ejau anorthositic gabbro; 2 = Core portion, Otjanjasemo troctolite; 3 = Basal portion, Otjanjasemo troctolite; 4 = Pure dunite, Etengua body; 5 = Anorthositic dunite, Etengua body;

free, recalculated t o 100% --

Vol. norm

Modal analyses (vol. %)

2

1

2

3

3

5

6

48.3

tr.

72.4

59.2

8.5

36.2 3.2 0.3 1.1

8.5 tr, 67.6 15.4 -

67 N.D. 40.8 -

N.D. N.D. 91.5 -

-

-

7

50.4

19.9

67.2

63.2

6.1 42.6 0.9 -

49.5 1.2 26.4 3,O -

10.6 18.1 2.3 18

1.3 -

-

-

-

33.9 0.3 1.3

51.5 0.2 -

70.0 30.0 -

17.6 9.0

66 66 49.6 48.84 24.77

100 96 80.1 30.42 1.00

58-64 N.D. 32.8 -

N.D. N.D. 36.8

68-72 N.D. 51.7 -

N.D. N.D. 100.0 -

56 80 27.6

41.62

4.26

-

-

-

6 = Norite, Ombuku body; 7 = Serpentinite, Ombuku body. Tr = Trace; N.D. = Not determined; * After Von Gruenewaldt, 1973a.

-

-

-

-

-

7

-

581 pyroxenite) character are commonly separated from the main body of the complex and contained within the country rocks. Table 9.VI shows the available chemistry and mineralogy of these rocks. Alternately, rocks with similar characteristics intrude the main complex as plugs or dykes and one could invoke the postulate of a separate ultramafic to mafic magma to explain occurrences of this type. The main complex. The main rock types of the complex are pale massive anorthosites and dark melanocratic troctolites with olivine anorthosite and noritic rocks possibly representing gradational facies between the main types. The petrochemistry and mineralogy of these rocks are shown in Tables 9.VII and 9.VIII. The pale massive anorthosite is a leucocratic palegrey, medium- to coarse-grained rock which shows a progressive chalky whiteness with increasing saussuritization of the plagioclase near surface. The troctolites are dark rocks which owe much of the colouration to metamorphic clouding of the feldspars. They are generally finer-grained and less altered than the anorthosites and commonly exhibit tabular-feldspar igneous lamination. The alteration of these rocks could be a function of the metamorphic environment (see above) since it is clear from their distribution (Fig. 9.18) that the anorthositic and troctolitic rocks respectively predominate in the southern and northern portions of the complex. The interrelationship of the two main rock types is controversial. Simpson and Otto (1960) considered the troctolitic rocks t o occupy elevated ground stratigraphically above the massive anorthosites. Kostlin (1967) regards the pale massive anorthosite as metasomatised alteration products of the country-rock gneisses into which the troctolites and more ultrabasic rocks were sequentially intrusive. Simpson (1970) considers that the metasomatic evidence and argument are not entirely convincing but would support a later sheet-intrusion of the troctolites into the anorthosites. The latest evidence is overwhelmingly in favour of anorthosite complexes being of an intrusive nature (see below, also Carmichael et al., 1974) and an origin for the diversity of rock types has been sought in gravity-controlled crystal fractionation. Thus Stone and Brown (1958) first suggestested early ultrabasic to basic differentiates with fractionation leading to silica, sodium, iron and apatite enrichment. Silva (1972) suggested repetitive gravity stratification of alternating anorthosites, olivine anorthosites (of troctolitic affinity) and troctolites although he failed to provide a sequence of cyclicity and could obtain no idea of the cryptic variation as indicated by the chemistry of 26 samples over a horizontal distance of some 39.5 km of traverse across the central part of the complex in Angola (Picolo area). G.F.W. Menge (pers. commun., 1975) supports a similar mechanism and considers that the troctolites underlie the anorthosites with a rhythmic alternation in 25 subzones (cycles) in some areas.

TABLE 9.VII The petrochemistry of the Kunene Complex A: Average raw analyses of t h e Kunene Complex and differentiation indices calculated on water-free basis

i

Northern Complex (After Silva, 1972) Troctolite Avek S.D. Si02 Ti02 A1203 Fe203

FeO MnO MgO CaO Na20 K2O p205

H2O+ H2O-

L.O.I.

49.48 k1.16 0.35 k0.20 22.91 k1.57 1.74 f0.50 5.46 k1.78 0.09 20.03 6.40 k1.3 8.97 20.82 3.84 k0.20 0.39 20.08 0.06 kO.01 0.61 k0.59 0.04 k0.02 -

No. of Samples

3

Larsen Ind. Mafic index Felsic index Ca Index Mg Index

-4 .O 52.9 32.3 69.8 46.7

C.V. %

2.3 57.1 6.8 28.7 32.6 33.3 17.6 9.1 5.2 20.5 16.7 96.7 50.0

-

Olivine anorthosite

Noritic anorthos it e

Anorthosite

Avek S.D.

Avef S.D.

Avef S.D.

50.87 k0.49 0.69 k0.58 24.36 f0.62 1.38 f0.22 4.06 k0.81 0.05 k0.03 3.92 20.08 9.46 k0.75 4.38 20.45 0.48 $0.49 0.12 k0.16 0.37 k0.20 0.05

k0.20 -

9

C.V. %

1.o 84.0 2.5 15.9 19.9 60.0 2.o 7.9 10.3 102.1 133.3 54.0 400.0 -

49.86 f0.41 1.86 f0.23 19.70 24.59 1.61 f0.12 7.46 f3.14 0.09 fO.09 4.89 f1.50 8.48 f0.57 4.15 f1.32 0.93 f0.32 0.11 fO.ll 0.79 k0.55 0.15 20.10

-

2 0

58.5 34.0 66.0 41.5

-3.3 65.0 37.2 62.8 34.7

C.V. %

0.8 12.4 23.3 7.4 42.1 100.0 30.7 6.7 31.8 34.4 100.0 69.6 66.7 -

53.72 f0.70 0.34 f0.35 27.13 k0.78 0.66 f0.41 0.66 k0.55 0.01 k0.03 0.97 f0.70 10.10 f0.76 4.99 f0.48 0.59 f0.14 0.04 f0.05 0.62 k0.29 0.06 k0.03 -

11 6.8 58.3 35.4 64.5 41.7

C.V. %

1.3 102.9 2.9 62.1 83.3 300.0 72.0 7.5 9.6 23.7 125.0 46.8 50.0 -

583

Southern Complex (After Kostlin, 1967) Quartz diorite

Ave k S.D.

C.V.

49.56

-

1.83

-

16.88

-

2.43

-

11.09

-

0.37

-

2.88

-

4.85

-

3.72

-

2.26

-

0.51

-

1.46

-

%

Troctolite

Ave k S.D. 49.26 kl.ll 0.25 k0.48 25.10 k1.81 1.15 k0.42 3.58 f1.29 0.06 f0.02 5.28 51.97 11.28 k4.13 3.50 k0.51 0.43 k0.18 0.01 f0.009 -

1

-

0.82 f0.44 14

0 82.3 55.7 44.2 17.2

C.V. %

2.2 192.0 7.2 36.5 36.0 33.3 37.3 36.6 14.6 41.9 90.0

-

0.06

-

Noritic anorthosite

-3.4 47 .O 25.6 74.3 52.5

53.6

Anorthosite

Ave k S.D.

C.V.

49.62 k0.22 0.22 20.07 28.56 k0.05 0.59 fO.09 1.66 k0.41 0.03 50.007 1.59 k0.33 12.99 k0.02 3.08 f0.12 0.32 k0.007 0.04 k0.02 -

0.4

%

31.8 0.2 15.2 24.7 2.3 20.7 0.1 3.9 2.2 60.0

Ave k S.D.

C.V.

52.23 k1.29 0.13 k0.07 26.39 52.44 1.12 21.04 0.56 k0.46 0.04 10.03 0.66 k0.54 11.97 21.63 4.91 k1.31 0.42 k0.30 0.01 k0.006

2.5

0.36 f0.15 2 0.6 59.0 20.5 79.5 41 .O

%

53.8 9.2 92.8 82.4 75.0

81.8 13.6 26.7 71.4 6.0

-

41.7

1.45 k0.87 6 4.6 54.2 30.7 69.3 29.2

60.0

TABLE 9.VII. (continued) Northern Complex (After Silva, 1972) Troctolite

Olivine anorthosite

Noritic anorthosite

Anorthosite

B: CIPW norms, Niggli values and differentiation indices Orthoclase Albite Anorthite Nephelene Diopside Wollastonite Hypersthene Olivine Ilmenite Magnetite Apatite Corundum a1 fm C

alk si k mg qz

DI MDI CI MCI Explanation Larsen Index: Mafic Index: Felsic Index: Ca Index: Mg Index: DI:

2.4 37.6 43.6 -

5.6 31.6 30.8 2.2 9.6 15.6 3.2 1.4

-

13.6 1.2 1.2 0.4 -

31.7 36.4 22.6 9.3 116.2 0.06 0.61 -21 .o

36.5 26.2 25.7 11.6 129.2 0.07 0.57 -17.2

28.9 36.9 22.7 11.5 122.0 0.13 0.49 -24.0

46.3 7.2 31.4 15.1 156.7 0.07 0.58 -3.7

35.2 41.5 58.5 65.2

41.2 45.2 54.8 59.1

41.3 50.7 50.8 52.3

47.8 47.8 52.0 52.9

2.4 32.0 41.6 0.4

3.2 44.0 48.8 -

0.4

-

21.6 0.6 1.4 -

-

0.8 1.6 0.6 0.6 -

-

( 1 / 3 S i 0 2 + K2O) - (CaO + MgO + FeO); 100 (FeO + F e 2 0 3 ) / ( M g 0 + FeO + F e 2 0 3 ) ; 100 ( N a 2 0 + K,O)/( Na2O + K 2 0 + CaO); 100 CaO/(CaO + Na2O + K2O); 1 0 0 MgO/(MgO + F e O + MnO + Fe,O,); Differentiation Index (Thornton and Tuttle, 1960);

585

Southern Complex (After Kostlin, 1967) Quartz diorite

13.6 33.6 21.2 19.2 5.8 3.2 2.2 0.9 0.3

Troctolite

Anorthosite

-

1.6 27.2 62.4 1.2 5.6 1.2 0.4 0.4 -

2.4 38.4 48.4 2.4 4.4 2.8 0.2 1.o -

27.7 43.7 14.5 14.1 138.1 0.28 0.27 -18.3

35.0 27.8 28.5 8.7 116.4 0.07 0.67 -18.4

44.1 11.0 36.5 8.4 130.0 0.06 0.56 -3.6

43.5 6.5 35.9 14.1 146.2 0.05 0.42 -10.2

50.6 78.7 30.7 36.1

30.8 35.2 65.0 69.9

29.0 31.6 67.2 68.4

43.7 43.7 53.5 53.7

MDI: CI: MCI: S.D. : C.V.:

2.4 26.4 48.8 1.6 4 .O

Noritic anorthosite

-

15.4 0.4 1.o -

Modified Differentiation Index (Von Gruenewaldt, 1973a); Crystallization Index (Poldervaart and Parker, 1964); Modified Crystallization Index (Von Gruenewaldt, 197 3a); Standard Deviation; Co-efficient of Variation.

586 TABLE 9.VIII The mineralogy of the Kunene Complex A: Compositional variation within the complex

Palgioclase Plagioclase Orthopyroxene Clinopyroxene Olivine

Mouta and O’Donnel, 1933

Stone and Brown, 1958

7 5-93%( 4 5-7 5% An)

5-100%(50-80% An) 0-20%(55% En) 7 0%-Less 75%-5%(85%-65% F o )

-

0.7% 0.16%

Silva, 1972 (42%-59% An) -

(85-55% F o )

B: The mineralogy of various rocks condensed from Silva (1972). (Note that the percentage indicates the number of slides out of the total in which a mineral was observed)

Rock (No. of samples)

Plagioclase

OPX

CPX

Olivine

Biotite

Opaques

Troctolites (5)

essential

accessory

acessory

other

other

47% An(42-55%)

80%

80%

essential F073(62-79%) 100%

100%

100%

Anorthositic troctolites (18)

essential

accessory

accessory

other

othei

51% An(46-56%)

83%

72%

accessory F065(52-85%) 100%

100%

100%

Noritic anorthosites ( 2 )

esse n t i a1 54% An

essential 100%

essential 100%

-

other 100%

other 100%

Anorthosites (22)

essential 51% An(46-59%)

accessory 32%

accessory 59%

rare acces. 36%

other 54%

other 91%

Quartz diorite to granodiorite (2)

essential

other

other

-

other

other

50%

50%

100%

100%

35% An

C: Cryptic variation within the complex Source

5 Complex Ctr. 4 3 2 1 Complex Edge

Stone and Brown, 1958 (North)

Simpson and Otto, 1960 (South)

Kostlin, 1967 (Traverse 1)

Plagiocl.

Olivine

Plag.

Oliv.

Opx.

Plag.

Olivine

50% An 55%An 66%An 76%An 80% An

65%Fo 80%Fo 84%Fo -

61%An 62%An 63%An 67%Fo 67%An

87%Fo 83%Fo 79%Fo 62%Fo 58%Fo

83%En 80%En 77%En 74%En

59,3%An 67.5%An 70.0%An 68.5% An 65.5%An

76,0%Fo 81.5%Fo 79.0%Fo 73.0% F o 64.5%Fo

587

Simpson and Otto, 1960; Simpson, 1970

Menge (Pers. Commun., 1975)

(Anorth) 100%(55-66% An) (Tract) 70-95%(54-80% An) 0-12%(85-21% En) 0-5% 1-30%(87-58% Fo)

(57-67% An)

-

(79-73% En)

-

(78-53% Fo)

Amphibole

Apatite

Calcite

Chlorite

Prenhite

Spinel

Sphene

Quartz

Orthoclase

Zircon

other

-

-

other

-

-

-

-

-

-

20%

80%

other

other

other

other

44%

22%

11%

28%

-

55%

-

-

-

-

other 54%

other 27%

other 27%

other 27%

other 27%

other 18%

other 23%

-

-

-

other

other

-

other

-

-

-

100%

microperthite 100%

other

100%

accessory 100%

Kostlin, 1967 (Traverse 2) Plag.

Olivine

66%An 66%An 64%An 67%An -

72%Fo 73%Fo 65%Fo 70% Fo

5 0%

other

100%

588 The calculation of various differentiation indices based on averaged analyses and on their norms does not provide a clear-cut confirmation (Table 9.VIII) of the cyclic sequence, especially with regard to the possible intermediate differentiates, although it does confirm Menge’s end-member sequence. The inadequacy of the indices can be related t o averaging of random sampling over a variety of cycles without the removal of cryptic variation and clearly needs to be undertaken for a single cycle. Limited studies of cryptic variation (Table 9.VIIIC) show conflicting tendencies and serve to aggravate uncertainties regarding differentiation trends of the dominant magma-type under a changing environment (see below). Fractionation towards sodium enrichment appears universal, with possible contamination or metamorphic sodification near the contacts (Table S.VIIIC, Kostlin traverses). Suggestions of iron-enrichment (Stone and Brown, 1958) are contradicted by upward magnesium enrichment elsewhere, although this could be related to random sampling without cognisance of cyclicity, since the centrally disposed titaniferous magnetite occurrences (Fig. 9.18) tend to favour iron-enrichment. The morphology and relationships of these titaniferous magnetite bodies is not clear and they could be either layered or plug-like. Widespread sampling of these rocks at 36 localities averaged Fe 49.56 f 1.53%; Ti02 18.71 f 3.49%; Si02 2.04 2.00%; A1203 4.22 ? 1.15%; MgO 1.92 f 0.97%; V205 0.35 f 0.20% and P 2 0 5 0.064 f 0.0031%. They are virtually confined to the stable and unaltered northern portion of the complex whereas the inexplicable correlative enrichment trends of sodium and magnesium in the south may be a result of metamorphic overprinting within the mobile conditions at the cratonic margin. The unavoidable truth is that our understanding of the geological and petrochemical aspects of this huge mass is currently too restricted with the limited data available. We need also to understand the environment of intrusion and the magma source in order to formulate predictions regarding the complex, as will be indicated below.

*

Possibly related acid rocks. Stone and Brown (1958) considered rafts of pyroxene-bearing gneisso-granites consisting of quartz, 50%; microcline + oligoclase, 40%; orthopyroxene (altered to amphibole and biotite) + apatite, lo%, to be contaminated rocks that formed an ephemeral roof to the complex. Silva provides aiz analysis of one rock (Table 9.VII) and descriptions of granodiorites to quartz diorites (Table 9.VIIIB) although the presence of normative and modal orthopyroxene suggests the charnockitic/mangeritic affinity commonly associated with anorthositic complexes. Cochran et al. (1975) consider that coarse red porphyritic t o pegmatitic granites and lesser pink fine-grained feldspathic microgranites and aplites should be assigned to the complex (2.6% of complex, Fig. 9.18) but their relationship has not been documented. They consider certain porphyritic granites (19.0%) and undifferentiated fine-grained pink biotitic granites (3%)to be either late phases of

589 the complex or post-complex types (Fig. 9.18). G.W.F. Menge (pers. commun., 1975) also relates granitic rocks to the complex and considers that some syenites and quartz syenites, previously considered as rheomorphic products, are acid differentiates of the complex.

Structure. Structural studies have been confined t o South West Africa/ Namibia (Kostlin, 1967; G.F.W. Menge, pers. commun., 1975) where the central anticlinal or dome-like feature contrasts with the interleaving structure of the low ground to the northwest. Menge cdnsiders that the entire complex may be characterized by structural differences of this type although both these structures appear t o be conformable with the regional (N75”E) grain of the enveloping gneisses. The dome-like feature dips outwards (north and south) at rather low (22”-25”) angles. Dykes follow north-northwestsouth-southeast and less frequently west-northwest-east-southeast directions while faults follow a northwestsoutheast direction in keeping with the regional stress pattern. The fault direction also determines the course of the Kunene River t o a great extent. Many authors have noted the disparity between the east-west structural grain of the smaller southern portion of the complex and the north-trending aspect of the rest of the complex. Recent mapping (Cochran et al., 1975) shows this separation to be tectonic, with a large east-west fault (Fig. 9.18) uplifting the “basement” to the north and with granitic rocks related t o the complex intruded ] arallel to, and at a fixed distance, north of the fault. This fault appears to be displaced by a northwest fault at the Kunene River (which the granite follows), the main displacement then continuing westwards along the course of the river. Little structural data exist for the north-trending portion of the complex but dykes and faults follow conjugate east-northeast and north-northwest directions. Reasons for the eastward swing of the complex at its northern extremity, are unknown. It will be shown below that the Kunene Complex has suffered at least four cycles of metamorphism and granite intrusion, to which should be added sodalite-bearing carbonatite dykes (G.F.W. Menge, pers. commun., 1975, advises that the Epembe nepheline syenite has an age of c. 1100 Ma based on the determinations of Burger, in press), pre-Permian porphyritic hornblende granites and microgranites (Vermaak, 1972a,b) and quartz dolerites of unknown age. It has been found (Bridgewater and Windley, 1973) that minerals in or near anorthositic masses have lost their radiogenic daughter elements so that existing dating (on biotite) could be untrustworthy in separating the various intrusive events around the complex. The Oendeolongo-Khoabendus (1250-1100 Ma), Bembe-Nosib (850750 Ma), West Congolean-Damaran (650-550 Ma) and Karoo (225-125 Ma) sedimentary cycles followed the complex, while Kalahari sands cover its eastern extremity. Only essential data are documented for these cycles

590 Table 9 .X) but detailed lithologies are available elsewhere (Mouta, 1954, 1955; Kopershoek, 1964; Stanton et al., 1963; Martin, 1965; Gij, 1970).

Genetic problems. The evidence is overwhelmingly in favour of all anorthosite complexes being of magmatic origin but their strikingly collective individuality of character poses unique problems concerning their genesis, magmatype and environment of intrusion. Their individuality is reflected by structural doming caused by forceful upward intrusion (reminiscent of granitic diapirs), by the invariable occurrence of mafic borders, by their dominantly anorthositic character, the presence of titaniferous magnetite ore minerals, and by the presence of a pyroxene-bearing (charnockitic, mangeritic) granitic, granodioritic or quartz syenite component. They are all Precambrian in age, with low initial strontium isotopic ratios (87Sr/86Sr= 0.703--0.706) suggesting a mantle derivation, and occur within mobile high-grade terranes. At least one of the regional metamorphic events has affected the constituent minerals of anorthosite masses (see also Table 9.VIIIB) with: (a) partial but extensive epidotization, recrystallization and cataclasis; (b) saussuritization, sericitization, chloritization and sodification of the feldspars to An30--35 ; and (c) serpentinization and hydration of the mafic minerals. Nevertheless two classes of anorthosite are recognized : (1)The Adirondack type with relatively sodic plagioclases (An42-45)and minor content of mafic minerals (<5%olivine). (2) A type intermediate between layerd intrusions such as the Bushveld or Stillwater complexes with plagioclases between 1 0 to 30% mafic minerals, prominent olivine in troctolitic anorthosites, and cumulate textures in cyclic gravity-controlled mineral stratification. The Kunene Complex clearly belongs to this intermediate type. Various genetic models are available from which anorthositic magmas can be derived from admissible deep-crustal or mantle sources low in radiogenic strontium. Yoder (1954, 1956, 1968) suggested a basaltic source subjected to high (5-10 kb) water pressure to produce the plagioclase-rich melt, with the residual alkaline liquids producing the charnockitic component, although the absence of primary hydrous phases within the actual rocks is difficult to explain. Green (1969) suggested partial melting of an anhydrous rock, chemically equivalent to andesite (quartz diorite), to produce the anorthosite which would have a complementary comagmatic quartz syenitic to granitic (charnockitic) magma derived from the same source. Others workers have postulated essentially a subalkaline basaltic source although Bridgewater (1967) and Upton (1971) have suggested a primary alkaline olivine basalt magma but magmas of this alkaline affinity may be restricted t o Greenland (Emslie, 1973). Hamilton (1964, 1966) noted a possible connecting link between the iron-enrichment (tholeiitic trend) and high-alumina (calcalkaline affinity) characteristics of the subalkaline anorthositic suites. He suggested that addition of 25-35% plagioclase of sodic t o intermediate com-

TABLE 9.IX Classification of the Kunene rocks according t o the computerized system devised by Irvine and Barager (1971) Object

1. Normative nepheline, normative hypersthene, (cation norm)

2. Si vs. Na,O + K 2 0 to indicate alkaline vs. subalkaline tendency

3. A1203 enrichment: calc-alkaline vs. tholeiitic tendency 4. Subalkaline rocks: K-rich or K-poor

5. Ternary Ne’-Q‘--Ol’, to indicate alkaline vs. subalkaline tendency

6. Alkaline rocks: separation of either sodic or potassic alkali olivine basalt series

Explanation:

South area (Kostlin Samples)

North area (Silva Samples)

Overall

Equations or inequalities (Irvine and Barager, 1971)

Nepheline in 44.4% and hypersthene in 55.6% of the analyses (27)

Nepheline and hypersthene each in 50% of the analyses (26)

Nepheline in 47.2% and hypersthene in 52.8% of the analyses (53)

(for symbols and explanation see below)

69.2% subalkaline 30.8% alkaline

82.7% subalkaline 17.3% alkaline

rock is subalkaline if: S > 13.3539 x A‘ + (1.2030 X As -(1.5188 X lo-’) A4 + (8.6096 X lo-’) A3 - 2.1111 A’ + 3.9492A + 39.0

100% calc-alkaline

96.2% calc-alkaline 3.8% tholeiitic

98.1% calc-alkaline 1.9% tholeiitic

rock is calc-alkaline if: A1203 2 1 2 + 0.08 P where P = 40-1 00%

8.3% are K-rich, 20.8% are K-poor, the rest are “average”

none are K-rich, 46.1% are K-poor, the rest are “average”

4.0% are K-rich, 34.0% are K-poor, the rest are “average”

rock is K-rich if: XAb’ < -(1.3482 x 10-6) X& + (2.1157 X X;, - (1.2000 x 10-2) XO,’ + (3.0554 x lo-’) Xhrr -3.6890 X&,, + 20.4400 Xorr + 0.35. rock is K-poor if: XAb’ 2 9 1 . 8 4 9 6 x x&r’ + (2.1187 x X5 ’ - (9.3497 x 10-3)x$:, + + (2.0202 x 10-1) x;,. 2.2698 X&,‘ + 13.4379 Xo,’, + 20.2. Given that XO,’ + XAb‘ + XAnt = 100.

96.2% subalkaline, 3.8%alkaline

i

55.6% alkaline, 29.6% subalkaline, 14.8% not applicable (of latter 75% contain Ne)

100% are extremely sodic

From cation norm: Q =Quartz; Or = Orthoclase; Ab = Albite; Ne = Nepheline; An = Anorthite; Hy = Hypersthene; 01 = Olivine;

73.1% alkaline 19.2% subalkaline, 7.7% not applicable (of latter 50% contain Ne)

64.2% alklaine, 24.5% subalkaline 11.3%not applicable

rock is subalkaline if: X N ~d’ 1.5 XQ’ where Xol’ = 40-100% X N ~<, 15 + 0.8889 XQ‘ where Xol’ = 0-40% given X N ~+, XQ’ + Xol, = 100.

96.2% are extremely sodic

98.1% are extremely sodic

rock is sodic if: XAb’ > (1.3162 x lo-’) x&,’ - (1.3525 X low3)Xbr, + (5.3479 x 1 0 - 2 ) x&, - (1.0154) X&,+ 9.2090 X&,, - 28.6217 XO,, given that xor’ + XAb’ + XAn’ = 100.

Ab’ = Ab + 513 Ne; 01’ = 01 + 3/4 Hy; Ne’ = Ne + 3/5 Ab; Q’ = Q + 2/5 Ab + 1 / 4 Hy; X = percentage in ternary plot;

P = 100 An/An + Ab’; S = SiOz, wt. %; A = NazO + K,O, wt.%.

Q

‘p u1

CD

co

I

UI

CD I&

595 positions to a basalt magma could produce highalumina basalts or andesites and that sequential differential separation of the anorthositic portion would leave a magma capable of fractionating along the typical tholeiitic trend. The main volume would thus be a troctolite-leucotroctolite mass with relatively small volumes of iron-enriched differentiates (Emslie, 1973). Morse (1972) envisaged the separation of both a large volume of silica-undersaturated plagioclase-rich magma and a saturated troctolite-syenite magma from a highalumina basalt source situated below the level of final emplacement. With the environment of emplacement of subalkaline magmas in mind, Emslie (1973) examined the chemical characteristics of anorthosite massifs, utilizing the descriminatory techniques employed for the separation of the various volcanic suites as formulated by Irvine and Baragar (1971). Emslie realized that anorthosite masses are probably of differentiated cumulate origin (consequently removed from parental compositions) but they are nevertheless products of igneous crystallization which would establish important constraints on source-magma compositions. The results indicate a chemical link between the anorthosite suite (including silicic rocks, without implying a cogenetic relation) with the calc-alkaline association in most respects, except for the iron enrichment relative to magnesia and alkalis. Emslie interpreted these data and the generally high K 2 0 levels, t o indicate that although magma generation originated at depth due to orogenic agencies (but possibly at considerable horizontal distance from the most active portion of the orogen), the magmas may be emplaced in crustal environments either near the zone of tectonic activity or in the cratonic hinterlands, the latter environment then being consistent with interpretation of the intrusions as being anorogenic. With the same object in mind, the available analyses of the Kunene rocks were examined. Both the CIPW and cation norms determined for the 53 analyses showed either nepheline (characteristic of the alkaline suite) or orthopyroxene (characteristic of the subalkaline suite) to be present in about equal proportions (Table 9.IX, section 1) in both the southern (Kostlin samples) and the northern (Silva samples) areas. Neither of these minerals is compatible with the alternate suite and their equal distribution in both areas precludes any analytical error. The results of the Irvine and Barager (1971) classification tests are shown in Table 9.IX which confirm the dual subalkaline (section 2; calc-alkaline, section 3, with low t o average potassium content, section 4) and alkaline (section 4,with a strong sodic alkali olivine basalt affinity, section 6) tendencies. However, it is of interest to note that 50% of the rocks containing normative hypersthene (cation norm) are alkaline according to the Q’-Ne’-01’ ternary (Table 9.IX, section 5). In K20/ Si02 plots (not shown) world-wide anorthosites show K 2 0 enrichment compared with calc-alkaline volcanic suites (Emslie, 1973) which the Kunene rocks fail to satisfy (in SiOz ranges of 5 0 - 6 0 wt. %, most anorthosites show up to 1.5 wt. % K 2 0 compared with up t o 0.8 wt. % K 2 0 for the Kunene

596

NORMATIVE PLAGIOCLASE

Fig. 9.19. A. Plot of total iron as FeO vs. MgO in rocks containing >85% normative feldspar. For explanation see text. B. Plot of colour index (sum of 01 + Hy + Di + Mt + I1 + He, cation norm) vs. normative plagioclase (100 An/An + Ab + 5/3Ne) for the Kunene rocks and their relation to rocks of the calc-alkaline suite. C. As B above but showing the relationship of the Kunene rocks to the alkaline suite. Note: open circles = Kostlin samples; filled circles = Silva samples.

597 rocks). Also total iron as FeO/MgO plots (Fig. 9.19) for Kunene rocks containing more than 85% feldspar do not show the iron enrichment (MgO/ FeO > 4 : 6) of the world-wide anorthosites (Emslie, 1973). FeO/MgO ratios are inferred to indicate the composition of the intercumulus fluid in equilibrium with a large amount of cumulus feldspars because of limited diffusion with the main magma volume, but this need not be true if the coarse grain size and low packing density of the cumulates are considered. Seen as a whole, the Kunene rocks are extremely alumina-rich (92.5% are >20 wt.% A1203;50.9%> 25 wt.% A l , 0 3 ) with a reasonably high Na20 content (4.0 ? 1.1 wt.%) compared with K 2 0 (0.47 ? 0.34 wt.%). It is, however an excess or lack of silica (Ave. 51.2 2.3 wt.% Si02) which determines whether the rock is calc-alkaline or alkaline. Unfortunately the data necessary to seek a cause for silica under-saturation in some samples are not available and would, in any event, be beyond the scope of this review. It should be noted, however, that whereas the world-wide volcanic association of tholeiitic and alkaline types is a common one, an association of calcalkaline and alkaline types is very rare. This probably lends some uniqueness to the Kunene rocks although these conclusions are based on limited analytical results. In both tholeiitic and calc-alkaline associations with alkaline rocks, the vexed problem of whether two independent magmas are involved or whether the one is a derivation of the other remains, although any notable departure or slight variation in either direction from the condition of silica saturation of the parental magma will normally persist in the composition of the end differentiates. It could also be suggested that the iron-enrichment postulated in the Kunene rocks could be a feature of their alkaline component rather than that of tholeiitic fraction.

*

Regional setting and age. Studies of the structure of the western seaboard of southern Africa indicate that major northeast t o east-northeast geosutures cross the coast and extend across the continent. Sediments in basins paralleling the coast (West Congolean) are diverted into these structures, the bestknown being the Damara sequence of central South West Africa/Nambia (Pan-African orogeny). Lesser-known and smaller geosutures include one originating around Luanda in Angola and another centred around the Kunene River (Kunene geosuture) which is the northeasterly extension of the offshore horst-faulted Walvis ridge. When extended on t o the continent these linear faults constitute what has been termed (Vermaak, 1972b) the N60"E Serpa Pinto line (controlling the intrusion of pre-Permian granites and lavas around Serpa Pinto in Angola) and the more northerly N50"E Mocamedes line (controlling the kimberlite-carbonatite intrusions of central Angola; see Figs. 9.18 and 9.20). In keeping with the major world-wide 1800-1200 Ma anorthosite event (Bridgewater and Windley, 1973), the intrusion of the Kunene Complex was fault-controlled (Serpa Pinto line) within a high-grade mobile belt terrane (Kunene geosuture) whose high geo-

598

0"

10"

2 0"

3 0"

1

I

10"

2 0"

I

1

30'

Fig. 9.20. The Kunene Complex within the tectonic framework of Africa, showing the Mocamedes kimberlite-carbonatie line and the Serpa Pinto line and their relationship to the Kunene Complex and Walvis Ridge.

thermal gradient was possibly responsible for generating the anorthositic magma. The southern portion of the complex thus follows the roughly eastwest grain of the mobile belt and was metamorphosed by continued mobility within it, while the northern extension of the complex on to the KasaiAngola craton was determined by its north-trending structural grain. Considering the scanty evidence regarding the character and sequence of geological events in the vicinity of the complex, most conclusions regarding age must be hypothetical and tentative. The earlier dating was on a muscovite from a pegmatite intruding the complex, which gave a Rb/Sr minimum age of 1260 90 Ma (Simpson and Otto, 1960), subsequently recalculated to 1190 Ma (Cahen and Snelling, 1966).Dating by De Carvalho (1969) suggested an age between 1700 and 2080 Ma, the latter being the age of the surrounding

*

599 metamorphic “basement”. Later work (interpretation of data from H.L. Allsopp and A.J. Burger by G.F.W. Menge, pers. commun., 1977) on a hornblende hyperite and a gneiss intruded by the complex, yielded a maximum age of 1500 Ma. Statistical studies by the author of 47 radiogenic dates (Cahen and Snelling, 1966; De Carvalho, 1969) from southwest Angola yielded frequency groupings and averages (column 2, Table 9.X) which closely fit the known metamorphic events and intervening sedimentary cycles in Angola and South West Africa/Namibia. Based on this assessment, a time-slot of 1600-1350 Ma not only fits the latest maximum 1500 Ma date (see above), but is also in keeping with the world-wide 1800-1200 Ma (average 1600 k 200 Ma) anorthositic event (Bridgewater and Windley, 1973). Thus apart from the early Precambrian, the intrusion of the Kunene Complex follows two metamorphic maxima (Table 9.X) at 2216 f 224 Ma (orogenesis 2250-2000 Ma) and the more widespread (statistical frequency) 1628 t 84 Ma of the Haubian episode (1750-1650 Ma) during the middle Precambrian. Bushveld and Kindred Complexes

*

The Bushveld Complex, a major source of a diverse suite of economically viable ores, underlies an area in the Kaapvaal craton of about 67,000 km2 of which some 45% is exposed. Formerly thought t o be a simple lopolith (Hall, 1932), the Complex is now regarded as comprising four separate intrusions; the northern (Potgietersrus) Compartment, the eastern Transvaal compartment, the western Transvaal compartment and the far western Transvaal (Goudini) compartment. These intrusions are located within the depositional basin of the Transvaal Supergroup. In the western compartment the intrusions are semi-concordant with the underlying sediments, except for the faulted area west of the Pilanesberg alkaline intrusive where the base of the complex trangresses the Magaliesberg quartzites of the Pretoria Group and rests on the underlying shales (Vermaak, 1970). Near Potgietersrus, the northern compartment transgresses some 1.75 km of a northward thinning Transvaal sedimentary sequence and eventually rests on basement granites and gneisses. In the eastern Transvaal, the Complex cuts down in a northerly direction through 5 km of the Transvaal succession from the Dullstroom lavas to below the Magaliesberg quartzites and uplifted the sedimentary sequence it transgressed, so that these now form part of the roof of the mafic phase of the Complex (Button, 1976). These transgressive relationships negate an extrusive origin for the Complex (Daly, 1928; Cousins, 1959; Feringa, 1959) and an origin linked t o in situ transformation of the Transvaal Supergroup (Van Biljon, 1949). Gravity data suggest definitely four but possibly as many as six feeders to

* By C.F. Vermaak and C.A. Lee.

TABLE 9.X A possible dating of the Kunene Complex in the light of Angolan and South West African geochronology (after Cahen and Snelling, 1 9 6 6 ; Clifford e t al., 1 9 6 2 ; Martin, 1 9 6 5 ; De Carvalho, 1 9 6 9 ; Vermaak, 1 9 7 2 b ; and G.W.F. Menge, pers. commun., 1975). Period

Metamorphic events (S.W. Angola)

Late Precambrian

514

_+

1 7 Ma ( 5 ) *

late sedimentation cycle 550-650 Ma

Middle Precambrian

Angolan stratigraphy

S.W.A. stratigraphy

Lithology, Metamorphism

Orogenesis 455-520 Ma

Orogenesis 450-550 Ma

(Pan-African)

West Congolean Damaran Schisto-greseux Mulden formation SchistoUpper Otavi calcaire (Tsumeb) Tilloid Chuos tillite Haut Shiloango Lower Otavi (Abenab) Tilloid Glaciation

quartzite, arkose, slate dolomitic dolomite, quartzite, shale -

727 5 27 Ma ( 3 )

Orogenesis 650-750 Ma

Orogenesis 580-660 Ma

early sedimentation cycle 750-850 Ma

BembeSansikwa

Nosib

1 0 5 6 5 9 8 Ma ( 4 )

Orogenesis 950-1100 Ma

Orogenesis 950-1100 Ma

late sedimentation cycle 1100-1250 Ma

OendeolongoLukosian

KhoabendusDordabis

Angola : basal conglomerate sandstones, grey Wac kes schists; upper quar+zites banded ironstones S. W . A . : quartzites, magnetite quartzites, slates, felsic lavas, tuffs

1 3 2 3 5 7 6 Ma ( 1 0 )

Orogenesis 1250-1 35 0 Ma

Orogenesis (?) 12 50-1 350 Ma

(Kibaran)

mafic lava, quartzite, arkose, schist, phyllite, conglomerate, grit, dolomite

601 TABLE 9.X (continued) Period

Metamorphic events (S.W. Angola)

Angolan stratigraphy

S.W.A. stratigraphy

1350-1600 Ma

Granites Kunene Complex (?)

Granites Kunene Complex (?)

1628 k 84 Ma (19) (1600-2000 Ma)

Orogenesis 1 650-1 750 Ma post-orogenic granites in northern Angola 1845-1980 Ma (Mayombian)

Fransfontein granite 1700 Ma

2216 f 224 Ma (6)

Orogenesis 2000-2250 Ma

Orogenesis (?) 2000-3250 Ma

early sedimentation cycle about 2250-2500 Ma

S.W. metamorphic series, Luiza. cycle 1910-2120 Ma

Epupua, Haub Marienhof formations

N . W . Angola. Zanidian >2500 Ma N.E. Angola. (Kasai) charnockitic cycle with sedimentation 2500-2650 Ma Luanzi (Kasai) cycle >3270 Ma

Abbabis for. mation (?) unknown

_______________Early Precambrian

Lithology, Metamorphism

(Haubian)

metasediments, felsic lavas (granulites), mafic lavas (amphibolites)

* Number of age determinations in parenthesis. the Complex but more information is needed in this regard. The form of a compartment overlying each feeder (feeders presumably connected in depth) is that of a funnel-shaped intrusion although this appears t o be a function of the level of crustal exposure (see later). In the eastern and western compartments, such funnel-shaped intrusions are tilted towards the centre of the main Complex, approximately half of the funnel being covered by the acid

602 phase and younger sediments in each case. In the exposed portions dips are invariably centripetal at an average of about 10" (5"-25", with local dips up to 85" due t o tectonic modification). Hunter (1976) has sketched many of the enigmas concerning the regional setting, level of emplacement and morphology of the Complex, which need not be reiterated here. One of the main problems of the mafic sequence concerns the origin of the diversity of rock types and their layering. An origin based on the differentiation of the magma at depth and subsequent intrusion of separate differentiates without further in situ fractionation (Lombaard, 1934; Schwellnus, 1956; Coertze, 1958,

Fig. 9.21. The geology of the eastern compartment of the Bushveld Complex (after Cameron, 1971; Molyneux, 1974; Schwellnus, 1956; Von Gruenewaldt, 1968, 1973b; and the authors).

603 1970) is neither consistent with the extreme lateral persistence of the layering and the lack of crosscutting relationships, nor with the geochemistry, petrology and demonstrable cyclicity of the rocks. It is thus generally accepted that the Bushveld mafic sequence consists of a pile of magmatic sediments either cyclically or non-cyclically deposited.

The Bushveld mafic sequence The average thickness of the eastern (Fig. 9.21), western (Fig. 9.22) and Potgietersrus (Fig. 9.23) mafic sequences is 8000 ? 750 m (Fig. 9.24). The stratigraphies of the eastern and western compartments are strikingly similar (Vermaak, 1976b) but that of the Potgietersrus area is very different (Van der Merwe, 1976). Modern sub-divisions of the Complex are based on its magmatic-sedimentary characteristics and on the decreasing-thickness ranked sequence: zone-subzone-suite-unit-layer. The basis of stratigraphic subdivision is the cyclic unit, the limits of which are defined by the basal appearance or upper disappearance of a given cumulus phase (Cameron, 1970; Vermaak, 1976b). The largest subdivisions represent major appearances or disappearances of cumulus minerals, which occur at the base and the top of a cyclic unit respectively.

Fig. 9.22. The geology of the western compartment of the Bushveld Complex (after Coertze, 1958; and the authors).

604

Fig. 9.23. The geology of the Potgietersrus compartment of the Bushveld Complex (after De Villiers, 1970; Van der Merwe, 1976).

Lower zone. Rimming the Complex is a marginal subzone of fine-grained norites up to 245 m thick, containing numerous metasedimentary inclusions. Where thin, this subzone may represent a contaminated chill forming a contact envelope of variable composition around the Complex depending on the stratigraphic position of the layered mafic rocks in contact with the metasediments; however, when thick it is differentiated into 8 cyclic units of ratio-layered orthopyroxene-plagioclase cumulates. The overlying basal subzone consists of beheaded cyclic units of olivine and orthopyroxene cumulates (harzburgites-pyroxenites) which gradually become less mafic upwards where plagioclase becomes cumulus; the final two units have a leuconorite capping. The bulk composition of the basal subzone is ultramafic, with komatiitic affinities (Vermaak, 1976b). In the eastern and western Bushveld, the subzone attains a thickness of 500 f 109 m. Poor outcrop at Potgietersrus allows no distinction of the basal subzone from the overlying transition zone (if this exists at all), but it is clear that the basal rocks are extremely thick (4000%000 m, Van der Merwe, 1976). This thickness of the mafic sequence and the presence of high Cr/Fe chromite seams (De Villiers, 1970) are unique, and suggest that the lower portion may have Great “Dyke” affinities (Vermaak, 1976b).

605 Stratigraphy of the Potgietersrus Compartn

I

Stratigraphy of

I

Explanation

the

ti I

Fig. 9.24. Generalized stratigraphy of the Bushveld Complex (data from Cameron, 1963, 1964, 1971; Cameron and Abendroth, 1956; Molyneux, 1974; Von Gruenewaldt, 1973b, and the authors).

In the eastern Transvaal an ultramafic magma either intruded or existed previously below the marginal subzone (Liebenberg, 1964; Willemse, 1969). The existence of a discrete ultramafic magma is proved by the presence of peridotite sills in the metasedimentary floor in the eastern Transvaal (Lee, 1977) and by ultramafic satellites separated from the main Complex at Potgietersrus (Van der Merwe, 1976). These rocks cannot represent differentiation products of the main Complex and the problem is whether they are correlatives of the basal subzone or an entirely separate intrusion.

606

Transition zone. This zone is defined by the basal disappearance of cumulus plagioclase and by the appearance of persistent cumulus chromite at the top. It is 800 225 m thick in the eastern and western Bushveld while the existence of this zone in the Potgietersrus area is unproved. The thick upper and lower pyroxenites are almost pure (95-98%) cumulus bronzite, (Ens4) with traces of chromite, plagioclase and clinopyroxene, with little compositional variation of the orthopyroxene. Furthermore, Ni geochemistry of these pyroxenites shows no variations attributable to cyclicity. Sandwiched between these pyroxenites, is a thick sequence of at least six olivine-orthopyroxene cyclic units (dunites, harzburgites, olivine pyroxenites, pyroxenites). A feature is limited compositional variation of the olivine and orthopyroxene (E.N. Cameron, pers. commun., 1976). The appearance of olivine in the transition sequence, after cumulus plagioclase at the top of the lower zone, may suggest a magma addition, although it can be demonstrated (Vermaak, 1976a) that olivine appears transiently throughout the lower, transition and critical zones. This large-scale cyclicity of olivine appearance is probably related to variations of parameters such as oxygen fugacity or total pressure. Chromite is a minor cumulus constituent of the transition zone (max. 3.1 ~01.96). Thus the occurrence of two chromite seams (Cr/Fe = 2.07) near the top of the ultramafic sequence at Potgietersrus (Fig. 9.24) serves to aggravate the problem of whether the transition zone exists in that compartment (see below).

*

Critical zone. The basal appearance of persistent cumulus chromite (>3 ~01.7%) and intercumulus plagioclase, and the appearance of cumulus clinopyroxene at the base of the succeeding main zone, delimits the critical zone. In the eastern and western Bushveld, the zone is 1300 rf: 310 m thick. It has received detailed study by Cameron and his co-workers (see references). The lower subzone consists essentially of feldspathic pyroxenite (cumulus orthopyroxene) interspersed with up t o seven chromite seams (lower group Cr/Fe = 1.47-2.13) although thin olivine cumulate layers occur associated with chromite seams near the middle of the sequence (Fig. 9.24). Cyclicity within this subzone is not unequivocally proven and the recurrent appearance and disappearance of cumulus chromite presents a problem. The most meaningful mechanisms invoked in the past t o explain this phenomenon include variations in oxygen fugacity related t o periodic sedimentary ingestion (Cameron and Desborough, 1969; Ulmer, 1969) and recurrent acidification contamination of the, mafic magma by convective overturn of a granite magma “floating” on top of the mafic magma (Irvine, 1974). These appear untenable considering the positioning of chromite layers in remarkably similar silicate stratigraphies of the separated eastern and western compartments (Vermaak, 1976b). Consequently “ e x t r a n e ~ u s ~influences ~ affecting chromite precipitation are rejected in favour of inherent chemical changes such as activities of either Si, Al,or Ca in the system (Meyer, 1959;

607 Vermaak, 1976b). Latest thinking has suggested chemical modification by frequent additions of parental magma (Irvine, 1977) or in changes of total pressure caused by the physical mechanism of magma intrusion (Cameron, 1978) but these would have t o assume an in-depth linkage of the feeders t o a central magma source. The appearance of cumulus plagioclase at the base of the upper subzone marks one of the major mineralogical changes within the Bushveld sequence since all the rocks below it are melanocratic (cumulus chromite, olivine, orthopyroxene) and all the rocks above are essentially leucocratic in spite of minor layers of dark rocks at the base of some cyclic units. It is possibly regrettable that this change was afforded no recognition in the original and currently traditional subdivision (Hall, 1932). The change occurs just above (although at no fixed stratigraphic distance) the main chromite seam and commences with cyclic units in which the middle group chromite seams form basal members. There is also structural evidence of magma motion at this point which is also manifest at the base of numerous cyclic units in the critical zone. Vermaak (197613) has shown that cumulus chromite-plagioclase “mix-layering” can be ascribed to mixing of the lower cumulus minerals (essentially chromite) of a cyclic unit with the upper cumulus minerals of the preceding cyclic unit, and the subsequent winnowing into separate layers by the magma temporarily in motion. A typical example is the lowermost (UG1) chromite of the upper group so magnificently displayed at the classical Dwars River occurrence and elsewhere. The predominant cyclicity of the upper subzone is simply ratiolayering of orthopyroxene and plagioclase. Nevertheless, up t o seven chromite seams (middle group, Cr/Fe 1.33-1.52 and upper group, Cr/Fe 0.97-1.35) form the basal layers of units that invariably display a magnificent cyclicity of cumulus chromite, orthopyroxene and plagioclase. The best examples are the Merensky Reef unit (in which olivine appears transiently) and the overlying uneconomic Bastard unit (with rare chromite) whose uppermost mottled anorthosite layer (the “Giant Mottled” marker horizon) marks the top of the critical zone. Vermaak (1976b) has demonstrated that anorthosite layers are the result of plagioclase floating in the magma and these layers trapped thin portions of magma below them. Since the sulphur-carrying capacity of melts is controlled by their iron content (Haughton et al., 1974), Vermaak (1976b) has suggested that within the limited volume of the above-mentioned entrapped magma, iron-rich chromite precipitation would increase f,, and cause the separation of immiscible sulphides. These would act as collectors for the platinoid metals accumulated via the intercumulus fluid from the base of the layered sequence. These unique conditions thus explain the nature and mineralization of economic horizons such as the Merensky Reef and the UG2 chromite layer high in the critical zone. It is possible that the upfloated plagioclase mattes also entrapped volatiles below them. Subsequent

608 crystallization would increase the pressure in the entrapped volatiles which would break through the matte, such as fluid- and gas-escape structures observed in soft detrital sediments. The void left by these volatiles could probably account for the slump structures (“potholes”) observed at the top of the critical zone. A t Rustenburg and south of the Steelpoort faults the layered sequence below the Merensky Reef is different from the rest of the Complex. The sequence which would equate with the lower, transition and lower critical zones, consists of interlayered anorthosites, norites and pyroxenite compared t o a generally ultramafic character elsewhere. In both areas there is a sequence of chromite seams, with the lower seams appearing as an attenuated sequence in pyroxenite. Correlation of chrome seams and sequences north and south of the Steelpoort fault is not possible. Cameron (1971) suggests that movement along the fault during emplacement of the Complex could account for these differences. The thick sequences of leucocratic rocks interlayered with pyroxenites are problematic and are not yet accounted for. At Potgietersrus, the critical zone is reportedly 400 m thick and the economic horizon (“Platreef”) has been assigned to the main zone (Van der Merwe, 1976). Reasons for this subdivision are not clear and, compared with the subdivision elsewhere, it is probably erroneous. Thus whereas the four chromite seams in norite (Cr/Fe = 1.7) have been assigned to the critical zone, those occurring in pyroxenite below (Cr/Fe = 2.07) have been placed within the lower zone, even though the chromites of the lower critical subzone (in feldspathic pyroxenite) have a similar Cr/Fe ratio. Unless there is any valid mineralogical reason for attributing the Platreef and the overlying anorthosites to the main zone, they should be retained in the critical zone (Fig. 9.24). Main zone. The base of the zone is marked by the appearance of cumulus clinopyroxene and its top by the appearance of cumulus magnetite and apatite in the succeeding upper zone. In the eastern and western Transvaal the zone thickness averages about 3000 m. A thickness of almost 4000 m south of the Steelpoort river in the Eastern Transvaal is attributed, on slender evidence, to strike faulting (Von Gruenewald, 197313). At Potgietersrus, the main zone is about 1500 m thick. Until recently the main zone was regarded as a monotonous gabbro sequence, but ratio-graded cyclic units (plagioclase and orthopyroxene) are laterally persistent and capable of correlation over long distances. The lower subzone consists of norite t o gabbronorite interlayered with up to ten anorthosite sequences ranging from one to 50 m thick. The middle subzone has been divided into five topographically distinct gabbros (Molyneux, 1974). Where exposed, the contact between the gabbro units is characterized by leucocratic horizons up to 15 m thick, which show a variety of sedimentary features including cross-bedding, washouts and

609 inclusions. Towards the top of the middle subzone there is a sequence of ratio-layered light and dark minerals resulting in leuco and melagabbros. The base of the upper subzone is marked by a two metre thick websterite layer (the pyroxenite marker, also present at Potgietersrus) overlain by a gabbro sequence containing primary orthopyroxene, in contrast to the inverted pigeonite of the underlying sequence. Throughout the main zone, there is a general increase in iron (En7sEns8) and decrease in calcium (An78-An55)but the trend is reversed at the pyroxenite marker, which leads several workers (Von Gruenewaldt, 1970I. Molyneux, 1974; Van der Merwe, 1976) t o postulate a magma influx. In the thicker anorthosites, the An trend reverses upward in the layers, suggesting that plagioclase again floated in the magma. At Potgietersrus, the Platreef is overlain by a 300 m-thick norite succession followed by four pyroxenite bands and a gabbro sequence (Van der Merwe, 1976). The pyroxenite bands could be correlatives of the Bastard Reef situation, so that the appearance of cumulus clinopyroxene in the gabbro above them would initiate the main zone. A 110 m-thick olivine norite developed 1100 m above the Platreef is unique t o Potgietersrus. Upper zone. The base of the zone is marked by the appearance of persistent cumulus magnetite in the ferrogabbros, although magnetite is recorded in the main zone. The zone is some 1800 m thick in the eastern and western Bushveld and 1100 m thick at Potgietersrus. The upper zone is divided into three lower gabbroic subzones and an upper dioritic subzone. It is characterized by cyclicity between cumulus magnetite, olivine ( Fa49-Fa70), orthopyroxene (En,o-Enso), clinopyroxene and plagioclase (Ans9-An45). Up t o 5%apatite is present in the top subzone (fluorine-rich at the base, becoming hydroxyl-rich upwards, and finally fluorine-rich; Von Gruenewaldt, 1973b; Molyneux, 1974). A feature of the upper zone is the presence of 24 magnetite seams (20 at Potgietersrus) with a threefold grouping depending on upward increasing Ti02 (13-20%) and decreasing V20s (2.5--0.3%).Olivice occurs sporadically in the lower subzones but becomes common, and almost pure fayalite, in the upper subzone. The lower subzone repeats the cycle started in the main zone of primary orthopyroxene t o inverted pigeonite. An antipathetic relationship between cumulus clinopyroxene and olivine is related t o silica activity and the overall mineral chemistry indicates decreasing oxygen fugacity up the sequence. The sequence at Potgietersrus is similar with alternating layers of magnetite, magnetite gabbro, gabbro and anorthosite capped by fayalite diorite. The main magnetite seam contains 1.6%V 2 0 5 and 12.7%Ti02 and the uppermost seam 0.3% V20s and 21.1%Ti02. A noteworthy aspect of the upper zone is its apparent transgression across the preceding sequence in the northern position of the western compartment (Coertze and Schumann,

610 1962; see Fig. 9.22). It should be stressed, however, that this concept of a “transgression” may be an oversimplification since limited data based on recent enigmatic drilling results suggest that the geology and structure of the area are highly complex. Nevertheless such apparent transgression, together with structural, mineralogical, geochemical and isotopic evidence from other levels within the intrusion (Cameron, 1971; Von Gruenewaldt, 1973b; Molyneux, 1974; Gijbels et al., 1974) have suggested magma influxes, presumably of parental magnesium-rich magma. Initial 87Sr/86Srratios increase in a stepwise fashion upwards in the intrusion; these steps occur at horizons which show irregularities in the cryptic variation which suggests that the mafic phase was emplaced as a succession of magmatic influxes (Hamilton, 1977). In any event, a single intrusion of the enormous volume of Bushveld magma is difficult t o conceive. On the other hand, mineralogical and geochemical changes, unless supported by structural evidence, could simply mean a major convective period within the chamber before the accumulation of new crystals. It is also equally diff?cult t o visualize major parental magma influxes which were so thoroughly mixed with the magma remaining in the chamber so as not to cause a drastic alteration t o the overall fractionation path.

Pipes and plugs. Various pipes and plugs are related t o a pegmatoid phase within the layered sequence. They consist, in a descending stratigraphic order, of vermiculite pegmatoids, magnetite pegmatoids, diallagite pegmatoids, platiniferous hortonolite-dunite to dunite pegmatoids and bronzite pegmatoids (Willemse, 1964). Some of the bodies are replacement types (Cameron and Desborough, 1964) but somewhat rarer types are considered to be a product of violent fluid emplacement (Ferguson and McCarthy, 1970). The country rocks, including chromites, are converted t o iron-rich rocks by the replacing fluids. Although a structural control of the distribution of the pipes is not readily apparent, this aspect has not been studied. There is no evidence of the Bushveld having domed its roof but evidence is available from the layering (Vermaak, 1976b) that the weight of the Bushveld was accommodated by a sagging floor, causing considerable stretching of the pile. One explanation for the pipes is that failure of the rocks due to stretch could have caused downward migration of localized ironrich rest-fluids of the mafic sequence. The “bronzite pegmatoids” in the structurally disturbed area west of the Pilanesberg, are xenoliths of recrystallized country rock in the nickel pipes (Vermaak, 1976a). Although the oxides and silicates show iron-enrichment, this is unrelated to the ironenrichment of the other pipes. The nickel pipes are related t o faulting which is radial t o and the result of the emplacement of the Pilanesberg alkaline intrusion. Therefore the nickel pipes are younger and are only partly related to the Bushveld Complex in which they occur (Vermaak, 1970,1976a).

611 The Bushveld acid phase The acid rocks of the Complex comprise granophyres and granites. Spatially associated with these is the 4 km thick Rooiberg felsite succession. The stratigraphic position of these felsites is open for debate. Willemse (1954) and Hall (1932) considered them t o be the epicrustal phase of the Bushveld Complex, genetically related to the granites. Rhodes (1974) demonstrated the wide compositional scatter of the felsites and concluded that they are not related to the Bushveld granite. Button (1976) has suggested that felsites associated with the Dullstroom lavas of the Transvaal Supergroup may represent a forerunner of the Rooiberg felsite which was uplifted t o its present position to form a roof t o the mafic phase. The felsites are intruded by both the mafic sequence and the acid rocks of the Complex thus implying a time interval between them and the Bushveld granites at least sufficiently long to accommodate the emplacement of the mafic sequence. Acid phase granophyres have diverse field relations which have given rise to several hypotheses as to their origin. The granophyres occur as a finegrained facies in the upper part of the Bushveld granite, and at the contact between the mafic sequence and the felsites where the granophyre formed as a consequence of partial melting of the felsite (Von Gruenewaldt, 1968). Other granophyres have formed through melting and recrystallization of Bushveld granite adjacent t o mafic sills. Mineralogically the granophyres consist of quartz-alkali feldspar intergrowths with textures ranging from coarse granophyric t o micrographic. The granophyre occurring as a roof phase of the Bushveld granite and regarded as a product of magmatic crystallization has quartz and alkali feldspar phenocrysts surrounded by fine-grained granophyre intergrowths (Rhodes, 1975a). Granophyre considered t o have formed from melted felsites has quartz and potash feldspar intergrowths set with zoned (An, + An4) plagioclase phenocrysts (Von Gruenewaldt, 1968). Generally there is little variation in the granophyre mineralogy, although there is a wide scatter of the Na20/K20 ratio. The dominant rock of the acid phase is the stratiform 2.8 km-thick Bushveld granite which intrudes the felsite roof, the Transvaal sediments and the mafic sequence. The granite is composed of potash feldspar, oligoclase, finegrained quartz, with accessory biotite and hornblende. In the eastern Bushveld Lenthall (1975) has described an increase in the amount of mafic mineral in a zone following the contact of the granite with the mafic phase. Layering in the granite is defined by variations in colour, grain size and mafic mineral content; a number of different textural varieties have been described. Stocks of red, medium- t o coarse-grained granite composed of quartz and alkali feldspar with variable amounts of chlorite, minor biotite and accessory fluorite and carbonate cut across the Bushveld granite. Economic deposits of cassiterite occur as disseminations in the younger granite or in zoned pipes in the upper parts of the younger granite stocks, the BushveId granite and granophyre.

612 Rhodes (1974) has demonstrated that analyses of the Bushveld granite lie close to the thermal minimum in the Qz-Ab-Or system and that generally the composition is chemically homogeneous,although there is some variation in Ca content. McCarthy and Hasty (1976) present a model of Bushveld granite crystallization, in which they show that distribution of Ba, Rb and Sr in the coarser-grained granites define a single continuous trend. They conclude that the granites formed by in situ fractional crystallization of a single magma. They calculate that the coarse red granite (younger granite stocks) formed late in the fractionation sequence; this granite is depleted in Ba and Sr, enriched in Rb and contains economic tin deposits. The initial 87Sr/86Sr ratios for the Bushveld granites range between 0.7153 and 0.7240 (Davies et al., 1970; Hamilton, 1977) from which an origin by partial melting of the crust is argued for the granites. The age is given at 1940 Ma, with which U/Pb studies on zircon (Burger et al., 1967) and Rb/Sr ages for mica (Nicolaysen et al., 1958) are in agreement. Kindred complexes (a) The Trompsberg Complex, with a diameter of some 27 km, was intersected by drilling beneath 1000 m of Palaeozoic cover (Ortlepp, 1959). It intruded marbles correlated with the Transvaal sequence. The rocks are predominantly coarse leucogabbros and anorthosites, melagabbros, olivine gabbros, troctolites and titaniferous magnetites. Although a feature of the mineralogy is the coexistence of titaniferous magnetite with forsteritic olivine, the sequence resembles the upper zone of the Bushveld complex. Magnetic surveys suggest that the mafic rocks are concentrically disposed around an overlying central granophyric granite. The age of the acid rocks (Davies et al., 1970) is 1372 k 142 Ma with an initial 87Sr/86Srratio of 0.7066. The mafic rocks show an initial ratio of 0.7043. (b) The Losberg Complex is a sheet-like intrusion, 130 m thick, intrusive into the Pretoria Group of the Transvaal Supergroup. The sheet, differentiated from basal harzburgites to upper augite granophyre, shows the normal sodium-enrichment of the plagioclase and iron enrichment of the ferromagnesian minerals. The regular trends of trace elements across the phase layering suggests the in situ differentiation of a single magma (Abbott and Ferguson, 1965; Danchin and Ferguson, 1970). (c) The Kuffirshruul Pyroxenite, is described by Frick (1975) as a funnelshaped intrusion with a steeply-dipping porphyritic norite marginal zone. The main rock-type is a clinopyroxenite which dips towards the centre at 4", the top of the succession being capped by an 8 m-thick magnetite seam. Witwatersrand quartzite and shale xenoliths occur along the outer margins while norite or gabbro xenoliths are randomly distributed. Frick regards the mass as a zoned ultramafic complex originating from a multiple intrusion of magmas differentiated in depth. The first magma yielded gabbros and norite and the second more mafic magma resulted in magnetite-rich rocks, disrupting the first intrusion.

613

614 (d) Other related occurrences, of which little is known, are the Vogelstruisfontein pipe-like mass of mafic rocks within the Witwatersrand Supergroup, and the Koringkoppies chromiferous pyroxenite (Fig. 9.25). The Uitloop layered pyroxenite is regarded as a satellite body (of which there are others) of the Potgietersrus compartment. Tectonic setting The emplacement of the Bushveld Complex and kindred layered intrusions in the Kaapvaal and Rhodesian cratons should be related t o the geological development and tectonic framework of Southern Africa (Fig. 9.25). The Kaapvaal and Rhodesian cratons appear t o have deveIoped increasing stability and crustal thickness under tensional conditions with dominant vertical tectonics, which contrasts with the compressional tectonic stacking and thrustcontrolled interleaving proposed, as in Greenland, t o account for crustal thickening and increasing stability of the early cratons (Bridgwater et al., 1974). The great volume of plutonic and volcanic rocks in the stratigraphic record of the Kaapvaal/Rhodesian craton (of which the Bushveld and related complexes are a manifestation) is the result of this pervasive extension. We would also concur with Hunter ( 1 9 7 3 ~that ) the emplacement of these complexes should be regarded as the culmination of a successive evolution of geological events. These will be considered in terms of chelogenic cycles interspersed by disruptive cycles (Sutton, 1967, 1973). The pre-Bushueld history The initial differentiation and the first chelogenic cycle occurred between 4500 and 3800 Ma ago during which time primitive sialic crust developed by bimodal dacitic-tholeiitic volcanism (Barker and Peterman, 1974) as small, distinct, nuclei co-existing with a hydrous simatic lithosphere. The Kaapvaal and Rhodesian cratons are thought to have formed a single cratonic nucleus and available evidence, especially from Zimbabwe (Bliss, 1962; Stowe, 1968; Wilson, 1968a, 1973), suggests that the original tectonic axes had a northerly strike. This structural grain could be related t o the “edge effect” (Ramberg, 1967) of the oblong shape (longer axis striking between N15” and N25”) of the composite cratonic nucleus. As the cratons stabilized, dyke intrusion occurred and shelf-type deposition of a world-wide correlative sedimentary type (quartzites, banded ironstone, carbonates and aluminous pelitic sediments) accumulated as a primordial supracrustal sequence in localised shallow basins. The basement and cover sequence were then intruded by characteristic anorthosites dismembered from extremely fractionated layered intrusives emplaced after the sedimentation. A long disruptive cycle, extending from 3800-3200 Ma led initially t o the development of a whaleback-type crustal swell (the northeasterly aligned Limpopo Belt) bounded by parallel faults causing graben preservation of the central zone which retains the characteristic sedimentation (Messina Forma-

615 tion), anorthositic intrusions and structure of the preceding cycle. The effect of tectonism was t o separate the original composite nucleus into the Kaapvaal and Rhodesian cratons while at the same time overprinting a strong ENE anisotropism for great distances on t o both cratons. Although this imprint extended 200 m southward into the eastern part of the Kaapvaal craton, the western part appears to have been protected from the overprint by the NWtrending Malopo swell, so that the early northerly grain (i.e. Kraaipan trend, Fig. 9.25) of the original tectonic grain has been preserved. The second event during this disruptive cycle was the development of greenstone sequences, during which large volumes of ultramafic, mafic and felsic lavas were extruded and eventually covered by the sediments tenninating the event. It is probable that the tensional conditions originated by the Limpopo swell (and other belts surrounding the cratons) favoured the generation of large quantities of magma and that separation occurred along the pre-existing northerly and northeasterly tectonic grain. Periodic reactivation of the mobile belt caused the belt to spread and encroach on to the cratons and rework the greenstone material. Granitic emplacement between 3200 and 3000 Ma terminated the disruptive cycle and initiated the second chelogenic cycle during which the Kaapvaal craton developed a measure of crustal stability. The sequential accumulation of thick sedimentary sequences is a measure of this stability although some instability can also be recognised in the vertical tectonics attributed to intermittently rising granite domes flanking the early sedimentary basins. This suggests a mobile substratum which caused the rise of granite domes as a result of crustal loading. Sedimentation in the successive epicontinental basins of the Kaapvaal craton was controlled by : (1)Northern provenance areas and an alignment determined by pre-existing tectonic grain. (2) A constant cyclic pattern of sedimentation and volcanism (Pretorius, 1966), a sequential northwestward migration of successive basin axes (Annhaeusser, 1973) and a decreasing energy level with time as reflected by the sediments (Hunter, 1974a,b). The interplay of existing structural trends led to the basins occupying synclinal troughs interspersed with anticlinal ridges along which active granite domes were aligned. In the light of Ramberg’s models (1967), it is probable that the granite doming initiated the pattern of troughs and ridges and thereafter, especially during the early sedimentation, played an important role in the vertical tectonics which determined the configuration of the basins and the characters of the sedimentation within them. The sedimentation immediately preceding the Bushveld event was that of the Transvaal Supergroup consisting of a localized lower clastic-volcanic group, a middle chemical sedimentary group and an upper pelitic-clasticvolcanic (Pretoria) group which constitutes the “floor” or “wall” to the

616 Bushveld Complex. The Pretoria Group is extensively invaded by diabase and dolerite sills, the latter being intruded (with strong differentiation) near the base of the Bushveld Complex, although their true relationship to the Complex remains obscure (Willemse, 1959, 1964; Frick, 1973; Hunter, 1974). A tilloid near the base of the Pretoria Group provides valuable environmental implications regarding the position of the Kaapvaal/Rhodesian cratons in pre-Bushveld times (see below). This unique period of crustal stability (referred to by Cloud, 1976 as the Zuluan Wedge) was terminated by the Bushveld disruptive episode which deserves our closer scrutiny below. What we have attempted to stress in the foregoing is the great importance of early structure and tectonics within the tensional environment of the Kaapvaal/Rhodesian cratons which have been constantly responsible for determining the patterns of succeeding events leading up to the Bushveld disruptive cycle.

The Bushveld disruptive cycle (2500-1300 M a ) Several hypotheses have been presented to account for the emplacement of the Bushveld Complex, namely : (a) A response to the central collapse of a vast geanticlinal structure produced by tangential forces (Daly, 1928). (b) Alignment along a major lineament that is most prominently marked by the Great Dyke (Hall, 1932; Brock, 1956,1957; Cousins, 1959). (c) A response, in the cratonic environment, t o deformation in the adjacent Limpopo mobile belt (Crockett, 1969). (d) Emplacement of mafic magma along an east-northeast axis as a result of tension in an orogenic environment (Truter, 1955; Visser, 1957). (e) The product of extra-terrestrial impact (Dietz, 1963; Hamilton, 1970; Rhodes, 1975b). ( f ) A response to superposition of north-northwest and east-northeast crustal warps coincident in time with a period of maximum continental inundation and stability, the warping being associated with the vertical movement of crustal blocks (Hunter, 1975). The Bushveld Complex is one of a number of mafic layered intrusives that are developed in the Kaapvaal/Rhodesian craton, namely : NORTH : Great “Dyke” 2514 f 16 Ma. (Hamilton, 1977) Bushveld Complex 2095 f 24 Ma. (Hamilton, 1977) 1881 2 282 Ma. (Davies et al., (1970) Losberg Complex SOUTH: Trompsburg Complex 1372 f 142 Ma. (Davies et al., 1970) In addition t o the above, a significant positive gravity anomaly (the Northern Cape anomaly, Smit et al., 1962), in the extreme western part of the Kaapvaal craton which shows alignment parallel to the Kraaipan trend, may mark the site of a mafic intrusive not unroofed by erosion.

617 If the layered complexes listed are t o be related t o one disruptive episode of intrusion, then the following would have to be explained: (1)The significant difference in the morphology of individual complexes; (2) Their alignment along a NNE-trending lineament; and (3) The intermittency of intrusion as marked by their significant southward younging. The shape of a layered intrusion would appear t o be linked to the crustal level at which it was intruded and eventually exposed. It can be assumed that the lowest crustal level would be dominated by the oldest tectonic trend and that of the highest level by the youngest structures. Consequently the Great “Dyke” and the central portion of the Potgietersrus compartment which are flanked by basement, assume a dyke-like linear configuration parallel t o the oldest northerly structural grain of their environment. By contrast, the overprinted structures i.e. those of the Limpopo trend and the Malopo arch in the supracrustal sediments of the Kaapvaal craton respectively determine the structural sill-like configuration of the Bushveld Complex and its western limit. It is also significant that the northern Cape anomaly across the Malopo arch assumes an elongation parallel to the Kraaipan trend. A good example of the effect of the presence of a supracrustal sequence is seen in the Potgietersrus compartment where the complex has a dyke-like attitude in the basement granite but is sheeted where it intrudes the Transvaal Supergroup. The tectonic development of the Kaapvaal/Rhodesian cratons and the emplacement of the Bushveld and Kindred intrusives can be considered in terms of the movement of lithospheric plates in north-south directions across the oblate spheroid of the earth (Oxburgh and Turcotte, 1974). Two models of movement and the membrane stressed involved are: (a) Motion of plates towards the poles induces compressional stresses in the centre of the plate and tensional (tangential) stresses on the periphery of the plate. (b) For a plate moving from the pole to the equatorial regions, the interior of the plate is in a tensional condition and the peripheral stresses are compressional. The first membrane tectonic model is used t o explain continental break-up, limited horizontal extension (e.g., Red Sea trench) or very limited finite extension systems without evidence of either continental split or disruption (e.g., East African Rift). The model accounts for the southward younging of Mesozoic and later rifting and volcanism. The second model provides a possible solution to the interrelated Bushveld episodes. The presence of a tilloid in the Pretoria Group suggests that the cratonic pair may have been located around the (South) pole just before the Bushveld event. If the Kaapvaal/Rhodesian craton pair moved northwards towards the equator, the initial peripheral stresses would be compressional and the centre of the Rhodesian craton under tension, favouring proto-

618 rifting and permitting intrusion of the Great “Dyke” as a linear complex determined by the oldest structural grain in the crust. Tensional fracturing would cause stress-release magma generation within the mantle. No intrusion of Great “Dyke” magma would occur in the Limpopo mobile belt as it .would be in a state of compression. With continued northward movement of the cratonic pair, tensional conditions would develop in the Kaapvaal craton, permitting the emplacement of the Bushveld and related magmas along the fundamental north-trending tectonic grain of the craton, although their morphologies will be determined by the supracrustal sediments into which they were intruded. According to the scheme outlined here a progressive southward activation of the tectonic grain should be reflected by a southward younging of the mafic complexes. The near coincidence of the Great “Dyke” age and that of the earlier metamorphic event in the peripheral Limpopo mobile belt (2550-2600 Ma) could reflect northerly movement of the Rhodesian craton. Similarly the 2000 Ma Limpopo event is close to the age of the Bushveld Complex possibly reflecting the later northward movement of the Kaapvaal craton. The post-Bushveld Waterberg sediments have been tectonically affected along an ENE direction probably caused by failure along the greenstone belts, but greater tectonic disruption is shown along a broad belt parallel to the Great “Dyke”-Bushveld Complex-Trompsburg Complex lineament (Jansen, 1975), providing additional evidence for the existence of an ancient potentially active mid-cratonic rift. D. Sedimentary sequences of uncertain correlation THE PIRIWIRI, DEWERAS AND LOMAGUNDI GROUPS

*

An arcuate belt of folded sedimentary rocks of contrasting bulk lithologies occurs in the northwest of Zambabwe, stretching in a NNE direction from the Umniati River, west of Gatooma, to the Zambezian Escarpment. These rocks were originally grouped together in a loosely defined Lomagundi System but are now subdivided into a Deweras Group in the east comprising arkoses, greywackes, argillites and basaltic lavas, unconformably overlain by a Lomagundi Group consisting of quartzites, slates and greywackes with a basal orthoquartzite-carbonate association and adjoined on the west by an extensive Piriwiri Group composed mainly of phyllites and greywackes (Fig. 9.26). The lithological groupings of these rocks have been defined several times during the piecemeal mapping of the belt, and the Piriwiri and Deweras have been referred to both as separate entities and as series and formations of the

* By. J.G. Stagman.

619 YOUNGER COVER SIJARIRA GROUP

a

PARAGNEISSES OF POST BASEMENT AGES

7URUNGWE GRANITE LOMAGUNDI GROUP 3 DEWERAS GROUP PlRlWlRl GROUP

Fig. 9.26. Distribution of the Piriwiri, Deweras and Lomagundi groups in Zimbabwe.

original embracing Lomagundi System. Accounts of the chequered history of the classifications which evolved are contained in the many Zimbabwean Geological Survey publications on the subject (see references). Recent completion of regional mapping of the belt, which has involved the tracing of the lithological and structural units, together with the adoption of a loose lithostratigraphic system of nomenclature has resulted in the establishment of three separate groups. They are composed of variously named formations which possess certain distinctive or combinations of distinctive features, possibly of a genetic character, and are subdivided into members themselves of distinct lithological character. The exact inter-relationships of the groups are imperfectly known for throughout most of the belt normal stratigraphic relationships are obscured by folding and thrusting, and there is an almost total absence of exposure in the critical areas. The uncertainty which has been termed the "Lomagundi Puzzle" is compounded by high-grade metamorphism in the north and a cover of Lower Karoo and eluvial deposits in the south. The Deweras Group is regarded as having been deposited in a graben formed by fracture of the eastern margin of the Piriwiri basin whereafter further subsidence and marine incursion led t o the differentiated sedimentation of the Lomagundi Group in a series of axially adjoining basins along the same general northerly trend.

620

South of latitude 17"s the groups display no evidence of having been affected by the -500 Ma Miami metamorphism. Nevertheless the radiometric age determinations have been inconclusive in attributing absolute or even relative ages t o the three groups. Galena from the Copper Queen Mine in the Piriwiri Group yielded an age of 2250 5 32 Ma and a lead age of 1980 f 33 Ma from a prospect in Lomagundi dolomite is recorded. Wholerock, K-Ar ages from Lomagundi slates and Piriwiri phyllites range between 1655 Ma and 1975 Ma. Furthermore the Urungwe porphyritic granodiorite intrusive into the Piriwiri with the production of a thermal aureole, has a whole-rock, Rb-Sr age of 21805 130 Ma whereas its biotite age is only 833 ? 40 Ma. Numerous explanations for these diversities have been offered but the picture remains obscure other than t o say that the groups are of Early Proterozoic age and were only mildly metamorphosed where not involved in the Miami (-500 Ma) event, which is described in the chapter dealing with the Zambezi Metamorphic Belt. Piriwiri Series was the name given by Molyneux in 1919 t o a monotonous succession of phyllitic rocks forming a belt 50 km wide in the country west of Sinoia. In the west they pass unconformably under Karoo or Sijarira red beds cover or have been intruded by the Urungwe granodiorites. Beyond this granitic core they are recognised as remnants, at higher metamorphic grade, infolded with various older paragneiss formations in the Zambezi Metamorphic Belt; here with northeasterly rather than NNE strike and extending as far west as the Wankie District where they occur as the tin-bearing schist belts of the Kamativi-Dett inlier. Although Molyneux recorded younging from west t o east of the Kanyaga, Chidomo Sandstone and Graphitic divisions which he distinguished in his series and which are throughout in faulted contact with the west-dipping Striped Slates of the Lomagundi System, he nevertheless placed them as the uppermost members of that system. This practice has been followed by many subsequent workers in the field and the Piriwiri has been described as the deep water or flysch facies coeval with a miogeosynclinal or shelf facies assemblage of dolomites, orthoquartzites and slates to the east. Essentially the Piriwiri Group, as it is now named, consists of phyllites, greywackes, black graphitic and ferruginous slates, cherty quartzites, other subordinate argillaceous and arenaceous members plus rare volcanics. In areas of higher metamorphic grade calc-silicate rocks are present but their parent type has not been recognised elsewhere. Structural analysis around and east of the Copper Queen Mine led t o the distinguishing of four phases of deformation. These and the accompanying metamorphism were termed, perhaps rather extravagantly, the Magondi orogeny. The practical result was folding of the group principally on NNE axes curving north and eventually NNW and the impression of a strong, usually vertical, regional foliation often coincident with the bedding which has led to the sculpting of strikingly linear terrane unmistakable on aerial photographs and contrasting strongly

621

with the photoexpression of the Deweras and Lomagundi groups. In the thermal aureole of the Urungwe granite, Piriwiri rocks have been converted into garnetiferous mica schists and micaceous quartzites and in the Miama area, by the much later 500 Ma metamorphism, progressively to sillimanite schists and gneisses, but for the greater part low greenschist facies prevails. Many attempts have been made at stratigraphic subdivision, often based on the proportion of various constituent varieties and the presence or absence of impersistent volcanic members. For instance along the Umfuli River section (southeast of the Copper Queen Mine), three divisions have been erected on somewhat tenuous grounds. An Umfuli Formation comprises phyllites and greywackes with minor interbedded cherty quartzites and grits plus, in the south, graphitic and pyritiferous slates at the base. The Chenjiri Formation, also consisting of phyllites and greywackes, is characterized by thin beds of tuff, chert and pyritiferous slate and thicker beds of micaceous quartzite are prominent. Thirdly, the Copper Queen Formation has predominant phyllites identical t o those of the other formations but lacks interbedded chert and tuff. Additionally i t contains the contact aureole of the Copper Queen and King domes where hornfelses, skarns and amphibolites, probably of Basement Complex origin, are circumscribed by garnetiferous mica schists and micaceous quartzites. Due to thinning out of members and subtle lateral facies changes, such subdivisions can seldom be sustained beyond local map boundaries. As a consequence in the most recent analysis it has been necessary to revert to a two-fold division namely a Kanyaga Formation equating largely t o Molyneux’ division of the same name and composed of phyllites, greywackes and micaceous quartzites with a very fine, silty, buff t o pinkish soil cover and a Chitena Formation with prominent graphitic phyllites, cherts and black ferruginous slates. Both are penetrated by sodic diatremes and contain thin beds and lenses of agglomerate, tuff and associated narrow, cherty albitite dykes. As previously stated, opinion regarding the relationship between the Piriwiri and Lomagundi groups has been contentious. For more than 220 km of strike the exact contact is nowhere displayed. Its approximate position is usually marked by a topographic break and the appearance of graphitic schists. However, there are numerous factors which, although not individually conclusive, strongly indicate that Lomagundi deposition postdated the formation of the Piriwiri Group. Among these are different folding styles; the almost ubiquitous presence of concordant white quartz veins in the Piriwiri absent from the adjacent Lomagundi striped slates; the generally higher metamorphic grade of the former; and the occurrence within the Piriwiri Group only of Basement Complex type of hydrothermal gold mineralization, as at D Troop Mine, copper-lead-inc mineralization at the Copper Queen Mine and low-grade copper-gold mineralization in association with quartzsiderite veins that characterizes the so-named old Piriwiri mineral belt.

-

622

In addition to those described there are three other categories of mineralization within the group. In the zones of higher metamorphic grade intrusions of pegmatite are common and many have been mined for their tin, tantalum and beryl content. Late phase quartz veins have produced wolframite and several of the pegmatites contain beryl, topaz and tourmaline of gem quality formed as the result of late pneumatolytic processes. High-grade book or sheet mica once the basis of a flourishing industry has deteriorated into scrap recovery for ground mica production. The mica-bearing pegmatites are demonstrably confined t o sillimanite grade terrane which also contains deposits of kyanite, sillimanite and graphite ore. Manganese often in associztion with flatdipping quartz veins within graphitic phyllites is characteristic for long stretches of the contact zone between the Piriwiri and Lomagundi groups where thrusting is suspected. The Deweras Group is seen as a deposit of medium- to coarse-grained unworked sediments of a continental environment. Alluvial fans situated along the fault-bounded margins of crystalline massifs are considered t o have distributed the sediments into intermontane troughs with finer-grained and sometimes calcareous varieties being the product of evaporative conditions in temporary basins of water along the axes of the troughs. The bounding faults, as well as controlling uplift of the provenance, probably acted as channels up which tholeiitic magma was introduced. The group was, in 1936, first defined in the Lower Umfuli area t o the northwest of Hartley, where it was regarded as an uppermost series of the Basement Complex although separated by major unconformity from all older formations. Traced southwards beyond the Umniati River into the Mafungabusi area, it became apparent that it had a much greater affinity with the overlying Lomagundi Group albeit here too with distinct unconformity . Proceeding north the Deweras passes underneath the Lomagundi for some 40 km to emerge as an antiformal exposure 10 km wide flanked by hills of Lomagundi dolomite and quartzite which extend nearly to the Mangula Mine. North beyond the mine there is a swing in strike from north t o NNW and the arkosic rocks, heavily intruded by dolerite, continue as a belt some 5 km wide to the Zambezi Escarpment. They rest on granite, in the main unconformably but with intrusive relationship adjacent to Mangula Mine and dip west below the Lomagundi dolomite, quartzite and slate which are succeeded by the Piriwiri Group, in this area all apparently conformably. South of the Umfuli River exposure is relatively good and although the rocks have been involved in two major episodes of deformation accompanied by thrusting and extensive faulting, bedding dips are flat t o moderate and it has been possible to divide the sequence with confidence into three formations. A Lower Arenaceous Formation, sporadically developed, attains a maximum thickness of 150 m. Grey-green t o reddish brown grits, arkoses and thin shales are subordinate t o a distinctive basal conglomerate with Base-

623 ment Complexderived clasts of tonalite, porphyry, quartz, chert and lavas. This is unconformably overlain by the Volcanic Formation as much as 1000 m thick in places and composed of flows of amygdaloidal tholeiitic lava generally green or purplish in colour and often strongly epidotized. The flows, which provide evidence of gravity differentiation, are often separated from one another by thin impersistent mud seams. In scattered localities agglomerates occur at the top of this succession. The very thick Upper Arenaceous Formation rests with a fossil lateritic base upon reddened and spheroidally weathered lava. It is composed of greywackes of variable lithology and containing a chaotic boulder bed, followed by an argillite member with reddish shales and thin dolomite horizons; in turn capped by considerable thicknesses of medium- to coarse-grained, cross-bedded arkoses. The formation is grey-brown in the lower parts becoming paler and pink in the upper parts. Although there is no reason to believe that conditions of sedimentation were much different in the section northwards of the latitude of Sinoia, no similar division of the group into formations and members is possible. For the most part metamorphic grade remains low but it is suspected that deformation, principally on north t o NNW-striking axes, is very much more intense. The terrane is little dissected and natural outcrops are rare. Except in the vicinity of some of the copper mines, where silicification has rendered the rocks more resistant to weathering, there are only small outcrops often several kilometres apart and separated by fine, pale sandy soil. On the basis of these few exposures the group is believed to consist of a great thickness of pink arkoses representing an ill-defined Upper Arenaceous Formation with sporadically developed boulder and pebble beds. In the mine workings shales, argillites, chloritic schists and carbonated chloritic arkoses feature as well but presumably are not representative. The Volcanic Formation is known only from diamond drill core and small inliers of sheared amygdaloids with reddish clay soil cover in the environs of Shackleton Mine. Other than in the Silverside Mine outlier of possibly overthrust, sheared chloritic lavas underlying equally strongly sheared, arenaceous, feldspathic sediments, at the northern extremity of the Hunyani Escarpment (north of Sinoia), it has not been recognized further north. Indeed it appears that north of Shackleton Mine sills of dolerite, representing a deeper erosional level, replace the Volcanic Formation. Dolerite becomes more prominent within the Group proceeding northwards until some 20 km north of Mangula where, on entering the zone of higher metamorphic grade, it becomes amphibolitized and constitutes nearly half of the area of outcrop. The arkosic rocks of this zone are paler, harder and finer grained than usual and nearing the Zambezi Escarpment have been converted into feldspathic paragneisses. The arkoses of this ill-defined Upper Arenaceous Formation are the hostrocks of the important copper mineralization which characterizes the group. It has been described variously as syngenetic, hydrothermal and strata-

624

bound. It exhibits elements of each. If of biogenic origin, a popular modem concept, it presumably moved from a shaly or calcareous depositional environment into the more sandy one in which it is found and there became influenced by some thermal agency of concentration which led t o crosscutting relationships, alkali metasomatism and injection of pegmatite veinlets. Deposits currently being exploited are at the Mangula, Avondale, Shackleton and Angwa mines. The Silverside orebody is a hydrothermal quartz-carbonate lode. Exploration in the very much better exposed Upper Arenaceous Formation of the southem area, between the Umfuli and Umniati Rivers, has not disclosed orebodies of any consequence. Apart from copper with by-product silver and gold, uranium of academic interest at Mangula Mine and several small quartz veins that have been exploited for gold, the Deweras Group is apparently devoid of metallic mineralization. Outcrops of the Lomugundi Group extend from the Umniati River in the south to the Zambezi Escarpment 250 km away. For the entire distance its quartzites, dolomites and slates are in close association with arkosic rocks and in the south with basaltic lavas as well, of the Deweras Group. In the Nyamandhlovu District, 180 km southwest of the most southerly outcrop of the main belt, a borehole drilled through Kalahari Sand and Karoo cover passed, at a depth of 425 m, into a succession of dolomites, quartzitic dolomites and dolomitic schists with quartz-sericite and graphitic schist intercalations more than 400 m thick. Being on the direct strike extension and on the basis of characteristic carbon isotope ratios there is little doubt that this occurrence represents an isolated basin Lomagundi sedimentation. On account of the close parallelism of outcrop the Lomagundi and Deweras have in past classifications frequently been grouped together despite strongly contrasting lithologies and not taking fully into consideration the major unconformity between them which is clearly displayed in the Umniati River area although, due t o more intense folding and thrusting, this becomes less readily discernible as one proceeds north along strike. Between the Umniati and Umfuli Rivers division of the group into three formations has been effected, but the uppermost or Sakurgwe Formation dies out at the latter river and has not been recognized elsewhere; so that generally only two divisions persist as is illustrated in Table 9.XI. There is a conglomerate or coarse grit basal to the Arenaceous division both in the northern and southern areas. In the latter it is seldom deformed and is characterized by jaspilite and vein quartz clasts often in a dolomitic matrix. In the former, along the Hunyani Escarpment the basal bed is usually screecovered but where exposed is always a strongly deformed schistose grit with appreciable sericite and chlorite content, completely drawn out rock clasts and rare rounded quartz pebbles. It is mylonitic in character supporting the view that the Lomagundi has been thrust over the Deweras or Basement Complex along much of their contact. The quartzites of the formation are generally pure and fine-grained. At

625 TABLE 9.XI Sinoia Area northwards

Magondi-Mafungabusi area

Argillaceous division Sakurgwe Formation Greywacke + 600 m Striped Slate Formation striped slates mountain sandstone + 500 black slates

I I

Arenaceous division Sinoia Caves Formation dolomite + 300 m quartzites, with pockmarked variety + 600 m and thin dolomite schistose basal conglomerate 5 30 m

Nyagari Formation argillites striped slates + 600 m

Mcheka Formation main dolomite, 0-400 m upper quartzite, flOOO m pocked quartzite, 0-200 m lower quartzite, 0-300 m basal dolomite, 0-300 m basal conglomerate, 0-75 m

surface they are sugary and white in colour where not locally iron-stained. Fresh exposures in cuttings and quarries are hard, almost glassy and white, grey or pinkish in colour. These quartzites are thickly developed t o form the imposing Mcheka-wa-ka-sungabetarange (south-southwest of Sinoia) and the Hunyani Escarpment. A pock-marked variety of unusual appearance, believed to afford a bedding marker may not be such as it apparently transgresses t o near the base of the formation in the extreme southwest. Additionally to dolomites, thin grits and sandy shales are interbedded with the quartzites but they, like the dolomites, are seldom exposed except on steep slopes. All along the eastern edge of the Lomagundi belt bedding dips average between 30" and 40" t o the west with small local reversals and practically all of the rocks, the argillaceous ones particularly, have a near vertical, schistose foliation trending parallel to bedding strike. Stromatolitic and calc-arenite structures in the dolomite, suggesting a shallow-water evaporitic environment, are well-preserved in the country southwards of Sinoia. These features are less apparent elsewhere. The calcareous rocks are very fine-grained, white, pink, grey or buff in colour with darker weathered skin, the roughness of which increases with degree of silicification, brecciation and cherty interlamination. In the Sinoia Caves area a 10 m-thick bed of pure, pale pink dolomite occurs near the base of the succession and an upper band 300 m thick containing a bed of coarse, grey argillaceous grit is at the top, dipping directly beneath the Striped Slate Formation. Due to lack of exposure it is seldom that such bands may be traced more than a few kilometres along strike.

626

This expression of a little contorted, relatively regular sedimentary sequence is not repeated along the anticlinal limbs of the Arenaceous Formation which arch over the Deweras Group of the Umboe Valley between the Alaska and Norah mines and in the country to the north. The quartzites, dolomites, minor grits and shales have undergone such severe tectonic distortion, coupled with brecciation and silicification, that they are inextricably intermingled and so attenuated that the outcrop width of the whole formation is often reduced t o less than 500 m. The west-dipping belt from Mangula to beyond Shamrocke Mine is similarly attenuated and north of the mine the sediments are largely replaced by interstratified dolerite now converted into amphibolite. The dolomites in this area have been changed into coarse-grained marble and quartz-tremolite rock, the argillaceous beds into garnetiferous biotite schists and the quartzites coarsely recrystallized. The only other sector where metamorphic grade of the Lomagundi rocks rises above low greenschist facies is some 15 km south of Mangula where latestage folding has depressed the formation into an oblique synclinal belt with northeasterly axis and in which the dolomite has been thermally metamorphosed t o massive tremolite rock by a large underlying mafic intrusion. The Argillaceous division, usually referred to as the Striped Slates, comprises the Nyagari and Sakurgwe formations in the southern third of the area, the Zhonzi Formation in the centre and again the Nyagari Formation in the far north. In the south this last provides evidence of unconformable relationships with Mcheka dolomites and quartzites. It consists there of striped slates, argillites, minor beds of andesitic lava, tuff and agglomerate and occasional lenses of dark, carbonaceous dolomite. The equivalent Zhonzi Formation has been divided into black basal shales overlain by striped slates. The prominent but lenticular Mountain Sandstone, here an impure grey quartzite, is interbedded near the top of the black shales. The Sakurgwe Formation, a monotonous sequence of fine-grained, unbedded greywacke and minor argillite is known only from the large synclinorium, of which it forms the uppermost unit, centred on the river of the same name. Mineralization of the Lomagundi Group is lean compared with that of the underlying Piriwiri and Deweras groups. The only known primary occurrence of importance is the copper deposit at the Shamrocke Mine which lies within rocks occupying the stratigraphic position of the Mcheka or Sinoia Caves formations. In this area of high-grade metamorphism the equivalents of the dolomite and quartzite are overlain by calcareous graphitic schists which contain complexly folded beds of meta-arkose, a lithological sequence differing considerably from the norm. Copper as chalcopyrite is disseminated in the meta-arkose. Other copper deposits, of secondary nature, which have been mined in severely faulted zones within the Lomagundi dolomite in the Alaska Mine area were almost certainly derived from adjacent Deweras arkose. Meagre lead mineralization and small hydrothermal gold veins are known but do not warrant exploitation.

627 However, the dolomites are good aquifers and locally provide sufficient water supplies for irrigation. In addition, though frequently siliceous, they contain unlimited amounts of pure material suitable for aggregate, for building and ornamental purposes and are a source of ground and burnt lime for agricultural use. Within the Striped Slates there are several localities where the coincidence of bedding and metamorphic foliation have resulted in the formation of slate which is quarried for many building purposes too but does not cleave finely enough to be of roofing quality. Volcanic centres. Several assemblages of highly carbonated vent agglomerate with subordinate crystal tuffs, fine-grained bedded tuffs and pyritiferous felsite are irregularly located in a northeast-striking zone over a distance of some 60 km and with one exception within phyllites of the Piriwiri Group. The exception is the most northerly or Nyamakari Centre, due south of D Troop Mine, where the country rocks appear t o be Lomagundi slates but where field relations are obscure. Each occurrence occupies an area of one to two square kilometres. In some instances the tuffs and felsites are intercalated with the phyllites but in others the agglomerates, containing fragments of tuff, felsite and other country rocks in addition t o obviously foreign blocks, are apparently intrusive and post-date the deformation of the phyllites. Chemical analyses of the felsites, ultra fine-grained rocks of cherty aspect, show them t o be almost pure albitites and the fragmental rocks as well are highly sodic and contain large proportions of magnesium carbonate. Although never directly associated, it is clear that the hydrothermal veins of quartz with iron-magnesium carbonate of the area, which contain low-aade copper and gold mineralization, are genetically related to the volcanic activity. REFERENCES

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629 Haughton, S.H. (Editor), 1964. The Geology of Some Ore Deposits in Southern Africa. Volume I: Gold Deposits of the Witwatersrand Basin. Geol. Soc. S. Afr., Johannesburg, 616 pp. Hobday, D.K. and von Brunn, V., 1976. Evidence of anomalously large tidal ranges in the early Precambrian Pongola Supergroup. S. Afr. J. Sci., 72: 182. Humphrey, W.A. and Krige, L.J., 1932. Explanation of Sheet 102 (Vryheid). Dept. Mines, S. Afr., GeoI. Surv., 6 0 pp. Hunter, D.R. (compiler), 1962. The mineral resources of Swaziland. Swaziland Geol. Surv. Bull., 2: 111 pp. Hunter, D.R., 1963. The Mozaan Series in Swaziland. Swaziland Geol. Surv. Bull., 3: 516. Hunter, D.R., 1974. Crustal development in the Kaapvaal craton. 11. The Proterozoic. Precambrian Res., 1: 295-326. Key, R.M., 1977. The geochronology of Botswana. Trans. Geol. SOC.S. Afr., 80: 31-42. Klein, G. deV., 1971. A sedimentary model for determining paleotidal range. Geol. SOC. Am. Bull., 82: 2585-2592. Knowles, A.G., 1966. Palaeocurrent Study of the Ventersdorp Contact Reef at Western Deep Levels, Limited, on the Far West Rand. M.Sc. Thesis, Univ. Witwatersrand, Johannesburg (unpubl.). La Berge, G.L., 1966. Altered pyroclastic rocks in South African iron-formations. Econ. Geol., 61: 572-581. La Berge, G.L., 1973. Possible biological origin of Precambrian iron-formations. Econ. Geol., 68: 1098-1109. Malherbe, S.J., 1970. Flat pebble conglomerates in the Dolomite Series in the northern Cape Province, Dept. Mines, S. Afr., Geol. Surv., Ann., 8: 89-94. Martini, J.E.J., 1977. A copper-bearing horizon in the Pretoria Group in the northeastern Transvaal. Abstr., 17th Congr. Geol. Soc. S. Afr., Johannesburg, pp. 57-59. Mason, T.R. and Von Brunn, V., 1977. 3-Gyr-old stromatolites from South Africa. Nature, 266 : 47-49. Matthews, P.E., 1959. The metamorphism and tectonics of the pre-Cape formations in the post-Ntingwe thrust-belt, southwestern Zululand, Natal. Trans. Geol. SOC.S. Afr., 62: 257-323. Matthews, P.E., 1967. The pre-Karroo formations of the White Umfolozi inlier, northern Natal. Trans. Geol. SOC.S. Afr., 70: 39-63. Matthews, P.E., 1972. Possible Precambrian obduction and plate tectonics in southeastern Africa. Nature, 240: 37-39. Matthews, P.E. and Scharrer, R.H., 1968. A graded unconformity at the base of the early Precambrian Pongola System. Trans. Geol. SOC. S. Afr., 71: 257-271. McGregor, J.M., TruswelI, J.F. and Eriksson, K.A., 1974. Filamentous algae from the 2300 m.y. old Transvaal Dolomite. Nature, 247: 538-539. Nagy, L.A., 1974. Transvaal stromatolite: first evidence for the diversification of cells about 2.2 X l o 9 years ago. Science, 783: 514-516. Pretorius, D.A., 1975. The depositional environment of the Witwatersrand goldfields: a chronological review of observations and speculations. Miner. Sci. Eng., 7 : 18-47. Pretorius, D.A., 1976. The nature of the Witwatersrand gold-uranium deposits. In K.H. Wolf (Editor), Handbook of Strata-bound and Stratiform Ore Deposits. Elsevier, Amsterdam, Vol. 7 : 29-88. Truswell, J.F. and Eriksson, K.A., 1975. A palaeoenvironmental interpretation of the Early Proterozoic Malmani Dolomite from Zwartkops, South Africa. Precambrian Res., 2 : 277-303. Tyler, N., 1976. The pre-Chuniespoort Group succession around the Makoppa Dome, northwestern Transvaal. In: D.A. Pretorius (Compiler), Eighteenth Annual Report for

630 1976. Econ. Geol. Res. Unit, Univ. Witwatersrand, Johannesburg, pp. 16-17. Van Niekerk, C.B. and Burger, A.J., 1964. The age of the Ventersdorp System. Dept. Mines S. Afr., Geol. Surv. Ann., 3: 75-86. Van Niekerk, C.B. and Burger, A.J., 1968. Pb-isotope dating of the Zoetlief System, South Africa. Earth Planet. Sci. Lett., 4 : 211-218. Visser, H.N., Krige, L.J. and Truter, F.C., 1947. Explanation of sheet 6 4 (Ermelo). Dept. Mines, S. Afr., Geol. Surv., 9 1 pp. Visser, J.N.J., 1971. The deposition of the Griquatown glacial member in the Transvaal Supergroup. Trans. Geol. SOC.S. Afr., 74: 187-199. Visser, J.N.J. and Grobler, N.J., 1972. The transition beds at the base of the Dolomite Series in the northern Cape Province. Trans. Geol. SOC.S. Afr., 75: 265-274. Von Brunn, V., 1974. Tidalites of the Pongola Supergroup (Early Precambrian) in the Swart-Mfolozi area, northern Natal. In: A. Kroner (Editor) Contributions to the Precambrian Geology of Southern Africa. Chamber of Mines Precambrian Res. Unit, Univ. Cape Town, Bull., 15: 107-122. Von Brunn, V. and Hobday, D.K., 1976. Early Precambrian tidal sedimentation in the Pongola Supergroup of South Africa. J. Sediment. Petrol., 46 : 670-679. Von Brunn, V. and Mason, T.R., 1977. Siliciclastic-carbonate tidal deposits from the 3000 m.y. Pongola Supergroup, South Africa. Sediment. Geol., 18: 245-255. Walraven, F., 1977. The Transvaal-Waterberg stratigraphy in the central and eastern Transvaal. Abstr., 17th Congr., Geol. SOC.S. Afr., Johannesburg, pp. 142-146. Walter, M.R., 1972. Stromatolites and the biostratigraphy of the Australian Precambrian and Cambrian. Special Papers in Palaeontology, 11. The Palaeontological Association, London. Whiteside, H.C.M., 1970. Volcanic rocks of the Witwatersrand Triad. In: T.N. Clifford and I.G. Gass (Editors), African Magmatism and Tectonics. Oliver and Boyd, Edinburgh,pp. 73-87. Whiteside, H.C.M., Glasspool, K.R., Hiemstra, S.A., Pretorius, D.A. and Antrobus, E.S.A., 1976. Gold in the Witwatersrand Triad. In: C.B. Coetzee (Editor), Mineral Resources of the Republic of South Africa, Dept. of Mines, S. Afr., Geological Survey, Pretoria, 5th ed., pp. 39-73. Winter, H. de la R., 1963. Algal structures in the sediments of the Ventersdorp System. Trans. Geol. Soc. S. Afr., 6 6 : 115-121. Winter, H. de la R., 1965. The Stratigraphy of the Ventersdorp System in the Bothaville District and Adjoining Areas. Ph.D. Thesis, Univ. Witwatersrand, Johannesburg (unpu bl .). Winter, H. de la R., 1976. A lithostratigraphic classification of the Ventersdorp succession. Trans. Geol. SOC.s. Afr., 79: 31-48. Wyatt, B.A., 1976. The Geology and Geochemistry of the Klipriviersberg Volcanics, Ventersdorp Supergroup, south of Johannesburg. M.Sc. Thesis, Univ. Witwatersrand, Johannesburg (unpubl.).

Part B Broderick, T.J., 1976. Explanation to the geological map of the country east of Kariba. Rhod. Geol. Surv. Short Rep. No. 43. Button, A., 1976. The Strata of the Umkondo Basin, Southeastern Rhodesia. Econ. Geol. Res. Unit., Univ. Wits. 17th Annu. Rep. Carta Geologica, Provincia de Mopmbique 1 : 250,000 Folha SUL-E-36/J Lourenco Marques 1968. De Vries, W.C.P., 1969. Stratigraphy of the Waterberg System in the southern Waterberg area, North-western Transvaal. Ann. Geol. Surv. S. Afr., 7: 43-56. De Vries, W.C.P., 1973. Sedimentary structures in the southern and central portions of

631 the Waterberg area, north-western Transvaal. Geologie Mijnb., 52( 3): 133-140. Du Plessis, M.D., 1972. The relationship between the Bushveld Complex and the Waterberg System in the area between Loubad and Warmbad. Ann. Geol. Surv. s. Afr., 9: 85-87. Jansen, H., 1969. The structural evolution of the southern part of the Waterberg basin. Ann. Geol. Surv. S. Afr., 7: 57-62. Jansen, H., 1970. Volcanic rocks and associated sediments in the southern portion of the Waterberg basin. Ann. Geol. Surv. S. Afr., 8: 53-61. Jansen, H., 1976a. Precambrian basins on the Transvaal craton and their sedimentological and structural features. Trans. Geol. Soc. S. Afr., 78(1): 25-33. Jansen, H., 1976b. The Soutpansberg trough (Northern Transvaal) an aulacogen. Trans. Geol. SOC.S. Afr., 78(1): 129-136. Jones, D.L. and McElhinny, M.W., 1966. Palaeomagnetic correlation of basic intrusions in the Precambrian of Southern Africa. J. Geophys. Res., 71: 543-552. McElhinny, M.W. and Opdyke, N.D., 1964. The Palaeomagnetism of the Precambrian Dolerites of eastern Southern Rhodesia, an example of Geologic Correlation of Rock Magnetism. J. Geophys. Res., 69: 2465-2475. Meinster, B. and Tickell, S.J., 1976. Precambrian aeolian deposits in the Waterberg Supergroup. Trans. Geol. SOC.s. Afr., 78: 191-199. Slater, D., 1965. Structural and Stratigraphical Relationships in the Umkondo System of Eastern Southern Rhodesia - a Study of Rocks along Part of the Western Margin of the Moqambique Orogenic Belt. Ph.D. Thesis, Univ. Leeds (unpubl.). Slater, D., 1967. A Correlation of the Gairezi Series with the Umkondo System, Rhodesia. Trans. GeolSoc. S. Afr., 70: 177-87. Stocklmayer, V.R., 1975. Revised Stratigraphic Nomenclature for the Umkondo rocks. Ann. Rhod. Geol. Surv., I: 23-34. Swift, W.H., 1962. The Geology of the Middle Sabi Valley. South. Rhod. Geol. Surv. Bull., 52. Tickell, S.J., 1976. Braided river deposits in the Waterberg Supergroup. Trans. Geol. Soc. S. Afr., 78(1): 83-88. Tyndale-Biscoe, R., 1957. The Geology of a portion of the Inyanga District. South. Rhod. Geol. Surv. Short Report, 37. Vail, J.R., 1954. The Umkondo System: Late Precambrian sedimentation along the western edge of the Moqambique Orogenic Belt in Southern Africa. XXII Int. Geol. Congr. New Delhi, Vol. 10, Pt. 12. Vail, J.R., 1966. Zones of Progressive Regional Metamorphism across the western margin of the Mozambique Belt in Rhodesia and Mozambique. Geol. Mag., 103: 231-239. Watson, R.L.A., 1964. Contribution to discussion on Vail’s paper. XXII Int. Geol. Congr. New Delhi, Vol. 10, Pt 12. Watson, R.L.A., 1969. The Geology of the Cashel, Melsetter and Chipinga areas. Rhod. Geol. Surv. Bull.. 60.

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638 Van Biljon, S., 1949. Transformation of the Pretoria Series in the Bushveld Complex. Trans. Geol. SOC.S. Afr., 52: 1-172. Van Breemen, 0. and Dodson, M.H., 1972. Metamorphic chronology of the Limpopo belt, southern Africa. Geol. SOC.Am. Bull., 83: 2005-2018. Van der Merwe, M.J., 1976. The layered sequence of the Potgietersrus limb of the Bushveld Complex. Econ. Geol., 7 1 : 1337-1351. Vermaak, C.F., 1967. The Sominang iron ores of the Kunene Complex, Southern Angola. Internal report, J.C.I. Johannesburg. Vermaak, C.F., 1970. The geology of the lower portion of the Bushveld Complex and its relationship to the floor rocks in the area west of Pilanesberg, Western Transvaal. Geol. Soc. S. Afr., Spec. Publ., 1: 242-262. Vermaak, C.F., 1972a. Record of past work on the copper occurrences, Vila Serpo Pinto area, Angola. Internal report, Companhia de Minas Angolana, Luanda. Vermaak, C.F., 1972b. The economic significance of Angolan geology and geochronology. Internal report, Companhia de Minas Angolana, Luanda. Vermaak, C.F., 1976a. The nickel pipes of Vlakfontein and vicinity, Western Transvaal. Econ. Geol., 7 1 : 261-286. Vermaak, C.F., 1976b. The Merensky Reef - thoughts on its environment and genesis. Econ. Geol., 71: 1270-1298. Visser, D.J.L., 1957. The structural evolution of the Union. Trans. Geol. Soc., S. Afr., 60: xiii-xlix. Von Gruenewaldt, G., 1968. The Rooiberg felsite north of Middelburg and its relation to the layered sequence of the Bushveld Complex. Trans. Geol. Soc., S. Afr., 71: 153172. Von Gruenewaldt, G., 1970. On the phase change orthopyroxene-pigeonite and the resulting textures in the Main and Upper Zones of the Bushveld Complex in the Eastern Transvaal. Geol. SOC.S. Afr., Spec. Publ., 1: 67-73. Von Gruenewaldt, G., 1973a. The modified differentiation index and the modified crystallisation index as parameters of differentiation in layered intrusions. Trans. Geol. SOC.S. Afr., 76: 5 3 4 1 . Von Gruenewaldt, G., 1973b. The Main and Upper Zones of the Bushveld Complex in the Roossenekal area, Eastern Transvaal. Trans. GeoI. SOC.S. Afr., 76 : 207-227. Willemse, J., 1959. The “floor” of the Bushveld Igneous Complex and its relationships, with special reference to the eastern Transvaal. Proc. Geol. Soc.,S. Afr., 62: xxi-ixxx. Willemse, J., 1964. A brief outline of the geology of the Bushveld Igneous Complex: In: The Geology of some Ore Deposits in Southern Africa. Geol. Soc. S. Afr., 2: 91-128. Willemse, J., 1969. The geology of the Bushveld Igneous Complex, the largest repository of magmatic ore deposits in the world. Econ. Geol. Monogr., 4 : 1-22. Wilson, A.H., 1976. The Petrology and Structure of the Hartley Complex of the Great “Dyke”, Rhodesia. D. Phil Thesis, University of Rhodesia, Salisbury, Rhodesia (unpubl). Wilson, J.F., 1968a. The geology of the country around Mashaba. Bull. Geol. Surv. Rhod., 62: 239 pp. Wilson, J.F., 1968b. The Mashaba Igneous Complex and its subsequent deformation. Trans. Geol. Soc. S. Afr. 71, Annex: 175-188. Wilson, J.F., 1973. The Rhodesian Archaean craton - an essay in cratonic evolution. Philos. Trans. R. Soc., London, Ser. A, 273: 389-411. Wilson, J.F., 1979. A preliminary reappraisal of the Rhodesian Basement Complex. Geol. SOC.S. Afr., Spec. Publ., 5: 1-23. Worst, B.G., 1958. The differentiation and structure of the Great Dyke of Southern Rhodesia. Trans, Geol. SOC.S. Afr., 61: 283-354. Worst, B.G., 1960. The Great Dyke of Southern Rhodesia. Bull. Geol. Surv. South. Rhod. 47: 234 PP.

Worst, B.G., 1964. Chromite in the Great Dyke of Southern Rhodesia. In: S.H. Haughton (Editor), The Geology of some Ore Deposits in southern Africa. Geol. SOC.S. Afr., 2: 209-224. Wright, E.P., 1961. The Geology of the Gaberones District. D. Phil. Thesis, Univ. Oxford, 265 pp. (unpubl.). Yoder, H.S., 1954. The system diopside-anorthite-water. In: Annu. Rep. Director, Geophys. Lab., Carnegie Inst., Wash. Yearb., 53: 106-107. Yoder, H.S., 1955. Diopsideanorthite-water system at 5000 bars. Geol. Soc., Amer. Bull., 6 6 : 1638-1644. Yoder, H.S., 1968. In: Y.W. Isachsen (Editor) Origin of Anorthosite and related Rocks. N.Y. State. Mus. and Sci. Mem., 18: 13-22.

Part D Bliss, N.W., 1965. The Deweras Formation South of the Umfuli River, Rhodesia. M.Sc. Thesis, McGill Univ., Montreal (unpubl.). Bliss, N.W., 1970. The geology of the country around Gatooma. Rhod. Geol. Surv. Bull. No. 64. Bliss, N.W., 1971. The Deweras Group between the Umfuli and Umniati rivers. Trans. Geol. SOC.S. Afr., 71: 133-147. Bliss, N.W. and Tennick, F.P., 1964. Deweras and Lomagundi rocks between the Umfuli and Umniati rivers, Southern Rhodesia. Proc. 7th Annu. Congr. Geol. Soc. S. Afr., Salisbury, 1964, p. 47. Kirkpatrick, I.M., 1976. The geology of the country around Tengwe, Lomagundi District. Rhod. Geol. Surv. Bull. No. 75. Leyshon, P.R., 1969. The geology of the country around the Copper Queen. Rhod. Geol. Surv. Bull. No. 66. Macgregor, A.M., 1931. The geology of the country around the Norah, Molly and Umboe copper claims, Lomagundi District. South. Rhod. Geol. Surv. Short Rep. No. 25. Maufe, H.B., Lightfoot, B. and Molyneux, A.J.C., 1923. The geology of the country west of Sinoia, Lomagundi District. South. Rhod. Geol. Sum. Bull. No. 9. Molyneux, A.J.C., 1919. Preliminary report on the geology of the country west of Sinoia, Lomagundi District. South Rhod. Geol. Surv. Bull. No. 6. Phaup, A.E. and Dobell, F.O.S., 1938. The geology of the Lower Umfuli gold belt. South. Rhod. Geol. Surv. Bull. No. 34. Schidlowski, M., Eichmann, R. and Junge, C., 1976. Carbon isotope geochemistry of the Precambrian Lomagundi carbonate province, Rhodesia. Geochim. Cosmochim. Acta, 40: 449-455. Stagman, J.G., 1959. The geology of the country around Mangula Mine, Lomagundi and Urungwe Districts. South. Rhod. Geol. Surv. Bull. No. 46. Stagman, J.G., 1961. The geology of the country around Sinoia and Banket, Lomagundi District. South Rhod. Geol. Surv. Bull. No. 49. Stagman, J.G., 1962. The geology of the Southern Urungwe District. South. Rhod. Geol. Surv. Bull. No. 55. Sutton, E.R., 1979. Geology of the country around Mafungabusi. Rhod. Geol. Surv. Bull. No. 81. Tennick, F.P., 1976. The geology of the country around Magondi, Lomagundi, Hartley and Gatooma Districts. Rhod. Geol. Surv. Bull. No. 65. Thole, R.H., 1974. The Geology of, and Controls to, the Distribution of Copper a t the Shamrocke Mine near Karoi, Rhodesia. Ph.D. Thesis, Univ. of Rhodesia, Salisbury (unpubl. ). Thole, R.H., 1976. The geology of the Shamrocke Mine, Rhodesia: a stratiform copper deposit. Econ. Geol., 71: 202-228. Wiles, J.W., 1961. The geology of the Miami mica field, Urungwe District. South. Rhod. Geol. Surv. Bull. No. 51.

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641

Chapter 10

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

A. The Limpopo Belt

*

INTRODUCTION

The Limpopo Belt is a strip of high-grade polymetamorphic and polytectonic terrane which strikes east-northeast for some 680 km and separates the Rhodesian and Kaapvaal cratons by over 260 km (Fig, 10.1). The main tectonic and metamorphic events cover the time span 3200 to 2000 Ma, though the belt has subsequently acted as a site for faulting, sedimentation and igneous activity. The northern and southern limits of the belt are gradational t o the two cratons. In the west it may pass under young cover rocks of the Kalahari Basin, where the belt probably dies out. In the east it is truncated by the north-trending Moqambique Belt but this junction is obscured by Proterozoic and Mesozoic cover. The Limpopo Belt lies in a region of low rainfall so the rocks are remarkably fresh, though large areas are peneplaned and partly masked by Cenozoic sands. Though considerable regional mapping has been done by the Geological Surveys of Botswana, Zimbabwe and South Africa, coverage is incomplete. Cox et al. (1965) were the first t o note a megascopic zonation t o the belt; a Central Zone with structural trends oblique t o the belt are flanked on each side by the Northern and Southern marginal zones with structural trends roughly parallel to the belt. Mason (1970, 1973) drew attention to the granulite terranes of the marginal zones and t o the major shear belts which separate the marginal zones from the Central Zone, which is in turn characterized by a repeatedly metamorphosed basement and the metasupracrustals of the Limpopo Group. This broad subdivision of the belt is still satisfactory (Fig. 10.2 inset).

* By I.D.M. Robertson and M.C. du Toit.

642

20' S

Fig. 10.1. The Limpopo Belt in relation t o the cratons and metamorphic belts of southern Africa.

THE NORTH MARGINAL ZONE

This zone falls entirely within Zimbabwe and Botswana. It consists mainly of reworked granite-greenstone terrane, now at the amphibolite or granulite facies (Worst, 1962; Gerrard, 1965; Robertson, 1968; Mason, 1970; Bennett, 1971; Odell, 1975) which grades into the Rhodesian craton t o the north and passes south, via the Tuli-Sabi Shear Belt, into the Central Zone. Craton margin and marginal granites

The northern limit of the Limpopo Belt is difficult t o define. Where the granulite terrane passes transitionally into the Rhodesian craton the orthopyroxene isograd is commonly taken as the northern margin of the Limpopo Belt but tectonic and metamorphic elements of this belt persist to the north of the isograd. A transitional region can be recognized in Zimbabwe north of the isograd, containing greenstone belt rocks, similar to those of the craton,

647 which have been converted to highly deformed amphibolite facies schist belts, aligned parallel t o the northeast tectonic trend of the North Marginal Zone (Fig. 10.2) (Worst, 1962; Robertson, 1973a,b). In Botswana, there is no orthopyroxene isograd but metamorphic and tectonic effects increase progressively southward from the southern edge of the Tati Greenstone Belt (Mason, 1970; Bennett, 1971). North of, but parallel to, these craton marginal schist belts lies a chain of massive, even-grained granite plutons (Fig. 10.2, The Tuli, Chibi and Zimbabwe granites) which extend some 400 km parallel to the Limpopo Belt, and appear to have been derived by partial melting from within the belt (Robertson, 1973a,b; Hickman, 1976; Hickman, 1978). South of these plutons and the schist belts is an unusual Porphyritic Granite which occurs both as discrete sheets (Robertson, 1973a,b; Odell, 1975) and as zones of similar porphyroblastic augen gneiss (Mason, 1970; Bennett, 1971; Key, 1977). Chemically this “granite” is characterized by low silica (68%) and high iron, calcium, titanium and phosphorus contents. South of the Mweza Schist Belt the lit-par-lit contacts of this “granite” lie roughly parallel to the southdipping grain of the host gneisses and are marked by a strong marginal south-southeastdipping foliation and downdip quartz and feldspar rodding. The intrusions have relatively massive cores cut by transverse shears. The Porphyritic Granite within the granulite terrane near the Tuli-Sabi Shear Belt (Odell, 1975) is coarser grained, has interfingering contacts with the host gneisses, a marked foliation and no truly massive core. In general Porphyritic Granite intrusions show increased strain and metamorphism on passing from the craton into the Limpopo Belt where they become augen gneisses. The feldspars become darker, clinopyroxene occurs in some, foliation is more marked and the interdigitated contacts become tighter. Sites of high strain seem to favour emplacement of Porphyritic Granite, which was probably emplaced late-syntectonically, just after the second amphibolite-granulite metamorphic peak, as a series of southdipping pods (Robertson, 1973a,b; James, 1975). It is possible that these granites predate the even-grained granites of the craton margin (Hickman, 1976). Amphibolite-granulite terrane This terrane consists of granite-greenstone material metamorphosed to granulite or amphibolite facies. Swift (1962) and Worst (1962) suggested that, in the east, the granulites were separated from the craton by a major thrust. James (1975) and Odell (1975) described mylonites developed in cataclastic-textured microcline-biotite gneisses along this contact. Robertson (1968, 1973c) working to the southwest, and Mason (1970) and Key et al. (1976), still further west in Botswana, were unable to find a major fault and considered the contact to be a structural and metamorphic gradation.

648 The granulite terrane in Zimbabwe (Fig. 10.2) consists mainly of a sea of granitic granulite gneiss * surrounding charnockites, enderbites and scattered, deformed and folded metasedimentary and meta-igneous inclusions. The charnockites and enderbites occur both as narrow, infolded gneissic bands or as large; relatively homogeneous bodies (Odell, 1975). The granulite gneiss becomes areally far more important towards Botswana. In the east l't tends to be gametiferous, possibly implying metamorphism at a deeper crustal level. The inclusions are thought t o represent remnants of greenstone belt material (Swift, 1962; Worst, 1962; Robertson, 1968, 1973a,c, 1974). The most common inclusions are of mafic granulites. These granoblastic rocks have a general assemblage of pyroxene-plagioclase ? hornblende biotite. They have been equated t o the amphibolites of the craton margin schist belts and to basaltic rocks of the greenstone belts. Banded ironstones occur as lensoidstructured quartz-magnetite-grunerite rocks in the craton margin, the grunerite being dehydrated to hypersthene in the granulite terrane. They form ribbons in the granulite gneiss or are closely associated with mafic granulites. Rare quartzites also occur. M a r k are represented by laminated, compositionally variable, skam-like calc-silicate gneisses with granulite or near granulite mineral assemblages, and are associated with other greenstone belt remnants. Rather rare sapphirine-bearing aluminous rocks occur within the granulite zone. Robertson (1973a) suggested they had an argillaceous origin. Granulite facies remnants of mafic-ultramafic complexes, which are discussed later, have been recognized in this terrane. The metamorphic facies in the North Marginal Zone decreases towards the southwest of Zimbabwe towards Botswana (Fig. 10.3). First the charnockiteenderbite suite is replaced by pyroxene-free granulite gneiss and then the mafic granulites pass into amphibolites. As no detailed petrographic work has been done in this zone, it is not yet clear whether this is due t o regional retrogression of the granulites or that granulites were not developed here in the first place. In eastern Botswana metamorphism increases progressively southwards, across the craton margin, from the greenschists of the Tati Greenstone Belt to amphibolites and diopside-bearing amphibolites south of the Tati pluton (Mason, 1970). South of here, north of the Monachane Structure and Mmadinare Shear (Figs. 10.2, 10.4j, poorly foliated biotite gneisses, set with a few magnetite quartzites, contorted amphibolites and talc-tremolite schists, probably represent reworked cratonic material. To the west, assemblages analogous t o the Limpopo Group and the Tonota and Matsitama schists pass t o the northwest into what was previously known as

*

* A rather contradictory but useful field term introduced by Worst (1962) for a granitic gneiss which is pyroxene-free but contains high-temperature feldspars, otherwise typical of the granulite facies.

649

COVER ROCKS AMPHIBOLITE FACIES

n

GRANULITE FACIES

0

CRATONIC GNEISSES GREENSTONES AND IGNEOUS COMPLEXES

Fig. 10.3. The distribution of metamorphic facies in the Limpopo Belt, as seen at present.

the Shashe Belt by Crockett (1968), Bennett (1970) and Mason (1973). Key and others (1976), Coomer et al. (1977) and Key (1977) regard these rocks as forming a sequence of sediments and metavolcanics which underlie, but are an integral part of, the Tati Greenstone Belt. Ma fic-ul tramaf ic complexes A number of remnants of pre-metamorphic mafic-ultramafic complexes occur scattered among the granulites of the North Marginal Zone (Worst, 1962; Robertson, 1973a; Odell, 1975). In some places mesocopic slivers of metapyroxenite, metadunite or serpentinite are all that remain; in others the ultramafic assemblage occurs together with laminated, coarse-grained metagabbros and metanorites and rarely with laminated or mottled metaanorthosites. Though the textures of these nickel- and chromium-rich rocks are metamorphic, their chemistry, structure and association are typically igneous. Small pods and lenses of chromitite occur within some of these complexes. The component rocks of the complexes are also pod-like, due to tectonic interslicing, and some are closely associated with, and regionally conformable to, magnetite quartzites. They appear t o have been emplaced into penecontemporaneous greenstone belt material and were later metamorphosed along with it. They differ from the complexes of the Central Zone, in having relatively large ultramafic and very minor anorthosite components and their associated chromitites have higher Cr/Fe ratios. This could

650 reflect an ultramafic (komatiitic) initial magma as opposed to the mafic parent of those of the Central Zone.

Structure and metamorphism There is a strong east-northeast to northeast structural grain throughout most of the North Marginal Zone. Regional foliation trends, lithological boundaries and major fold axes lie parallel or sub-parallel t o this grain. Early isoclinal folds are refolded by major open folds which in places close into domal structures. To the south of Bangala Dam the structures become progressively more linear, charnockites are rodded and a concentration of dolerite dykes in a strongly developed regional foliation herald the approach to the Tuli-Sabi Shear Belt. Within the area around Bangala Dam, James (1975) recognized a main prograde granulite facies metamorphism, estimated at about 750-800" C at >5 kb, which seems to have obliterated all evidence of any previous metamorphic history. It was an essentially static metamorphism with some partial melting and intrusion of charnockite and has been dated (Rb-Sr errorchron) at 2930 f 120 Ma by Hickman (1976) who considered that these rocks had only a brief history prior to that date. The granulite metamorphism was followed by a deformation which produced major and minor northeast trending upright folds but no fabric. A later deformation produced variable strain, in some places inducing only a quartz-ribbon fabric, in others a cataclastic fabric. It produced several east-northeast-striking major braided shear zones which widen to the north and the rocks affected by them have been retrogressed, also with increasing intensity to the north. The Porphyritic Granites may have been emplaced about this time. It is possible that this was accompanied by another amphiboliteegranulite facies metamorphism (Hickman, 1976) at about 2700 Ma. Coward et al. (1976) explained this latter deformation, and an associated deformation in the Rhodesian craton, by movement of the Rhodesian craton to the southwest, parallel to the Limpopo Belt, producing heterogeneous zones of shear separated by less deformed lenses. A swarm of unmetamorphosed and undeformed mafic dykes related to the Great Dyke, errorchron dated at 2600 f 120 Ma (Robertson and Van Breemen, 1970), transect the granulites t o the North Marginal Zone (Fig. 10.2) and provide a minimum age for the high grade metamorphism and deformation. THE TULI-SABI SHEAR BELT

The Tuli-Sabi Shear Belt (Fig. 10.4) extends east-northeast for the length of the Limpopo Belt and is characterized by large-scale cataclasis. On a regional scale it seems to form a series of roughly en-khelon shear zones each of limited strike length. Both granulite and Limpopo Group rocks have

651

Fig. 10.4. The distribution of major shear belts, younger, unmetamorphosed cover rocks and igneous complexes in the Limpopo Belt.

been identified in the shear zone, though most are refoliated and high-temperature mineral assemblages have been cataclastically retrogressed to micas and chlorites (Swift et al., 1953; Mason, 1970). The Mmadinare Shear of eastern Botswana terminates in the west, just south of the Lethlakane Fault (Key and Hutton, 1976) and to the east near the Zimbabwean border. It consists of a zone of bands of steep, southdipping mylonite about 5 km wide which grades northward over a further 10 km, via flaser and augen gneisses, into the cataclastically affected rocks of the Monochane Structure (Mason, 1970). The mylonites consist of a succession of flaggy gneisses with black flinty crush interlayers containing scattered kernels of white porphyroclastic feldspar. Early isoclinal folds have been largely obliterated by intense refoliation. The isoclinal fold axes lie in the cataclastic foliation and are parallel to shallow dipping rod, mullion and other linear structures. A similar shear to the Mmandinare Shear is displaced en4chelon to the south and continues eastward close to the Karoo contact (Fig. 10.4). The Tuli-Sabi Shear Belt remains largely unmapped in Zimbabwe but reconnaissance work (Mason, 1970, 1973; Coward et al., 1976) indicates that it breaks up into a zone of fault splays. Here the cataclasites dip shallowly south. Their northern edge is heterogeneous, consisting of lozenge-shaped blocks of unsheared material separated by intensely sheared gneiss and intruded by dolerite. A quartz-leaf and feldspar-augen linear fabric is gently

652 eastward dipping. Rotation of porphyroblasts and curvature of earlier foliations into the shear belt suggest dextral movement. Mason (1973) and Wakefield (1974) also noted some vertical and rotational components. Coward et al. (1976) argued that the shallow-dipping shear belt in Zimbabwe represents a zone of d6collement on which the Central Zone slid westwards on the rigid granulites. They proposed progressively greater displacements t o the east, the discrepancies being taken up in refolding of Central Zone rocks. Mason (1973) doubted if strain calculations are altogether valid in a ductile shear belt of the magnitude and complexity of the Tuli-Sabi Shear Belt. On a regional scale, in Botswana, Key and Hutton (1976) noted that latestage overfolding occurs t o the east (north of the shear-zone) and to the west (south of the shear zone). They suggested generation of these overfolds by movement along the shear zone. It seems more probable that the shear was a late result of the same couple which produced the overfolds, continuing to act on rocks of decreasing plasticity (metamorphic grade), rather than being their cause. Key and Hutton argued for termination of the Limpopo Belt at the western extremity of the Mmandinare Shear. Such an abrupt termination of the Limpopo Belt seems unlikely in view of the discontinuous, en-kchelon nature of these shear zones on a regional scale, though it is probable that both the shear belt and the Limpopo Belt itself may die out under the Kalahari Basin (Reeves and Hutchins, 1975). Hickman and Wakefield (1975) considered that Rb-Sr ages from minerals and thin-slice isochrons grouped around 2000 Ma in the Pikwe area, adjacent to Mmadinare shear zone, probably date the polyphase deformation which ended in the transition from ductile t o brittle conditions which produced the Tuli-Sabi Shear Belt. However, it is possible that the Tuli-Sabi Shear Belt originated much earlier as a zone of weakness during the shear deformation of the granulites at about 2600 Ma but only developed its present mylonitic characteristics much later at about 2000 Ma. The Lethlakane fault was formed still later. T H E CENTRAL ZONE

The Central Zone is a distinctive unit flanked t o the north and south by major shear zones and granulite terranes. It is characterized by complex relationships between what appear t o be a repeatedly metamorphosed basement and an intensely metamorphosed cover sequence. Major anorthositic complexes intrude near the cover-basement interface. It is also characterized by large scale, relatively open folds and dome and basin structures which generally lie across the trend of the Limpopo Belt. The cover sequence is problematic. Sohnge et al. (1948) introduced the term “Messina Formation” which they restricted to a sequence of high-grade metaquartzites, magnetite quartzites, dolomites and marbles near Messina in South Africa. This term was later broadened by Mason (1965, 1967),

TABLE 10.1 Nomenclature and sample stratigraphies of the Central Zone ,ithostratigraphic nomenclature houp

Subgroup

Sample stratigraphies

Formation

Messina

Pikwe

Mt. Dowe me ta-quartzite

quartzite with lenses of magnetitequartzite, mafic granulite and amphibolite

marble, impure quartzite, aluminous gneiss, magnetite-quartzite and amphibolite

Artonvilla cordierite gneiss

aluminous gneiss with lenses of calcsilicate gneiss and magnetite-quartzite

aluminous gneiss

Messina amphibolite (diopside-garnet) Alldays marble

impopo

marble, c a lc silic a te gneiss and magnetite-quartzite

hornblende-biotite gneiss with aluminous horizons

ore-bearing amphibolite banded pink gneiss with intercalations of aluminous gneiss hornblende gneiss and amphibolite

anorthositic gneiss and amphibolite

anorthositic gneiss

Maryland lcucocratic gneiss

garnetiferous leucogneiss with intercalations of garnet-biotite gneiss, calcsilicate, amphibolite and magnetitequartzite

banded quartzo-feldspathic gneiss

Magdala biotitegarnet gneiss

calc-silicate gneiss and magnetitequartzite

~

Shanzi

dacuville

nebulitic t o charnockitic granite-gneiss basement

porphyroblastic granite-gneiss ment

base-

Lithostratigraphic nomenclature from unpublished work of the South African Committee for Stratigraphy and applies mainly to rocks of the Messina area. Similarities with the Pikwe area (Botswana) are implied only in the broadest sense. Messina after Bahnemann (1972), and Pikwe after Hickman and Wakefield (1975). Meta-igneous rocks in italics.

Q,

cn w

654 working in Botswana, t o include quartzo-feldspathic paragneisses, amphibolites, and other rocks intimately associated with and having similar structural elements t o the “Messina Formation” (his Baines Drift Formation). The latest lithostratigraphic nomenclature for these rocks is given in Table 10.1. The old term Messina Formation of Sohnge et al. (1948), which has been retained in Zimbabwe (Watkeys, 1976), now has subgroup status. Underlying this is the Shanzi Subgroup, equivalent t o the Leucocratic Gneiss Formation in Zimbabwe. Both subgroups now comprise the Limpopo Group, also variously known as the Beitbridge Group in Zimbabwe (Watkeys, 1976), Assemblage IIA or the G2 Paragneisses of Botswana (Bennett, 1971; Key et al., 1976, respectively). All these rocks rest on a basement of Macuville Group (Zimbabwe and South Africa) and appear t o rest on Assemblage IIB (Botswana). The structural complexity, high metamorphic grade and variability caused by facies changes over a wide area, has made unravelling of the stratigraphy difficult. I t must be stressed that the stratigraphy is generally an apparent stratigraphy as conclusive facings are rarely found. In compiling the map (Fig. 10.2) it was not possible to distinguish the Shanzi Subgroup in South Africa and Botswana, so the Limpopo Group encompasses much of this element, although the distinction has been made in Zimbabwe.

Ancient basement Rocks basal to the Limpopo Group (Macuville Group) occur as rather illdefined inliers within the supracrustals (Sohnge, 1946; Mason, 1970, 1973; Bahnemann, 1971, 1972; Bennett, 1971; Key et al., 1976; Barton et al., 197713; Key, 1977; Light et al., 1977). Correlation between them is not yet possible. In comparison t o the Limpopo Group they tend to be relatively homogeneous, rather nebulitic, generally tonalitic gneisses with extremely complex structures and contain altered mafic dykes and a few relics with granulite facies assemblages. Messina Subgroup assemblages (defined later) are characteristically absent. Mason (1973) regarded them as possible granitegreenstone terrane, now remetamorphosed granulite. The former is difficult to prove, the latter very probable. Mason (1970) and Bennett (1971) described extensive tracts of relatively homogeneous, migmatitic granitic gneisses in Botswana which are regionally conformable to, and are interdigitated along their margins with, the cover assemblages. Mafic and pyroxene-bearing granulites and associated mafic t o ultramafic rocks form. a very small but significant component. Bahnemann (1972) described an ancient crystalline basement, the Bulai Belt near Messina, consisting partly of Bulai “Granite” (discussed later) and partly of various associated nebulitic granulite gneisses, some containing hypersthene (Fig. 10.5). Light et al. (1977) described the Macuville Group as banded, migmatitic, generally nebulitic, garnet-poor orthogneiss with a quartzdioritic to granodioritic composition. Bands of mafic pyroxene-bearing granulite crosscut the foliation. Bahnemann (1971) and Barton et al. (1977b) have noted an occurrence of

Fig. 10.5. Part of t h e Central Zone, a photogeological interpretation of the Messina Area after Sohnge (1946) and Bahnemann (1972), showing the nebulitic “Bulai Belt” basement (northwest), the crossfold belt of Limpopo Group and meta-anorthosite (centre) and t h e linear zone (southeast).

cn

656 ancient basement gneisses in the bed of the Sand River (Fig. 10.5) some 10 km south and southeast of Messina. The Sand River rocks consist of steep-dipping, layered, peraluminous, medium- to coarse-grained, grey to leucocratic gneisses showing small-scale interference fold patterns. They are preserved in the core of a nappe-fold unit. The gneisses are cut by early quartzo-feldspathic dykes and by at least two generations of discordant but deformed amphibolite layers interpreted as mafic dykes. Neither the quartzo-feldspathic nor the mafic dykes have been recognized in the surrounding metasedimentary rocks. The Sand River gneisses have a Rb/Sr isochron age of 3858 f 116 Ma and the two dyke suites have isochron ages of 3643 2 102 and 3128 f 84 Ma (Barton et al., 1977b). The relationship between cover and basement is very variable. In places the junction between the two can be recognized on a regional scale but, in others, complex infolding coupled with partial melting have rendered this junction indistinct (Mason, 1973). Poor exposure and a lack of detailed mapping throughout large portions of the Central Zone have contributed to this problem. The most common relationship is where partly remobilized and remelted basement has pushed upward into the cover, producing major domal structures separated by tightly folded synclinal zones. The cores of the domes are occupied mainly by basement material but also contain infolded and xenolithic fragments of resistant cover, typically metaquartzite or marble. Locally slices of basement occur tectonically pushed into the cover and infolded with it (Mason, 1973). Bahnemann (1972) interpreted the narrow northwest-trending synclinal folds of the Limpopo Group (his cross-folded belt), which penetrate the Bulai Belt, as reflecting a ductility contrast (Fig. 10.5). He proposed that the dry competent, crystalline Bulai Belt or nebulitic basement tended to rupture while the comparatively wet, mobile Limpopo Group metasediments tended to flow into the zones of weakness. To the southeast the Messina Subgroup rocks were highly mobile and produced a zone of strong linearity characterized by very tight folds. Limpopo Group A profusion of metaquartzites, aluminous and calc-silicate gneisses and carbonate rocks characterize the Limpopo Group. Sample stratigraphies at Pikwe and Messina are given in Table 10.1. Most of this group consists typically (from the base) of leucocratic gneiss, marbles, calc-silicate gneiss, slightly aluminous gneiss and quartzite with minor intrusion of anorthosite, serpentinite and gabbro. Locally the stratigraphy is complex, as at the Monte Cristo and Messina (Fig. 10.2), where the calcareous rocks are partly substituted by magnetite quartzite and highly aluminous, garnetiferous and cordierite-bearing biotite gneisses called the Artonvilla Cordierite Gneiss Formation, which could represent an abnormal depositional environment in these two localities.

657 Bahnemann (1972) suggested that the Messina Subgroup is a shelf or platform assemblage with shale at the base, overlain by calcareous and ferruginous shale with banded ironstone and beds of dolomite near the middle and clean quartz sands with interfingering of feldspathic sands, sandy shale and calcareous shale at the top. It has not yet been possible t o date the Messina Subgroup directly, although a minimum age of 3221 ? 48 Ma has been provided by anorthosite complexes intruded into it near Messina (Barton et al., 1977a).

Anorthositic complexes The anorthositic complexes comprise a range of rocks including metaanorthosite, anorthositic metagabbro, metagabbro, amphibolite, hornblendite, minor metapyroxenite and serpentinized metadunite that occur as layered bodies lying concordantly t o semi-concordantly within or near the base of the Limpopo Group. Sohnge et al. (1948), Van Eeden et al. (1955), Hor et al. (1975) and Barton et al. (1977a) recorded chromitite and magnetite layers in serpentinites, hornblende gneisses and hornblende-bearing metaanorthosites. Most regard these as remnants of layered complexes intruded into supra-crustal rocks and subsequently involved in high-grade metamorphism. Significant meta-anorthosite bodies occur at Pikwe, northeast of Maasstroom, east of Alldays and near Baines Drift though the most carefully studied are a number of layers which comprise the Messina Intrusion which extends across the Limpopo River near Messina. Though the Messina Intrusion has been extensively recrystallized by multiple metamorphism, Hor et al., (1975), Barton et al. (1977a) and Light et al. (1977) have been able to distinguish an igneous stratigraphy, starting with what was layered pyroxenite, passing through gabbroic anorthosite, hornblende anorthosite to garnet anorthosite and have also recognized remnants of megacrysts, ophitic textures, primary cumulate textures, and cumulate settling cycles. The Messina Intrusion (Fig. 10.5) has been complicated and repeated by nappe folding (Barton et al., 1977a). Competent meta-anorthosites tend to be boudinaged but their associated amphibolites and ultramafics have been deformed plastically. Plugs, and oval bodies of serpentinite with rare meta-pyroxenite occur throughout the Central Zone, in places associated with fold noses and hinge lines, and could be of a later age (Light et al., 1977). Bahnemann (1972) has provided convincing evidence of a sedimentary origin (siliceous dolomitemagnetite quartzite association) for some serpentinites. Variable-sized, mesoscopic to megascopic inclusions of ultramafic and mafic rocks (not shown on Fig. 10.2) form a broad 230 km-long zone from near Mahalapye northwards to beyond Pikwe in Botswana, within which lie the Pikwe-Selebi Ni-Cu deposits. The largest group (Lear, 1971) is in the

658 west, at or near the junction between basement and cover. They consist of homogeneous or composite boudinaged remnants of what were probably much larger differentiated masses. Most are lenticular and are conformable to, and have been folded with, the enclosing gneisses, though some are pluglike and discordant. The majority of these inclusions occur within gneisses which are rich in supracrustal remnants and only rarely occur as xenoliths in the more homogeneous remobilized basement. Rock types such as serpentinized metadunites and metaperidotites, olivine metapyroxenites and tremolitic to hornblende amphibolites are well represented. Many are mantled by amphibolites, biotite schists and hornblende gneisses and are injected by pegmatitic material. Their textures are generally xenomorphic. In one of the larger, relatively undeformed complexes in the east Lear (1971) recognised a differentiated ultramafic to mafic sequence complete with layered chromitites and in a nearby similar complex he reported traces of original cumulus textures. Both complexes have a rather uncertain relationship to nearby meta-anorthosites.

Bulai “Granite” Sohnge (1946) defined the Bulai “Granite” (axtually a granitic, in places augen-rich, gneiss) as a coarse-grained, porphyritic, red-brown to grey biotite granite which occurs mainly in the area north-west of Messina (Fig. 10.2). Bahnemann (1972) referred t o the “Bulai Belt’’ in which he included the Bulai “Granite Gneiss” proper as well as other gneisses which he regarded as a basement t o the Messina Formation (Fig. 10.5). Both Sohnge (1946) and Light et al. (1977) reported numerous, partly assimilated xenoliths of the surrounding gneisses and metasupracrustals within the Bulai “Granite” but Bahnemann (1972) interpreted these as boudinaged fold noses, ripped off and included in the basement. Recent Rb/Sr dates ranging from 2690 to 2772 Ma (Van Breemen, 1968; Barton and Ryan, 1977) clearly showed that the Bulai “Granite” is younger than the Limpopo Group. It was probably derived syntectonically by anatexis of the basement at about 2700 Ma but, due to its water content, was incapable of rising significantly. The Bulai “Granite” is structurally, petrographically, chronologically and chemically analogous to the Porphyritic and Matok granites of the North and South Marginal Zones, respectively, and is also very similar to various coarsegrained porphyroblastic augen gneisses with variable t o strong fabrics which occur both in the North Marginal and Central Zones in Botswana (Mason, 1970; Bennett, 1971).

Singelele “Granite ” The Singelele “Granite” (also a granitic gneiss) was defined by Sohnge (1946) as a relatively homogeneous red t o pinkish-grey even-grained rock

659 consisting of microcline-perthite, quartz and subordinate hornblende, plagioclase and garnet. It typically occurs in the cores of anticlinal zones as a series of progressive, regionally cross-cutting phacoliths near Messina (Fig. 10.5). A biotitic variety occurs on the flanks of refolded anticlines and both an intermediate, hornblende-rich variety and a red leucocratic, mesoperthitic variety occur in the cores of refolded anticlines. It also occurs as a facies marginal to the Bulai “Granite”. Bahnemann (1972, 1973) has suggested that the Singelele “granite gneiss” was derived by partial melting of a tonalitic basement. Light et al. (1977) regarded the Singelele “gneiss” as a definite stratigraphic horizon which occurs within the Shanzi Subgroup. This controversy can be resolved if the Singelele “Granite” is regarded as part of the Shanzi Subgroup which was particularly prone to melting.

Structure and metamorphism

(a) Pik we Wakefield (1974) has established three tectono-metamorphic events in the Pikwe area in Botswana (see also Key, 1977). A first fabric-producing event folded the rocks about a north-northeast-trending axis and was accompanied by high-grade regional metamorphism (estimated at 770”C and 7.6 kb) which produced cordierite and garnet in pelitic rocks and extensive partial melting. Granulite assemblages were not formed due t o the presence of large amounts of water. This metamorphism has been dated at 2660 k 60 Ma (Hickman and Wakefield, 1975). Though Key and Hutton (1976) have reported fold hinges predating this event, it is difficult t o assess the state of these rocks prior t o this high-grade metamorphism. A set of mafic dykes was emplaced which separates this event from subsequent ones. The next event started with large scale recumbent folding overturned t o the northeast. It was accompanied by amphibolite facies metamorphism producing andesine-hornblende-garnet assemblages in the mafic dykes, and localized partial melting. This event culminated in the development of flatlying cataclastic fabrics by intensification of the earlier recumbent folding. A third deformational event produced early dome and basin structures, was accompanied by lower amphibolite facies metamorphism, which annealed the previously formed cataclastic structures, and ended in shearing (Tuli-Sabi Shear Belt). There seems t o have been a transition from ductile t o brittle behaviour accompanied by a decrease in metamorphic grade from lower amphibolite t o greenschist facies. Thin slice and mineral ages of about 1900-2100 Ma (Van Breemen, 1968; Hickman and Wakefield, 1975) probably date this reheating event which may also mark uplift of these rocks. ( b ) Beit bridge Light et al. (1977) reported a long structural and metamorphic history in the Beitbridge area. The Macuville Group appears t o have been folded and

660 metamorphosed prior to deposition of the Beitbridge Group. Then both these groups suffered isoclinal recumbent similar folding on an east-northeast trend. The limbs of the folds became highly attenuated, resulting in horizontal shears and repetition of units. This was followed by relatively static high-pressure granulite metamorphism, producing garnet porphyroblasts included with randomly orientated kyanite crystals in some rocks. Later isoclinal folding and co-axial refolding on east-northeast- and northnortheast-trends were accompanied by sillimanite-granulite metamorphism, which culminated in basement anatexis and production of the Bulai “Granite” (2700 Ma). Southeast trending flexural slip folds and finally mylonites and flaser gneisses developed in flat-lying cataclastic zones. This was accompanied by amphibolite grade metamorphism which retrogressed some of the granulites and produced late staurolite in the mylonites. (c) Mahalapye The dominant structural trends of the Messina Formation swing t o the west or west-northwest in the extreme west end of the Limpopo Belt near Mahalapye in Botswana (Jennings, 1963; Crockett, 1965; Ermanovics, 1977). Here the Limpopo Group rocks have been folded into upright isoclines, disrupted and migmatized t o varying degrees and intruded by a series of relatively homogeneous, late-synkinematic t o postkinematic, slightly foliated to massive granites (Bennett, 1971; Ermanovics, 1977). Microcline is the dominant feldspar and garnet is distinctive. This zone of intense migmatization, characterized by relatively late structures (Ermanovics, 1977), is separated from the main body of the Central Zone in Botswana by the Sunnyside Shear Zone (Fig. 10.4). I t is regarded as part either of the South Marginal Zone (Cox et al., 1965; Ermanovics, 1977; Key, 1977) or of the Central Zone (Mason, 1973). Heterogeneous migmatites at Losi quarry have been dated at 2240 f 400 Ma (Van Breemen, 1968) with a minimum age of 2010 k 80 Ma. The error is high so the precise chronological position of this migmatization is not yet clear. THE SOUTH MARGINAL ZONE

The South Marginal Zone, lying to the west and south of Louis Trichardt is developed along, and shows a metamorphic transition to, the Kaapvaal craton. It is a structurally and metamorphically complex zone roughly analogous t o the North Marginal Zone but it is poorly exposed (Fig. 10.2). Its northern margin, adjoining the Central Zone, is largely obscured by Waterberg, Karoo and more recent cover. Mason (1973) regards the Soutpansberg fault zone as a reactivated basement fracture akin to, and a mirror image of, the Tuli-Sabi Shear Belt (Fig. 10.4). Highly sheared cataclastic gneisses and flinty crush zones are exposed in this zone north of Villa Nora, but the extent of this shear belt is not yet known.

661 Craton margin The southern limit of the South Marginal Zone appears t o be gradational to the Kaapvaal craton (Du Toit and Van Reenen, 1977; L. Wilsenach, pers. comm., 1975). Greenstone remnants become progressively migmatized, deformed and aligned t o the tectonic trend on passing northwards from the craton edge. Diopside, grunerite and ortho-amphibole appear in amphibolites, ironstones and metapelites respectively, followed by orthopyroxene in all as the grade increases. Van Reenen and Du Toit (1977) have distinguished two isograds adjoining the Kaapvaal craton. To the south lies a lower grade ortho-amphibole isograd, north of which ortho-amphibole and garnet are stable over a 10-15 km-wide strip, transitional t o the higher grade orthopyroxene isograd (Fig. 10.2).

Banderlierkop Formation The granulite terrane, from north of the orthopyroxene isograd to where it is covered by sediments of the Soutpansberg Trough, consists of disjointed and contorted, keel-like remnants of pelitic, quartzo-ferruginous, mafic and ultramafic rocks called the Baderlierkop Formation by Du Toit and Van Reenen (1977). They provisionally equated this with schist belt material as in the Sutherland Range, Pietersburg and other nearby schist belts (Fig. 10.2). They consist dominantly of metapelitic quartz-biotitegarnet-orthopyroxene-cordierite assemblages, though in similar rocks, between the isograds, ortho-amphibole replaces the orthopyroxene. The mafic rocks (orthopyroxene-clinopyroxene-plagioclase k quartz) tend to a komatiitic composition and the ultramafics are metapyroxenites and metadunites. The remnants of the Banderlierkop Formation are enclosed in a generally tonalitic, orthopyroxene-bearing, migmatitic gneiss. It is not yet clear whether this gneiss was a floor on which the Banderlierkop Formation was deposited or is, in part, later intrusive material.

Structure and metamorphism Detailed work on the limited good exposure has allowed Barton and Ryan (1977), Du Toit and Van Reenen (1977) and Van Reenen and Du Toit (1977) to outline the structural, metamorphic and some of the chronological history of the Banderlierkop Formation. Two early deformations produced plastic, isoclinal, co-axial folds with steep axial planes and two early metamorphisms produced granulite assemblages. It is not possible t o relate the first metamorphism to the early co-axial deformations. The second metamorphism post-dates the second deformation but predates the third and last deformation. It was not possible t o estimate the time span between these early metamorphisms. The first, high-pressure granulite metamorphism

662 generated garnet and hypersthene. During the second, low-pressure granulite metamorphism, garnet was unstable and formed cordierite and a second generation of hypersthene. Anatectic melts segregated during both metamorphisms but were more abundant in the second metamorphism. This second period of metamorphism and anatexis, which reset the Rb-Sr chronology, is provisionally dated at 2636 Ma and is identical to the age of the 2645 Ma-old Matok Granite. This porphyroblastic biotite granite, with its associated granodioritic to gabbroic phases, can be closely matched to the Bulai and Porphyritic “Granites” of the Central and North Marginal Zones, respectively (Fig. 10.2). The last metamorphism, which reached the amphibolite facies, caused the southern limit of the granulite facies t o regress northwards, leaving orthoamphibole in its wake and is associated with a third deformation which produced gentle, open, upright folds. Van Reenen and Du Toit (1977) have established metamorphic pressures and temperatures of 7.6 kb and 800°C from a study of reaction equilibria of the mineral pairs, garnet-cordierite in the pelitic rocks and diopsidehypersthene in the mafic rocks. These temperatures and pressures were attained over an area of some 3500 km2 south and west of Louis Trichardt. They regard the drop in metamorphic pressure in the second granulite metamorphism as marking the isostatic return of the previously deeply depressed Banderlierkop supracrustals. This was closely followed by generation of the Matok Granite at depth and its diapiric emplacement. Later but so far undated events are the emplacement of postorogenic, unfoliated granites and the pyroxene-rich syenites of the Schiel pluton (latter provisionally dated at >2100 Ma, J.M. Barton pers. commun., 1978. The South Marginal Zone has been cut by several linear, northeast-striking shear zones which have locally retrogressed the high temperature mineral assemblages to chlorite grade. These shears were probably generated at about 2600 Ma and are unrelated to the 2000 Ma shears in other parts of the belt. Comparison of the North and South Marginal Zones shows that they are not quite as similar as Mason (1973) has implied. Though both show evidence of multiple metamorphism to the granulite facies, contain the peculiar Matok-Porphyritic type of granite and seem t o be derived from pre-existing cratonic granite-greenstone terrane, the South Marginal Zone does not display such a linear fabric as that of the North Marginal Zone. The Matok Granite is not linear like the Porphyritic Granite and the South Marginal Zone does not possess the linear marginal granite plutons in such profusion (Tuli, Chibi, Zimbabwe) though the Mathlala, Moletsi and Lunsklip granites, lying north of the Pietersburg Schist Belt could be analogous (Fig. 10.2). This may be due to lack of the widespread shearing fabric at this (2600 Ma) stage of evolution of the South Marginal Zone, which undoubtedly influenced the emplacement of granites in the margins of the Rhodesian craton.

663 YOUNG UNMETAMORPHOSED COVER AND IGNEOUS COMPLEXES

After the Limpopo Belt was denuded t o approximately its present level, there were at least two periods of sporadic sedimentation and volcanicity, which began at about 1800 Ma, and were separated by further erosion. As a result comparatively young rocks of the Waterberg, Umkondo (Late Precambrian) and Karoo (Mesozoic) Supergroups now cover the extremities of the Limpopo Belt. To the east it passes beneath Umkondo and Karoo rocks and, in the west, it is obscured by Waterberg, Karoo and recent Kalahari cover. Two fault-controlled outliers cover critical sections along the margins of the Central Zone. The Tuli Trough, to the north, consists almost entirely of Mesozoic Karoo rocks but the Soutpansberg Trough in the south comprises a mixture of intensely faulted Waterberg and Karoo strata (Fig. 10.4). Karoo and post-Karoo igneous activity is rife throughout the Limpopo Belt with extensive basalt flows and intrusion of dolerite dykes along major Limpopo east-northeast and west-northwest trends. Early gabbroic intrusions followed by extensive granophyre sheets and later ring complexes of granite, granophyre, nordmarkite and nepheline syenite have been emplaced into the Karoo alkaline basalt-rhyolite succession at Mateke Hills and Chiwonje (Fig. 10.4; Cox et al., 1965) and into the rocks of the Limpopo Belt at Marangudzi (Gifford, 1961). Sporadic reactivation of large-scale faults of the Limpopo Belt has probably influenced the depositional environment and later preservation of Waterberg and Karoo strata and provided the necessary weaknesses for the mantle tapping conduits of their associated volcanism. ECONOMIC GEOLOGY

Significant deposits of nickel-opper (Pikwe-Selibe), copper (Messina) and chromium (Rhonda-Spinel) are worked within the Limpopo Belt. Other deposits comprise very limited amounts of a wide range of metals and minerals. Some deposits are characteristic of the rock types of the belt and others are related t o the later metamorphic and structural history.

Syngenetic deposits The occurrence of nickel-copper sulphides in a rather variable amphibolite host at Pikwe-Selebi (Fig. 10.2, Table 10.1) in Botswana, has been described by Gordon (1973). Lear (1971) suggested emplacement of early troctolitic sills with immiscible sulphides which were later subjected t o polyphase metamorphism and folding. This deformed and t o some extent re-sited the sulphides, though generally they lie within the amphibolite host (Wakefield, 1974). Though alumina and iron-rich chromitites occur sporadically in serpen-

664 tinites and meta-anorthosites throughout the Limpopo Belt, the only economic chromitites occur at the Rhonda and Spinel mines (Fig. 10.2) in Zimbabwe. They lie as small, irregular pods in tectonically disrupted and metamorphosed mafic-ultramafic intrusions, closely associated with magnetite quartzites (Worst, 1962; Robertson, 1973a). The sedimentary rocks contain small deposits of marble, dolomite and graphite and some magnetite quartzites are sufficiently iron-rich to constitute low grade iron ore deposits. Epigenetic deposits Apart from modifying syngenetic deposits, metamorphism has generated ores of its own, generally of pegmatite and vein type. Small and rare, erratic gold deposits occur in part of the North Marginal Zone of Zimbabwe in replacement bodies and quartz feldspar veins. They were probably emplaced during metamorphic retrogression (Phaup, in Odell, 1975). Small amounts of beryl, corundum, sillimanite, magnesite, talc and andalusite have also been worked. Apart from the usual gold vein deposits in the greenstone belts there is a concentration of tungsten, tantalum, niobium, tin, beryllium and lithiumbearing pegmatites spread symmetrically in the terranes flanking the Limpopo Belt. Emeralds occur at Klein Letaba in South Africa and high quality emeralds have been mined in the Mewza Schist Belt in Zimbabwe. Copper deposits of possible Karoo age near Messina in South Africa are associated with the Messina Fault (Fig. 10.5). At Campbell, Messina and Harper mines, bornite and chalcopyrite-bearing fissures and cross-cutting breccia pipes are surrounded by extensive haloes of hydrothermal alteration. The Artonvilla and Spence ore bodies are, t o a large extent, stratabound, gneisses (Sohnge, 1946; van lying in or near biotite-garnet-cordierite Graan, 1964; Jacobsen et al., 1976). Jacobsen and McCarthy (1976) and McCarthy and Jacobsen (1976) prefer an hydrothermal origin for the Artonvilla ores. SYNTHESIS

During the last decade geochronology has provided a numerical framework for correlation of igneous, metamorphic and structural events throughout the Limpopo Belt and adjoining cratons (e.g., Hickman, 1976 and Barton and Ryan, 1977). The results are summarized in Table 10.11 and need critical examination as their interpretation in ancient, polymetamorphic crust is difficult. K-Ar and Rb-Sr mineral ages do not penetrate an allpervasive 2000 Ma event in the belt and although Rb-Sr isochron techniques do so, their meaning can be ambiguous. _The most satisfactory method of unravelling the complex history of the Limpopo Belt is by coordinated struc-

665 tural, petrographic and geochronological studies of selected areas but so far such studies are few, and widely scattered. Sohnge et al. (1 48), Van Eeden et al. (1955), Jacobsen (1967), Mason (1970), Bahnemann (1971, 1972) and Key et al. (1976) have stressed that the attitudes of structural elements are remarkably constant over wide areas of the belt. Mason (1973) suggested a mechanism of deformation by a combin;?ion of convective and flow folding with interaction of stable blocks and shear belts playing a major part. He regarded the refolding as progressive and in some cases almost synchronous with stable, high-grade metamorphism and suggested that the Central Zone was mobilized separately and ground between stable granulite zones. Recent work has indicated a more complex series of events than Mason envisaged. All three zones have a large number of common structural, metamorphic and chronological elements (Table 10.111) so that separate evolution seems unlikely. The following is a tentative synthesis of recent work. The Central Zone has preserved the longest history, perhaps because it lacks the intense, very dry granulite metamorphism of the marginal zones which seems to have reset radiometric ages. The earliest so far recognizable event (1)was probably an early metamorphism of the Sand River gneisses at 3858 Ma followed by intrusion of mafic dykes at 3643 Ma (2). The errors of these events just overlap and they are broadly similar to errorchrons on ancient gneisses of the Rhodesian craton. The next event (3) was intrusion of anorthositic complexes and mafic dykes at 3221 and 3123 Ma, respectively. Depositon of the Limpopo Group, into which these complexes are intrusive, probably occurred between events 1and 3. It must be emphasised that rocks of the Limpopo Group may not be chronologically equivalent throughout the belt and the age determination errors involve large time spans. It is not yet clear when the early, high-pressure, granulite metamorphism, anatexis and folding took place in the Central and South Marginal Zones but it may have closely followed this 3200 Ma event and probably marked the initiation of the Limpopo Belt as a zone of high heat flow. An event (4) at 2930 Ma which dates metamorphism, anatexis, intrusion and folding in the North Marginal Zone, is the earliest event so far recorded there and is also reflected by granitic gneisses in the Central Zone. Hickman (1976) considered the gneisses that fall on the 2930 Ma isochron had only a brief prior crustal history. It is possible that more sampling in this zone may provide evidence of a longer crustal history in the form of high initial ratios, even if all rock units were isotopically reset by the granulite metamorphism. A complex series of events seems t o have occurred in all three zones at 2500-2700 Ma. It began with folding and high-pressure granulite metamorphism in the Central and South Marginal Zones at about 2700 Ma (5). In the North Marginal Zone there was shearing, thrusting and partial retrogression to amphibolite facies. Anatexis in all three zones produced the similar, Rapakivi-like (sensu lato), Bulai-Matok-Porphyritic “Granites”. Over-

TABLE 10.11 Selected Limpopo geochronology Event (see t ext )

Zone

1

'

Q,

a

Q,

Material

Approx. locality

Method

Age

Initial ratio

Ref.

C.Z.

Gneisses

Sand River Messina

R.S.I.

3 8 5 8 f 116

0.7012 f 0.0002

A

2

C.Z.

Early mafic dyke

Sand River Messina

R.S.I.

3643 f 102

0.7014 ? 0.0003

A

3

C.Z.

Anorthositecomplex

Messina

R.S.I.

3221C 48

0.70331 t 0.00006

A

4

4

N.M.Z. C.Z.

Gneisses Gneisses

Bangala Dam S.W. of Messina

R.S.E. R.S.E.

2930 f 120 2850 f 378

0.7014 -

5

C.Z.

Messina

R.S.I.

2690C 60

0.7038 f 0.0016

I

5

R.C.M.

Fort Victoria

R.S.I.

2660f

0.7025 C 0.0030

D

5 5 5

C.Z. C.Z. S.M.Z.

BulaiSingelele granites Victoria porphyritic granite Bulai granite Bulai granite Matok granodiorite

W. of Messina E. of Tshipise Banderlierkop

R.S.I. R.S.I. R.S.I.

2772 f 106 2751k 56 2657 f 143

6

R.C.M.

Fort Victoria

R.S.I.

2620f

80

0.7035 f 0.0080

D

6

R.C.M.

Fort Victoria

R.S.I.

2610k

60

0.7044 ? 0.0018

D

6 6

R.C.M. N.M.Z.

Chilimanzi granite Zimbabwe granite Lepidolite Retrogressed gneiss

Bikita Bangala and Ingoni

R.S.M. Rb-Sr

2 54 0-2 7 60 2650 f 150

-

G D

70

f 0.0012

D A

A A A

0.704

6 6 6

C.Z. S.M.Z. S.M.Z.

Gneiss Monazite Paragneiss

7 7

R.C.M. N.M.Z.

Great Dyke Great Dyke Satellite

? ?

C.Z. C.Z.

8

R.C.M.

8

N.M.Z.

8

C.Z.

9

10

Pikwe Banderlierkop Banderlierkop

R.S.I. U-Pb R.S.I.

2 6 6 0 f 60 2 5 5 0 k 90 2 6 3 6 k 98

0.7044 -

N.E. of Tods

R.S.I. R.S.E.

2 5 3 2 k 89 2600 -I 120

0.7024 k 0.0008 0.7049 f 0.0012

C F

Diabase dyke Gneisses

E. of Messina Losi Quarry

R.S.I. R.S.E.

2 2 6 5 f 64 2240 f 400

0.7012 0.7035

f 0.0002

A I

Chibi granite Feldspars and biotite Biotite

S.W. of Fort Victoria Tods Quarry

K-Ar (Overprint) R.S.M.

1 8 6 9 2 56

Artonvilla

R.S.M.

-

Basal Waterberg basalt

W of Tshipise

R.S.I.

1790 2 100

-

Nuanetsi Igneous Suite

R.S.M. K-Ar

200

-

2010-I 19402 1970k 19202

60 50 50 50

k 0.0008

E

B A

? 0.0047

J 0.724

H K

0.71079 f 0.00008

A

B

' Zone: Rhodesian Craton Margin (R.C.M.); North Marginal Zone (N.M.Z.); Central Zone (C.Z.); South Marginal Zone (S.M.Z.).

Method: Rb-Sr isochron (R.S.I.); Rb-Sr errochron (R.S.E.) Rb-Sr mineral age (R.S.M.). Decay constant for 8 7 R b taken a t 1.39 x yr-' . References: Barton and Ryan, 1977 ( A ) ; Burger and Coertze, 1973 (B); Davies e t al., 1970 (C); Hickman, 1976 ( D ) ; Hickman and Wakefield, 1975 (E); Robertson and Van Breemen, 1970 ( F ) ; Vail and Dodson, 1969 (G); Van Breemen, 1968 (H); Van Breemen and Dodson, 1972 (I); Wilson and Harrison, 1973 (J);Van Breemen e t al., 1966 (K).

lo-'

TABLE 10.111 Proposed sequence of events in the Limpopo Belt

*

South marginal zone

Approx. age (Ma)

thin slice isochrons and mineral ages; shearing, mylonite and flaser gneisses (P.M.); retrogression t o lower amphibolite and locally (shear zones) t o greenschist facies (P.M. ) ; dome and basin structures, annealing ( P )

shearing; n o dates available

1900-2100

-

-

2500 ( 7 )

Great Dyke

retrogression t o amphibolite facies (M.P.); large-scale recumbent and flattening cataclastic fabrics (P); upright flexslip folds tightening t o conical folds (M)

Schiel pluton? Palmietfontein granite; partial retrogression to amphibolite facies with gentle, open upright folds

undated

possibly related to but following 2600-2700 event

possible shearing

2600 (6)

Matok granite; high-pressure granulite metamorphism and anatexis

2700 (5)

amphibolite t o lowpressure granulite metamorphism with shear and fold deformation, anatexis and granite emplacement

North marginal zone

Central zone

thin slice isochrons and mineral ages, n o corresponding fabric.

intrusion of Great Dyke Satellites

marginal granites, Bikita pegmatites.

porphyritic granite. shearing, local thrusting and retrogression to amphibolite facies.

Bulai granite; near granulite (P) t o lowpressure granulite (M) metamorphism; tight upright folds (M.P.); isoclinal upright folds (M)

Overall event -

(8) isostatic uplift and development o f major sheer belts

major and minor upright folds; partial melting, charnockite intrusion; static granulite metamorphism.

possible metamorphism

-

2900 ( 4 )

? Northen granulite metamorphism

?

high-pressure granulite metamorphism; shearing of fold limbs; isoclinal recumbent folding (MI

high-pressure granulite metamorphism; uncertain relationship to two early deformations producing isoclinal, plastic, co-axial folds

undated

possibly related to 3200 event

intrusion of anorthosite complexes and concomitant metamorphism to granulite facies (M)

?

3200 ( 3 )

intrusion of anorthositic complexes

?

deposition of Limpopo Group

?

undated

initiation of Messina depository

?

intrusion of mafic dykes

?

3600 ( 2 )

?

metamorphism and anatexis of Sand River Gneisses (M)

?

3800 (1)

?

very early history deposition of Sand River rocks

?

undated

* P = Pikwe area; M = Beitbridge-Messina

area.

? early cratonic metamorphic event

670 lapping with this is event 6, centred at about 2600 Ma with further anatexis of 2900 Ma granulite material in the North Marginal Zone and emplacement of marginal granite plutons and pegmatites. These ages are broadly similar t o the ages of major Zimbabwean greenstone belts and their associated granites and were followed by intrusion of the Great Dyke and its satellites (7) which mark the end of major activity in the northern part of the belt. A further, rather obscure, event at 2200-2300 Ma involved anatexis of gneisses near Mahalapye, intrusion of aplite dykes in the Rhodesian craton margin and mafic dykes east of Messina. Though this age is similar t o that of the Ventersdorp Group volcanics of South Africa and t o parts of the Gaberones Granite of Botswana, the age determination of the Mahalapye remobilization has a large (400 Ma) error so it could alternatively be linked t o events 6 or 8. An event (8) found by most workers throughout the belt is centred at 2000 Ma and marks shearing, mylonitization and localized retrogression in the Tuli-Sabi Shear Belt. It could also include lower amphibolite facies metamorphism, annealing, folding and finally uplift and cooling of the belt. Reactivation of pre-existing basement faults, soon after, defined the Soutpansberg depository, and pillow lavas (9) dated at 1790 Ma were erupted at the base, followed by the rest of the Soutpansberg Group. Later reactivation of these and similar faults controlled deposition of Karoo rocks in the Tuli Trough, followed by eruption of basalts, intense intrusion of dykes and alkaline complexes at 200 Ma (10). Most workers agree that the polytectonic, polymetamorphic Limpopo Belt is intracratonic (ensialic). It seems probable that before about 3200 Ma the Rhodesian and Kaapvaal cratons were one but at about this time a belt of high heat flow developed and the Limpopo Belt became unstable and deviated from normal cratonic development. The first manifestation was deposition of the Limpopo Group, its burial and consequent high-pressure granulite metamorphism. This zone of weakness was later reactivated, the rocks being remetamorphosed at a relatively low pressure, and intruded by granites. At the same time the Kaapvaal craton apparently stabilized while granite-eenstone development continued in the Rhodesian craton during the latter history of the Limpopo Belt (2700-2500 Ma). As the Limpopo Belt attained rigidity its character changed from a zone of mobility t o one of shears and rifts (-2000 Ma) which were later reactivated in a minor way t o control deposition of cover rocks. Attempts t o explain the Limpopo Belt in terms of plate tectonics are unsatisfactory as the evidence is tenuous. I t may be hazardous t o interpret the early history of the Limpopo Belt in terms of modern global tectonics, though they may apply t o its later history. The early history of the belt is reflected in the deposition of the Limpopo Group which is not typical of a greenstone belt, a Proterozoic basin or a geosynclinal environment. A close link t o faulting and rifting appears throughout the latter history of the belt

671 so it is possible that proto-rifting could also have controlled its earlier history. Little is known of the three-dimensional subcrustal structure of the Limpopo Belt, although gravity evidence and the regional distribution of granulites (Fig. 10.3) both suggest a thin crust under the granulite zones. This and a progressive dying out t o the west, lends some support t o a rift hypothesis. Until there are more data any model as t o the cause of the Limpopo Belt must remain speculative.

B. The Namaqualand Metamorphic Complex

*

INTRODUCTION

When the Precambrian Research Unit started investigations into the Namaqualand Metamorphic Complex (NMC) in 1967, large tracts of the country underlain by granitic and paragneisses were geologically virtually unknown. The areas which had been geologically mapped were the eastern parts (Rogers and Du Toit, 1908; Von Backstrom, 1964), the Vioolsdrif area (Gevers et al., -1937), The Richtersveld (De Villiers and Sohnge, 1959), parts along the Orange River (Coetzee, 1941; Von Backstrom and De Villiers, 1972) and in the southwest (Brink, 1950; Jansen, 1960; Von Backstrom, 1960). Detailed unpublished geological maps were available for the copper mining district around Springbok (Benedict et al., 1964) and a number of reports were published on various unmapped areas. Martin (1965) gave a summary of the geology and discussed certain interesting aspects. In these reports structure and metamorphism were only discussed in outline, so that little was known about these aspects in 1967. Since then large areas have been mapped by various members of the Unit who also extended their activities t o include areas in Namibia and onto the Late Precambrian formations. Research was also conducted by the Department of Geology, University of the Orange Free State, and the Geological Survey continued with the regional mapping. Since the discovery of base metal deposits in Namaqualand, first at Prieska (Cu, Zn), then at Aggeneys (Cu, Pb, Zn, Ag) and Gamsberg (Zn), at the Haib (Cu) as well as the Pb-Zn ores in the Late Precambrian Gariep rocks, extensive exploration resulted. With few exceptions, all the rocks of Namaqualand returned isotopic ages of 1100 k 200 Ma (Nicolaysen, 1962; Nicolaysen and Burger, 1965; Burger and Coertze, 1973), but in the eastern Richtersveld ages ranging from 2000 t o 1800 Ma have been determined (De Villiers and Burger, 1967; Reid, 1976). In the east where the NMC abutts against the Kaapvaal Craton ages ranging up t o 2960 Ma have been found.

* By P. Joubert.

672

Fig. 10.6. Generalized map of the main geological elements of southern Africa showing the distribution of the Namaqualand Metamorphic Complex, the extensions in Natal and Namibia and the pre-Damaran rocks, part of which can possibly also be related to the NMC. Area enclosed by thick line is location of Fig. 10.7.

In any area of complex geology it is to be expected that there will be several interpretations of the geology and Namaqualand is no exception. There are disagreements on many aspects such as regional correlation, structure, metamorphism and on the interpretation of geochronological results. The area underlain by the NMC in the northern Cape Province is crescentshaped forming a strip in excess of 150 km wide from the west coast t o the western edge of the Kaapvaal Craton, a distance of sme 480 km. In Namibia the rocks of the NMC are exposed along a narrow belt stretching northeast from Ai-Ais and a large area t o the east of Luderitz. Exposures of these gneisses are extremely good in the belt along the Orange River and along the escarpment west of Springbok, but poor towards the south as the Karoo cover is approached (Figs. 10.6-10.8). LITHOLOGY

Geological sequences For descriptive purposes, the rocks of the NMC will be divided into five regions; Namaqualand (the western and central parts), eastern Bushmanland (the rocks of the Kheis belt), the craton-mobile belt junction, the Richtersveld (the area underlain by rocks of the Vioolsdrif Igneous Province) and the extension into Namibia.

673

Namaqualand The simplified lithological sequence deciphered in the western and central parts of Namaqualand is as follows: Western Namaqualand grey fine-grained biotite gneisses with mafic bands 5. quartzites, sometimes with thin crystalline limestones 4. aluminous gneisses and schists 3. thin quartzo-feldspathic gneiss 2. pink gneiss, underlain by mafic gneiss in places 1. augen gneisses 6.

Bushmanland Sequence

Springbok

Nousees mafic gneiss

(? grey granulite)

Pella quartzite

Springbok quartzite

Namies schist absent Haramoep gneiss

Springbok schist Springbok granulite absent

Aroams gneiss

Nababeep (Brandberg) gneiss

Joubert, 1 9 7 1 Joubert, 1 9 7 4 b Benedict e t al., 1 9 6 4

There is still disagreement on the correlations as outlined above as the geologists of the O’okiep Copper Company, who have the most detailed information of the area, do not consider the repeated lithologies or groups of associated lithologies as resulting from isoclinal folding. Moore (1976) describes a similar sequence for an area south of Pofadder, but equates the grey biotite gneisses with the augen gneisses and places them at the base of the succession. The gneisses of the NMC are not only banded in outcrop, but are also seen as belts of grey mafic rocks separated by layers of acidic and granitic gneisses giving rise to banding on a broad scale. The grey biotite gneisses and associated rock types lie along remarkably uniform bands and display notably consistent lithologies over most of Namaqualand and Bushmanland. Bodies of gabbro and norite, occurring along the mafic bands, are not found in the adjoining acidic gneisses. The main rock type comprising the mafic belts is a fine-grained grey gneiss with streaks and lenses of quartzo-feldspathic gneisses and amphibolites, the latter being represented by pyroxene granulites in areas which have undergone granulite facies metamorphism. These belts are frequently edged by distinctive lithologies such as the common association of quartzites and aluminous gneisses, sometimes with crystalline dolomites and other characteristic rock types. Calc-silicate rocks are commonly found within the grey gneisses and they attain their greatest distribution west of Kenhardt. Although the quartzites are relatively thin, they can often be followed for many kilometres along strike over most of Namaqualand and Bushmanland. These rocks become thinner towards the east, but

674 are most extensively distributed along the Kheis belt. Between these belts of supracrustal rocks, the main rock type is a granitic augen gneiss, known as the Nababeep Gneiss in the west, but with the so-called pink gneiss, a leucocratic, red-weathering quartzo-feldspathic rock, also prominent in some parts of Namaqualand.

Eastern Bushmanland Originally Rogers and Du Toit (1908) mapped the belt of rocks stretching in a northwestsoutheast direction from west of Prieska to beyond Upington as the Kheis “System”, consisting of the Marydale beds comprised of lavas and minor sediments at the base, overlain by the Kaaien quartzites and the Wilgenhoutdrif greenstones with intercalated sediments at the top. Vajner (1974) further subdivided the Kaaien quartzites and agreed with Rogers and Du Toit (1908) that the whole of the Kheis Group is older than the sedimentary and volcanic rocks of the Kaapvaal Craton. His subdivision is as follows:

Craton

Matsap Formation Ongeluk Volcanics Koegas Formation Asbestos Hills Iron Formation Campbell Rand Formation Black Reef Formation Seekoebaard Formation

I

Waterberg Group Middle Griquatown Group Lower Griquatown Group

Ventersdorp Group

Skalkseput Granite (2500 Ma) Draghoender Granite (2960 Ma)

Wilgenhoutdrif Formation Kaaien Formation Groblershoop Formation Marydale Formation

i

Transvaal Supergroup

i

Kheis Group

The conclusion that the rocks of the Kheis Group are older than those of the Transvaal Supergroup was already being questioned in 1910 (Schwarz) and, in view of the isotopic dating, Nicolaysen (1962) advocated reconsideration of the stratigraphic position of the Kheis rocks. More recent work of the Geological Survey in that area (Smit, 1973) indicates that the Kaaien Formation is the southerly extension of the upper part of the Matsap Formation, a view which is now generally accepted. The other members of the Kheis Group are also correlated with various formations lying on the craton and it is of interest t o note that part of the Marydale Formation is now equated with the volcanic rocks of the Matsap Formation dated at 2070 ?r 90 Ma and the rest with the Seekoebaard Formation.

675 The northern extensions of the Kheis rocks, consisting of a group of “reconstituted sedimentary rocks” (quartz-sericite schists, quartzites, quartzo-feldspathic rocks, calc-silicate rocks, amphibolites, aluminous gneisses and schists) and metalavas, were mapped by Von Backstrom (1964) and correlated with the Kaaien and Marydale Formations. Botha et al. (1976) renamed many of the lithological units there and they extended the correlation t o the west and equated rocks lying along the Kheis belt with those occurring within the NMC. The detailed investigations now being carried out by the Geological Survey should resolve the various contradictions.

Vioolsdrif area To the north of the Namaqualand gneisses an area, underlain by rocks belonging to an intrusive suite emplaced into lavas and minor sediments metamorphosed to greenschist facies, yields isotopic ages ranging from 2000-1800 Ma. The lavas and sediments, originally correlated with the rocks of the Kheis Group (De Villiers and Sohnge, 1959) and now known as the Vioolsdrif Igneous Province, have been subdivided by Blignault (1974) as follows : Orange River Group

Haib Subgroup Rosyntjieberg Formation De Hoop Subgroup

Reid (1976) lists the rocks of the Vioolsdrif Intrusive Suite as follows in order of relative age: (Youngest)

(Oldest)

Leucogranite Adamellite Granodiorite Tonalite Diorite Gabbro-peridotite

The relationship between the rocks of the Vioolsdrif Igneous Province and those of the NMC has been controversial for many years as the zone along which they meet is poorly exposed and has never been carefully studied. Gevers et al. (1937) recorded that the “Namaqualand Granite and Gneiss” are younger and intrusive into the granites of Vioolsdrif, an observation with which most later investigators agree. The relationships between the lavas of the Orange River Group and the rocks of the Bushmanland Sequence, however, are still uncertain. Gevers et al. (1937) grouped the lavas with quartzites as forming part of the Kheis without indicating their relative stratigraphic positions, while more recently Bertrand (1976) suggested that the quartzitic rocks represent a facies change away from the volcanic centre where the

677 lavas were extruded. The occurrence of sillimanite in rocks typical of the Bushmanland Sequence in localities far west of the position of sillimanite inversion as determined by Beukes (1973) seems to indicate that the quartzites and schists underwent metamorphism prior t o the extrusion of the Orange River Group lavas. There can, however, be no doubt that the rocks of the Vioolsdrif area become involved in deformation t o the north, east and south of the Richtersveld and are, in many respects, indistinguishable from some of the lithologies found within the Bushmanland Sequence. Martin (1965, plate 1) distinguished between areas of “Namaqualand granite-gneiss including Kheis metarocks” and the “Kheis metarocks intruded by and metamorphosed t o Grey Gneiss”. The term Grey Gneiss here refers to the Vioolsdrif granodiorite. The contact between these areas is shown to extend south of and following the course of the Orange River and recently this contact was again reconsidered (Joubert, 1976) and the correlation between the rocks of Vioolsdrif and those of the Kheis belt reiterated in spite of the fact that the lavas of Vioolsdrif are calc-alkaline in character (Reid, 1976) and those of the Kheis tholeiitic (Vajner, 1974). It is thus possible that the rocks now grouped as belonging to the NMC may be divided into a portion underlain by rocks typical of the Bushmanland Sequence where an east:west trend is most common and a zone of rocks typical of the Kheis belt which runs north from west of Prieska, along and north of the Orange River and into southern Namibia where the general strike is in a northwesterly direction (Fig. 10.7).

Namibia Toogood (1976) divided the rocks occurring north of the Orange River and Pofadder into pretectonic and post-tectonic groups with the former consisting of quartzo-feldspathic, biotite, aluminous, quartz-muscovite and hornblende gneisses as well as calc-silicate rocks, while the latter group is comprised of various granitic gneisses, norites and amphibolites. Around AUS,Jackson (1976) also described his Garub Sequence, consisting of impure marbles, calc-silicate rocks, aluminous gneisses, biotite schists, biotitehornblende schists, quartz-feldspar rocks, quartzites, magnetite and hematitebearing rocks, amphibolites, pyroxene granulites as well as cummingtonite and chlorite schists, as pretectonic. These rocks were intruded by pretectonic tonalites, followed by syntectonic intrusions, mainly of granitic composition. Other granites, granodiorites and diorites are listed as late syntectonic, while the mainly post-tectonic granites are correlated with the intrusives associated with the Sinclair Group. All the rocks, except those correlated with the Sinclair, are grouped with the NMC. To the west of the area investigated by Jackson, McDaid (1976) also described the occurrence of sediments and lavas which are tentatively correlated with rocks of the Sinclair Group. Granodiorites very similar t o those of the Vioolsdrif are have been found in the area of Luderitz.

678

Craton-mobile belt junction Judging by the literature, the position of the line of contact between the Kaapvaal craton and the mobile belt has not been settled, Most authors accept that the western edge of the granites immediately east of Marydale approximates the western limit of the craton. Kroner and Blignault (1976), however, showed this line a t least 15 km east of the town of Marydale while Vajner (1974) extended the cratonic area t o include the Kaaien quartzites as far as his “Namaqua front”, which is a line running west of the Kaaien Hills and Upington. Vajner (1974) thus included the rocks of the Kheis belt with those of the craton and described the Namaqua front as a “tectonic line or a zone up t o several kilometres wide which appears at different places along its length as a metamorphic transition, as an oblique-slip fault and as a general interface between areas yielding radiometric ages of 2.5 t o 2.9 Ga and 0.9 t o 1.25 Ga.” Kroner and Blignault (1976) regard the southern extension of their “Kheis Province” as separating the craton (“Kaapvaal Province”) from the “Namaqua Province” (NMC) and include the old granites dated at 2920 and 2650 Ma with the rocks of the Kheis belt. Vajner (1974) mapped small hills composed of dolomite rocks, 25 km north-northwest of Marydale, as outliers of the Campbell Rand Formation of the Transvaal Supergroup although Rogers and Du Toit (1908) included these rocks with those of the Kheis. The dolomites overlie metavolcanics and have been intruded by biotite granodiorite, which Vajner (1974) tentatively correlated with the 1000 Ma metamorphic event. The dolomites are separated by a major fault from the Kaaien quartzites just t o the west, but they have also been deformed by a t least two phases of deformation resulting in folds similar in style and direction to those seen in the quartzites. It seems clear that these rocks of the Transvaal Supergroup reached beyond the confines of the craton and were also involved in the tectogenesis of the mobile belt. This same feature can be seen further north where the rocks of the Matsap overlap to the northeast and become more deformed away from the contact and where they have been taken t o represent the Groblershoop Formation. Several large shear zones transect the granites as Marydale is approached, but a rather sudden development of a strong pervasive foliation occurs only at the change of lithology where the amphibolites, previously grouped with the Marydale Formation, appear. There does not appear t o be any doubt that the edge of the craton is marked by the western limit of the granite there and the presence of younger granodiorite t o the north of Marydale confirms this view. The 2960 Ma age for a granite t o the south of Prieska and west of the Doornberg Fault (Fig. 10.8) possibly indicates that this fault, usually considered as marking the western edge of the craton there, cuts across part of the craton which actually extends further west than the position shown in Fig. 10.8.

683 Intrusive rocks Apart from the intrusive rocks mentioned above, various intrusions in the shape of phacoliths, plutons, plugs, sills and dykes of compositions ranging from acid to ultramafic have been emplaced into the rocks of the NMC.

Granites Granites of the 100 M a event. Granites all yielding ages of 1100 ? 200 Ma have been emplaced ih the gneisses of the NMC, but there are some localities around Pofadder, for instance, where their distribution is limited t o a small number of minor intrusions. In most regions, however, they are of widespread occurrence and are known collectively as the Concordia type in the west and as the Straussberg Granites (previously Grey Gneiss) in the east, but a number of other names have also been employed in other areas. The distribution of these granites is irregular and they are frequently found as dyke or sheet-like bodies intruded along the foliation planes and banding of the host rocks. In many areas numerous small bosses occur while some granite bodies are accommodated in mild folds of late generation. Pegmatitic phases of the granite are emplaced in lit-par-lit fashion and result in pinch-and swell-structures. Contacts are sharp or gradational, but many are marked by the presence of aplites or pegmatites while tongues of the host rocks reach into the granite as tapering ribbons disappearing into clusters of garnets. The growth of porphyroblasts in mafic rocks intruded by the granites is sometimes quite striking, while recrystallization of the intruded paragneisses results in the destruction of the banding. The granites frequently enclose xenoliths and in these, the linear and planar elements maintain parallelism with those of the country rock. This feature, as well as the fact that thermal contact phenomena are absent, indicate that the granites were not forcibly injected and that the temperature of the invaded rocks must have been as high, or nearly as high, as that of the invading granite at the time of emplacement. The Concordia granitic gneiss is a white t o pinkish rock which, when the intrusive bodies attain a fair size, crops out as large smooth boulders, recognizable even from a distance. There is an apparent variation in the amount of mafic constituents in the various granites of this group and when they are finer grained, are indistinguishable from some of the paragneisses. Some varieties are porphyroblastic, but normally the granites consist of quartz, potassium feldspar and plagioclase as essential minerals, with minor amounts of biotite, ore, apatite, zircon and sphene. Hornblende and garnet are only sometimes present. The features displayed by these granites are generally interpreted as being consistent with late syntectonic emplacement. All gradations between leucocratic granitic gneisses t o darker and more rarely to mafic gneisses in which porphyroblasts grew, to schlieren of

684 country rock in the granites and homogeneous gneisses are t o be found. Migmatitic phases, as well as primary gneissic structure in the granite, merging into areas of homogeneous granite in which no directional fabric can be discerned, are features commonly associated with these rocks. It is, therefore, tempting to ascribe a replacement origin, with local mobilization only, to all these granites and undoubtedly a large proportion of the granitic gneisses evidently formed by coarse recrystallization. The occurrence, however, of metasediments essentially of the same composition as the granites, even in areas intensely intruded by granite and showing only local patches of recrystallization, as well as transgressive contacts in places, indicate that magmatic intrusion did occur on a fairly large scale.

Later granites. De Villiers and Sohnge (1959) described the Richtersveld Igneous Complex as a series of intrusives, occurring in the eastern Richtersveld, ranging in the composition from granites t o akerites and divided them into syenite and granite suites. Rocks from this group have been dated at 930 f 30 and 680 f 40 Ma (U-Pb, Burger and Coertze, 1973). A small body of granite (Lekkersing) found further south and intrusive into rocks of Gariep Group, probably also belongs to the Richtersveld Igneous Complex and has been dated as (Pb-Pb) k 800 Ma. The intrusions along the Kuboos-Bremen line were first described by Sohnge and de Villiers (1948) as the Swartbank-Frankfurt plutonic line along which the Swartbank, Kuboos and Tatasberg plutons were emplaced. Martin (1965) extended this east-northeast line t o include the Bremen Complex. The Swartbank mass consists of a coarse porphyritic granite,while the Kuboos pluton is composed of an earlier composite stock of granite and syenite enclosed by a later and much larger body of coarse porphyritic granite. The U-Pb age of a sample from Kuboos was determined as 550 f 20 Ma (De Villiers and Burger, 1967). The Tatasberg plutons occur as a composite stock of coarse porphyritic granite enclosing an older mediumgrained granite and another consisting of medium- to fine-grained granite only. The Bremen Complex has been shown t o consist of two distinct igneous events separated by the intrusion of the Gannakouriep dyke swarm. The plutonic rocks are represented by syenites, quartz syenites and granites and the younger phase was found t o be intrusive in rocks of the Nama Group. Several other smaller occurrences of igneous rocks have been found along this zone. Charnochitic adamellites Large areas around Keimoes on both sides of the Orange River are underlain by charnockitic adamellite porphyry. They appear to have been intruded as phacoliths, largely restricted to anticlinal and synclinal structures and although the adamellite is locally transgressive, they are regionally concordant and in structural harmony with the gneissosity and schistosity of the

685

,............_...,_.'

-

31O

Fig. 10.9. The metamorphic zoning in the Namaqudand Metamorphic Complex (see text), the zone along which pegmatites are concentrated and the distribution of some minerals. A = andalusite; K = kyanite;S = sapphirine; W = wollastonite.

invaded rocks. The contacts are thus concordant, highly transgressive or oblique to the foliation of the host rocks. Contact phenomena are not always present, but have been recognized in places as reaction aureoles, production of hybrid rocks or in the development of migmatites. The presence of sheets, sills and dykes of the adamellite proves the intrusive nature of these rocks and xenoliths of relatively small size occur throughout rocks of this group. One Pb-Pb age of +1100 Ma for the charnockite is available. A large massif of charnockitic composition also occurs in Namibia (Fig. 10.9) where the rock has a homogeneous appearance with a poorly developed fabric throughout. The normal charnockitic rocks consist of quartz, plagioclase, K-feldspar, ortho- and clinopyroxene, hornblende and biotite with accessory ore, sphene, apatite and zircon. The Namibian charnockites do not contain clinopyroxene and the K-feldspars are orthoclase and orthoclase-microperthite.

The intrusions of gabbro and norite, described below, occur as bodies in areas where the rocks have been subjected t o a grade of metamorphism no higher than upper amphibolite facies, as indicated by the mineral parageneses found in the surrounding mafic and pelitic rocks. There are, however, many rock types of charnockitic affinity in the area of granulite facies around Springbok and further south.

Mafic intrusives Numerous mafic intrusives have been emplaced along the major shear zones and although such bodies occur throughout the NMC, it is only along the major dislocations, with one small known exception, that ultramafic rocks have been found. Most of the smaller bodies have now been converted to amphibolites due t o continued movement along the shear zones. The best-known intrusive occurring along the Pofadder Lineament is that of Tantalite Valley in Namibia, described as a body of approximately concentric peridotitic-gabbroic intrusions, partly affected by metamorphism and deformation. Elsewhere along the shear zone there are also norite bodies as well as several small serpentinite occurrences. It is recorded that, especially in the case of the Tantalite Valley Complex, the situation of mafic bodies along the lineament is fortuitous because other large gabbroic bodies are frequently found elsewhere in the NMC and their emplacement predates the initiation of movement along the major crustal fractures. There are, however, two reasons for including the emplacement of some mafic bodies as part of the 1000 Ma event; these are firstly, that ultramafic rocks are found almost exclusively along major dislocations in the NMC and secondly, the concentration of small, medium-sized t o large gabbroic and amphibolites along the shear zones. Many bodies belonging t o the copper-bearing suite of rocks of the Springbok area, ranging in composition from anorthosite t o hypersthenite, are situated along so-called “steep-structures” which are related t o the deformational phase resulting in open folds. Whatever the origin of these rocks, they must have been placed in their present position during the 1000 Ma event. South of Kakamas and along the zone where rocks of the Kheis have been thrust over and onto the porphyroblastic gneisses of the NMC, several mainly lenticular bodies of norite have been emplaced along the fractures, but some of these have been retrogressed by continued movement along the thrust planes. Several ultramafic bodies, however, occur along the Kheis belt.

Dyke swarms Pretectonic, syntectonic and post-tectonic mafic and felsic dykes are encountered throughout the NMC. Dolerite dykes striking east-west occur sparsely but over a wide area and these have been correlated with the Karoo vulcanicity . Two dyke swarms have been found in the NMC. The one occurs in a wide-

spread zone along coastal Namaqualand stretching t o the north far into Namibia. The other occurs south of Kakamas and, as the rocks are poorly exposed and the dyke rocks all converted t o amphibolites, they have not been recognized as occurring in a swarm before.

Western Namaqualand. A wide variety of dykes ranging in composition from gabbro, perknite, quartz diorite, lamprophyre, quartz syenite, syenite t o bostonite occur over a wide area in westem Namaqualand. They generally strike north-northwest and north-south, but sometimes follow a northeasterly direction and were emplaced prior t o the deposition of the rocks of the Nama Group (k600 Ma). The K-Ar age determinations carried out on these rocks have given rather inconsistent results and the oldest date recorded is 878 41 Ma. Martin (1965) relates the dykes t o the Richtersveld Igneous Complex but their occurrence 250 km south of the Richtersveld and almost 600 km t o the north where they disappear under the Damara, negates this conclusion. It is almost certain, however, that the dykes were emplaced over a fairly lengthy period of time. The emplacement of the dykes was structurally controlled and was linked to the shearing deformation in the coastal region. Since their emplacement, many of the dykes have been sheared and converted to amphibolites, while others have been found with garnets forming idioblasts up t o one cm across. The dykes are vertical or have steep westerly dips, but occasionally small sills are encountered. In the east of the swarm, they are usually less than one metre wide, but towards the west they increase considerably in size and number. They are, however, absent near the coast. Even the thin lamprophyre dykes tend to extend over long distances and one of these dykes crops out intermittently over a distance of 130 km.

*

Kakamas area. Little is known about the dyke swarm in the east, not only due to poor exposure, but also because they have not previously been recognized as dykes. In most areas they strike parallel t o the regional banding, paralleling also the amphibolitic bands in the NMC and, as they have all been altered to amphibolite, the difference is not immediately evident. Many of the dykes are mildly curved by dextral movement along the zone where they are found and lines of boulders indicating the trend of the dykes on the surface define open Z shapes. The rocks are typical plagioclase amphibolites, sometimes garnetiferous, and could originally have been gabbroic in composition as they appear t o be very similar t o the sheared portions of the gabbroic bodies occurring in the same region. Pegma tites It is important t o note that, although pegmatites occur commonly throughout the NMC, they are most prolific in a belt largely straddling the contact between the Kheis and the NMC further west and the zone along the

688 Orange River, (Fig. 10.9) where they are known t o contain beryl, chrysoberyl, tourmaline, topaz, columbite-tantalite, spodumene, fluorite and rareearth minerals, amongst others. It is recorded that the majority strike westnorthwest, i.e. parallel to the main direction of shearing and the Pofadder Lineament, but there are also very many emplaced in a northwestern direction along the late fold direction, while in the Pofadder area, as elsewhere, they often parallel the axial traces of the large open folds. The vast majority of pegmatites dated yield ages of approximately 1000 Ma and must date the development of the structures along which they have been emplaced.

Kim berlites and melilite basalts Kimberlites and melilite basalt occurrences are found in profusion some 60 to 70 km southeast of Springbok, forming part of a north-northeasttrending line, stretching from Aggeneys to south of Garies, along which kimberlite and melilite basalt pipes are to be seen. Rogers and Du Toit (1908) also record the presence of kimberlites west of Prieska. The kimberlites of Namaqualand confirm the view that these rocks are nondiamondiferous in the more youthful orogenic belts. Younger formations The gneisses of the NMC are unconformably overlain by rocks belonging to the Sinclair, Gariep, Damara and Nama Groups, the Koras Formation as well as the Dwyka Formation of the Karoo Supergroup.

The 1200 Ma-old formations Koras Formation. Several large patches of almost undeformed sediments and lavas overlie the rocks of the Kheis Group and the granites intruded into them, with clear angular unconformity t o the east of Upington (Figs. 10.7 and 10.8). The sediments, derived from Kheis and Transvaal rocks, consist of quartzites, grits, conglomerates and shale and are intercalated with tuffs, agglomerates and lavas composed of quartz porphyry and basaltic andesite. Movements along north-south faults occurred while the Koras r w k s were being deposited so that unconformities are found within the sequence. The rocks have been faulted and mildly folded in a northwestsoutheast direction since deposition. Rather surprisingly these rocks, of which an acid lava was dated at 1280 50, in places lie on granite similar to others dated at +1200 Ma. Sinclair Group. The Sinclair Group of Namibia (Figs. 10.7 and 10.8) consists of coarse conglomerates, quartzites, shales and slates, intercalated with felsic and basic lavas and tuffs which accumulated in a fault-bounded trough. According t o Watters (1974) the accumulation of the Sinclair Group

689 occurred in three cycles, each initiated by basic magmatic activity, followed by felsic magmatism, vertical tectonism and the formation of fault troughs where clastic detritus was deposited locally. Watters (1974) correlated these rocks with volcanic, volcaniclastic and plutonic rock types occurring along an extensive curvilinear belt and suggests that this feature represents an ancient magmatic arc of active consumption of an oceanic plate during Late Precambrian times. Outliers of similar rocks have recently been mapped in an area to the east of Luderitz (McDaid, 1976) and have tentatively been correlated as belonging to the Sinclair Group.

The Gariep and Damara Groups Along the west coast of southern Africa a sequence of sediments of Late Precambrian age and metamorphosed to greenschist facies, encroaches upon the land in a broad sweep from south of Luderitz to Kleinzee in the south (Figs. 10.7 and 10.8). There is a gradual increase in the intensity of deformation from east t o west (Kroner, 1974) and the age of +800 Ma for a granite intruded into the Gariep rocks defines a minimum age for this group. This deformation took place during the Pan-African Orogeny of Late Precambrian to Early Palaeozoic age and was followed by the emplacement of the Kuboos Complex 550 t o 500 Ma ago. Kroner (1974) described a possible plate tectonic model for the orogeny whereby initial continental separation was followed by plate collision. The northeast-trending Damara orogenic belt appears from under the sands of the Namib Desert further north and consists of a great variety of metasediments. It has also been suggested that a plate tectonic model could be applied to the southern eugeosynclinal part of the Damara (Hartnady, 1975). The Nama Group The exact age of this generally flat-lying sequence of quartzites, shales and limestones unconformably overlying the NMC, is not known, but it is recorded (Germs, 1972) that the Nama Group is older than 510 Ma and younger than 719 f 28 Ma. These rocks cover an extensive area in the southern part of Namibia and the adjoining parts of the northern Cape Province (Fig. 10.6). They are generally unfolded except in some peripheral areas and it is interesting to note that they have been involved in the north-south shearing of the southeastern Richtersveld. Their correlation with the rocks of the Damara has been a point of controversy as these two formations have not been found in direct contact and Martin (1965) discusses three possibilities for NamaDamara relationship.

690 Karoo rocks North of the Vioolsdrif a large patch of Dwyka Tillite overlies the basement rocks and transgresses onto rocks of the Nama Group t o the northeast. The main Karoo basin overlies the Namaqualand gneisses in the south and numerous small outliers of Karoo rocks occur in many localities in southern Namaqualand and Bushmanland. Rocks correlated with the NMC crop out from under the Karoo formations on the east coast in Natal (Fig. 10.6). STRUCTURE

Four main Dhases of deformation have been described for Namaqualand (Joubert, 197413; Clifford et al., 1975) and for the Kheis belt (Vajner, 1974). In Namibia, however, a greater number of deformational events have been recognized and in the Aus area, Jackson (1976) distinguished no less than seven different foliations. Several directions of shearing are present, the most prominent being WNW-ESE and NW-SE, but along the west coast shearing in a NNE-SSW direction is most prevalent. The deformation suffered by the NMC is similar to that of gneissic terranes elsewhere and as the sequence of events recognized in Namaqualand has been substantiated by the subsequent work of a number of geologists and by the information gained on drilling of ore deposits, it will form the framework of the discussion below.

Structural events The earliest deformation (F,) can be seen in the consistently discontinuous banding and the tight intrafolial folds with disjointed limbs and sharp hinge zones deformed by all later folds. Although the existence of an early deformational phase has been questioned, there can hardly be any doubt of its presence as its effects have been confirmed in the three-dimensional studies of the ore bodies at Aggeneys.

The main deformation (F,) resulted in a variety of fold styles, but can usually be recognized as recumbent isoclinal folds on all scales. Sometimes the folds are open, but with the dominant foliation passing through the hinge zones and parallel to both or one of the limbs of the fold. The axes of the folds are parallel t o the dominant mineral lineation and such folds have been recognized in all parts of the NMC. The conclusion that these F2 folds resulted from a single tectonic event affecting the whole of the NMC is certainly questionable, but it is clear that if the cycles of metamorphism are to be correlated over the whole of the area, then it should also apply to the phases of deformation. The F2 deformation occurred at the time when the rocks were subjected t o the highest degree of metamorphism as minerals formed during this event crystallized along the direction of the dominant lineation.

691 In western Namaqualand the original trend of the axial planes of the F, folds, prior to subsequent deformation, was apparently in a WNW-ESE to NW-SE direction and, where the direction of translation can be determined, the minor F, folds indicate movement in a southerly to southwesterly direction over and onto the foreland to the south.

The third episode of deformation is the most easily recognized of all in being responsible for the formation of large open, generally east-westtrending folds and brachy-structures. The folds are often very nearly co-axial with F2 structures, but local variations result in curved 1, lineations. The F3 folds are of a flexural-slip type and much movement and shearing along the existing foliation planes during folding occurred. Sometimes a new foliation oblique to the existing s, foliation is developed as well as flexuralslip thrusts and tectonic slides along which pegmatite veins were emplaced. The large folds tend to die out in depth or upwards by differential movement along the foliation and "drag"-folds are common. Kink-folds of F3 generation are most commonly developed in the schists. Pegmatites are frequently emplaced along the axial planes of these folds and along the shear zones associated with this deformation. The fourth episode of deformation. resulted in minoclinal folds, shear zones and fractures striking northwest-southeast. Pegmatites and quartz veins are commonly associated with these structures, but the most interesting aspect of the monoclinal folds is that they have their shorter, steeper limb to the west in western Namaqualand and in eastern Bushmanland while the steeper limb is on the eastern side in the intervening area. Shear zones and tectonic slides have been mapped in all parts of Namaqualand. They vary greatly in size and extent of displacement as well as in mineral assemblages found along them. Large shear belts of consistent strike directions traverse the rocks and, although it can be demonstrated that the shear belts cut one another so that there appears to be a sequence in the order of the different directions, it is clear that the main movements associated with shearing were effected in all parts of Namaqualand at roughly the same time and that the shearing is related in some way t o the 1000 Ma event. The following directions of shearing have been measured in the NMC: (a) Coastal plains in the west

(b) Escarpment

N-S NNE-SSW (025") NNW-SSE (165") NE-SW (050") N-S NNE-SSW (025")

692

(c) Namaqualand (d) Bushmanland (east of the line through Kakamas and Kenhardt) (e) Namibia

WNW-ESE (115") ENE-WSW (065-080") NNW-SSE (160") WNW-ESE (115") NNE-SSW (025") NW-SE (120-140")

In the area of coastal Namaqualand, left-lateral shearing in a north-northeasterly direction is most common and is associated with conjugate shears in a north-northwesterly direction. The most pronounced shearing occurred along the Steenbok Shear Zone (Fig. 10.8) which can be followed from west of Kleinzee to the Richtersveld over a distance of more than 75 km and changes direction along its length in accordance with the directions of the conjugate shears of coastal Namaqualand. The shear zones are nearly always vertical or steeply dipping and the rocks produced along them are usually silvery phyllonites. All stages between unsheared rocks and phyllonites are encountered and earlier foliations and folds have been deformed in a variety of styles. The strong lineation associated with the shearing is seen as elongated mineral clusters or rodding and fluting of quartzites, usually plunging steeply down the dip and mildly curving in most zones. In belts of strong shearing where the closely spaced planes of slip produced paper-thin laminations, the lineations are seen to vary slightly in angle of plunge from one slip-plane to the next. In some places the phyllonites formed by shearing are crenulated along gently north-dipping planes resulting in well-developed kink-banding. Further deformation of the vertical shear foliation appears t o have been caused by continued and pulsating movements along the shear zone so that the newly developed foliation was further deformed. The general swing in the strike of the banding of the Namaqualand gneisses from east-west t o north-northeast is noticeable south of the Richtersveld and is most pronounced in the area west of Bitterfontein. This curvature of the banding and the dominant foliation trend was brought about by the large-scale sinistral movement along the shear zones in the coastal area. The fracturing of the coastal area was responsible for the formation of the trough in which the sediments of the Gariep Group accumulated as well as providing the channels for the dyke swarm occurring there. It can be demonstrated that the main shearing episode was terminated prior to the deposition of the Nama rocks, but that further movement also affected those rocks along the northern extension of the Steenbok Shear Zone. The largest shear zone mapped in Namaqualand up to the present is the west-northwesterly trending Pofadder Lineament (Fig. 10.8) which can be followed intermittently from east of Luderitz in Namibia, passing north of Pofadder and although not continuous further t o the southeast, the strike of the gneisses still swings to the north there in accordance with the dextral

693 drag along this lineament. Southeast of Pofadder another pronounced shear zone was mapped with small ultramafic bodies occurring along it and with the same strike direction. The Pofadder Lineament (or Tantalite Valley Mylonite Belt as it is known in Namibia), is between two and eight km wide, at least 550 km long and it has been estimated that the lateral displacement is in excess of 86 km. Numerous large and small mafic and ultramafic bodies have been emplaced along the Pofadder Lineament, many of which have been converted to amphibolites as a result of continued movement along the lineament. Some 200 km southwest of the Pofadder Lineament another major zone of strikeslip movement, followed for a distance of some 45 km, traverses the NMC immediately north of Garies. These wrench faults divide the main area in Namaqualand into west-northwesterly trending segments, but between these major shear belts there are numerous smaller parallel zones of slip with a similar sense of displacement so that the effects of large-scale dextral movements are to be seen throughout the area. Open folds cross the lineaments and these have curvilinear axial traces due to the drag along the shear zones. As the folds are of the same style as the F3 folds described above, there appears to be little doubt that they are of the same generation and have formed as a result of the movements connected with the wrench faulting. This is also indicated by the en-hchelon arrangement of some of the large open structures in parts of Namaqualand. Evidence for rejuvenation of shearing along existing zones is obvious in many of the shear belts. It is perhaps most obvious in the shearing of dykes emplaced along the shear zones, but can also be proved by the distortion and rupturing of folds, lenses, sheets of pegmatite and quartz veins formed along the belt of shearing. In some cases, mineral parageneses indicating a lower grade of metamorphism, normally found along a zone, are encountered. Tectonic slides display exactly the same features as the shear zones, but instead of the rocks being refoliated along vertical zones, the planes of slip in tectonic sliding are gently inclined to the north and the surfaces along which movement took place are gently curved. Tectonic sliding was only found associated with the F3 phase of deformation and the planes of schistosity derived from sliding are parallel, or nearly parallel to the preexisting layering. Shearing in an east-northeasterly direction is complementary to the major west-northwesterly trending zones and usually appears as highly foliated and “flattened” zones, often taking the place of antiforms between east-westtrending synforms. The movement along these shear zones is most commonly sinistral, but dextral displacements have been recorded. These structures are related to the so-called steep structures of the Okiep area (Benedict et al., 1964) along which the copper-bearing ore bodies are frequently found. The northwesterly trending faults across the Pofadder Lineament, north of Pofadder, are tension fractures and should according to the model, have

694 mainly vertical displacements, yet most of these fractures display pronounced dextral drag. In the eastern sector east of a line between Kakamas and Kenhardt, the general strike of the gneisses assumes a consistent northwest-southeast trend. This is an ancient zone along which dextral movement took place over an extended period of time which continued at least until the close of F4 times. In addition, the axial traces of the major F2 and F3 folds parallel this trend which is also followed by the numerous shear zones along the Kheis belt. On approaching this belt from the west, the general east-west banding and F3 axial traces swing parallel to the northwest-southeast trend while the l2 lineations curve into a northerly direction, with a concomitant steepening in the angle of plunge. This is a zone of convergent wrenching as indicated by the reverse faulting and overfolding to the west. The main shear zones in Namibia follow a northwest-southeast trend and as mentioned above, the probable continuation of the Pofadder Lineament can be followed to the west of Luderitz. There, however, there is no indication of large-scale rotation of older structures. It is interesting to note that the NMC in the area from Aus to Luderitz simulates the South African counterpart on a smaller scale in that the strike of the dominant foliation swings from northwest, through eastwest t o northeast when proceeding from east to west, while the rest of the area underlain by the NMC in Namibia, strikes northwesterly and parallel t o the Kheis belt. METAMORPHISM

Although the distribution of.the various mineral parageneses indicative of the grade of metamorphism have been documented in the past for many regions in Namaqualand, Bushmanland and Namibia, it is only in recent years that efforts have been made to organize the data in an attempt t o unravel the complexity of the metamorphic history of the NMC. The first attempts at determining metamorphic zonation (Joubert, 1971 ; Blignault et al., 1974) present a relatively simple picture which is clearly oversimplified. Since the main metamorphic phase, the rocks of the NMC have undergone at least one phase of deformation which was accompanied and outlasted by a phase of mainly thermal metamorphism.

Western Namaqualand The metamorphic zonation for western Namaqualand as determined by Joubert (1971) is shown in Fig. 10.9. The pelitic rocks in the most westerly zone (A) contain staurolite while, in the same rocks of zone B, kyanite also appears and in zones C and D, with one known exception, sillimanite is the only mineral of the A12Si05polymorphs in these rocks. Hypersthene enters the mineral assemblages of mafic rocks in zone C while both ortho- and

695 clinopyroxenes are found in the same parageneses in zone D. Cordierite is encountered in zones C and D, but it is only in zone D that garnet is a common associate of cordierite. Recent work in coastal Namaqualand has shown that the granulite facies extends nearer the coast than previously thought and that kyanite occurs very close to the coast at Hondeklip Bay. Clifford et al. (1975) concluded that metamorphic conditions in the Nababeep district exceeded temperatures of 800°C and pressures of 6 kb and that prograde metamorphism proceeded via kyanite reflecting an average geothermal gradient of less than 35"C/km.

Namibia Blignault et al. (1974) described the Namibian metamorphic zoning as increasing in grade both from the north and the south so that the metamorphic grade diminishes away from the central zone where chamockitic rocks are found (Fig. 10.9).The zones are recognized as follows: (a) The northern low-stage zone ( 1 ) is related to metamorphism in rocks underlying the Sinclair Group in the north. Northwest of Aus primary chlorite occurs in homblende-bearing schists. (b) In the northern medium-stage zone ( 2 ) calcareous rocks contain diopside, scapolite and occasionally anthophyllite, while the hornblende of the amphibolites is blue-green or green in colour. Plagioclase has an average anorthite content of 38%and epidote is also present. (c) In the high-stage zones ( 3 and 5 ) , muscovite and quartz are no longer in stable assemblage. (d) The orthopyroxene zone ( 4 ) , situated within the high-stage zone, includes the charnockitic rocks described above. (e) The southern medium-stage zone ( 6 ) is recognized by the absence of quartz + chlorite and the presence of cordierite, fibrolite and conformable leucosome banding. East of Vioolsdrif cordierite (+ spessartite), staurolite and andalusite have been found in this zone. (f) In thesouthern Zow-stage zone ( 7 ) ,comprising the bulk of the Richtersveld area, white micas, biotite, epidote, actinolite and hornblende form stable phases while quartz + chlorite are common associates. The indications are that the metamorphic gradation, in contrast t o westem Namaqualand (see Clifford et al., 1975), follows an Abukuma-type facies series. The metamorphic zonation as shown for Namibia, however, does not take into account the mineral paragenesis spinel + hypersthene + sapphirine + almandine + cordierite + biotite found by Beukes (1973) which clearly indicates high-stage conditions, but lies on the extension of the medium-stage zone north of the Orange River. Kyanite has so far only been found in one locality in the NMC of Namibia. It was seen in the drill core of a borehole at

696 Tantalite Valley where its presence has been ascribed to the emplacement of the intrusive (A.C. Moore, pers. commun., 1978). Jackson (1976) estimated temperatures of the amphibolite metamorphism to range from 650°C on the border of the area around Aus to 780-800°C in the centre and pressures ranging from 4 , 5 t o 6 kb with geothermal gradients of 30-50” C/km.

Bushmanland It is difficult to reconcile the metamorphic zonation as shown by Von Backstrom (1964), Vajner (1974) and Cornell (1975) with the mineral parageneses found in the eastern regions. From the early descriptions of Rogers and Du Toit (e.g., 1908), it is clear that there is an increase in metamorphism from east to west, a pattern which is also apparent in the area from east to west across Upington. If the orthopyroxene zone of Blignault et al. (1974) is extended from Namibia t o the southeast, it conforms well with the location of the charnockitic adamellites of Keimoes area. Here Schultz (1976) tentatively estimated a temperature range of 650-800°C and pressure of 4 1kb for the metamorphism. Further t o the west and around Kakamas, the grade of metamorphism is again lower, also conforming to the zoning in Namibia, but here the rocks are typically those of the types found in central and western Namaqualand . The presence of kyanite in areas where both sillimanite and andalusite are found, was recorded by Rogers and Du Toit (1908) to the west of Prieska, but according to B.J.V. Botha (pers. commun., 1978), this mineral is commonly found along the western side of the Kheis belt. Rogers and Du Toit (1908) also recorded the occurrence of small needles of an amphibole “allied to glaucophane” in an extremely fine, laminated groundmass of mylonites along the Doornberg Fault separating the craton and the Kheis belt. They suggested that the glaucophane developed in the most intensely sheared portions of a granite.

*

Metamorphism associated with the 1000 Ma event Not all geologists agree that the rocks of the NMC underwent more than one cycle of metamorphism. It is recorded by Clifford et al. (1975), for instance, that the F3 structures postdate the main metamorphism while subsequent metamorphism is not considered. There are, however, numerous indications that a metamorphic episode, largely thermal in nature, operated after extensive shearing and tectonic sliding had already taken place. Whether this applied to all the rocks described as belonging to the NMC is not clear at present, but it is certainly in accordance with the evidence gathered in central and western Namaqualand. From the mineral parageneses found in shear zones and from certain

features of these belts, interesting conclusions can be drawn. Since the shear zones cut across rocks ranging in metamorphic grade from greenschist t o granulite facies, they also contain a variety of mineral parageneses indicating more than one metamorphic grade even along the same lineament. The metamorphism due to shearing is naturally always retrogressive as a lower grade of metamorphism is indicated along the shear zone than that of the rocks traversed. It is also obvious that shear zones of similar trend and almost certainly of similar age, have mineral parageneses indicating metamorphic conditions ranging from greenschist t o amphibolite facies, depending on the metamorphic grade of the rocks traversed. Thus, along tectonic slides and shear zones such as the Pofadder Lineament, the Garies Shear and in the east, southeast of Kakamas, the presence of sillimanite has been recorded. Sometimes the needles of sillimanite are arranged parallel t o the recrystallization lineation common to the shear zone, indicating that the P-T conditions must have been such as to allow the formation of this mineral. In the west, however, where much lower metamorphic conditions prevailed during shearing, a zone parallel and west of the Garies Shear has been mapped where the rocks traversed have been converted t o phyllonites of the greenschist facies. Muscovite and chlorite are the main components of the phyllonites nearer the coast, but limited amounts of these minerals are sometimes encountered along shear zones across the granulite areas, probably indicating recurrence of movement at a later stage. The most convincing evidence indicating that the thermal metamorphism outlasted the shearing of the rocks of the NMC is the occurrence of corundum along a tectonic slide on the farm Rietfontein, some 27 km southeast of Springbok. There are two generations of several different minerals. Fine, often fibrolitic, crystals of sillimanite along bands defining the earlier schistosity are sometimes found folded during F3 deformation. A later and much coarser generation of undeformed crystals of sillimanite, sometimes with the needles paralleling the axial planes of tiny folds, are also seen. Cordierite also occurs in two generations with the earlier cordierite being densely pinitized. Large well-twinned porphyroblasts of cordierite enclosing fine needles of sillimanite folded during deformation, indicate growth of this mineral during or after deformation. Two generations of biotite in the NMC are evident in a very large proportion of the thin sections examined. The older biotite occurs in short stubby crystals frequently showing kinking of the flakes as well as greenish alteration, while the later generation forms large clear, strongly pleochroic flakes, sometimes seen along the axial planes of folds, or as unfolded flakes memitically crystallized around small fold closures. In the granulite-facies area south of the Springbok, rocks consisting of plagioclase, garnet and clinopyroxene have been found and this paragenesis would indicate a high grade of metamorphism but for the presence of

698

sphene. The grains of sphene, however, are always found to be surrounded by expansion cracks and are clearly of a later generation. Zoned garnets, often found in the rocks of the NMC, are sometimes described as having been derived from polyphase metamorphism. Some other minerals found rarely in Namaqualand and which would not have survived the higher pressures indicated by the mineral parageneses, include wollastonite and scapolite. As the possibility of an even later Pan-African metamorphism (+500Ma) in the coastal regions of Namaqualand exists, interpretations in those areas are suspect. Granites of the Concordia type, dated at 1000 Ma, however, occur within 20 km of the coast and just east of the eastern limit of the Late Precambrian formations there. The most reliable evidence for a late metamorphic event in the coastal area is the occurrence of large garnets in some of the representatives of the dyke swarm there. The presence of chloritoid was also recognized in two widely separated localities in phyllonites from shear zones of coastal Namaqualand. In the one case the foliation is curved around the metacrysts of chloritoid and formation appears to have taken place contemporaneously with the deformation, while in the other specimen, the mineral crystallized across the foliation and was possibly generated at a later stage. Staurolite in the aluminous schists of the coastal regions is often zoned, having dense cores with numerous inclusions similar t o the garnets found elsewhere. There are, apparently, also two generations of staurolite; one where the crystals are aligned parallel t o the foliation, the other being large idioblasts of staurolite lying across the foliation. Late staurolite was also found in an aluminous schist collected at the foot of the escarpment west of Springbok. In this rock sillimanite occurs with muscovite as sheaves of fine needles in thin schistose bands while the growth of large kyanite crystals disturbed the regular banding. The staurolite in this rock is clearly of a later generation as it lies across the foliation. In the Keimoes area, a finely schistose rock composed mainly of muscovite and quartz formed by shearing, can be seen containing large idioblasts of staurolite showing no directional orientation. The development of feather amphibolite seen on nearly all surfaces of slip in mafic rocks throughout Namaqualand and Bushmanland can only be ascribed t o the 1000 Ma thermal event. There are thus many indications that the main metamorphism responsible for the formation of the granulite facies parageneses of western Namaqualand was followed by a mainly thermal metamorphic event. Since this event was probably coeval with and outlasted F3 (and F4?) deformation, it does not appear likely that it is just due to a relaxation of the pressures operating during the main metamorphism, but should be considered as a separate episode. The more common occurrence of minerals like wollastonite, scapolite and

699 andalusite and the general absence of kyanite north of the Orange River and along the main Kheis belt, raises the question of whether the rocks underlying those areas were ever subjected to the type of metamorphism described for western Namaqualand. The origin of rock types consisting mainly of cordierite and anthophyllite or only of sillimanite and corundum has received much attention and is controversial, but it is generally agreed that once material of suitable composition is arrived at, only thermal metamorphism is required to produce those rock types. It is therefore necessary t o list these rocks as possible products of the 1000 Ma thermal event. The occurrence of cordierite-anthophyllite rocks has been recorded from south of Vioolsdrif, in the Tantalite Valley area of Namibia and from south of Pofadder. Apart from the sillimanite-corundum rock now being mined west of Pofadder, similar rocks have been mapped on Gamsberg and its vicinity. TIMING O F DEFORMATIONAL AND METAMORPHIC EVENTS

The majority of isotopic age determinations carried out on rocks from the NMC have been listed by Burger and Coertze (1973) and although numerous ages have become available since then, the picture has not materially changed since 1973. Apart from the Vioolsdrif and cratonic areas, all the ages determined in the NMC fall within the range of 840 t o 1180 Ma, with few exceptions. Many of the available ages have been determined on minerals from the pegmatites and only one pegmatite (from east of Vioolsdrif and south of the Orange River) yielded an aberrant U-Pb age (1426 60 Ma). The fact that the granites of the Concordia and Straussberg types all returned ages of 1100k 200 Ma and are of the same age range as the pegmatites, it is clear that there is some relation (Table 1O.IV). As the pegmatites are most frequently emplaced along F4 structures, their ages must reflect a minimum age for that event. The 1 1 0 0 k 200 Ma ages have been found from the edge of the craton (Fig. 10.7) in the east, across Namaqualand to the western exposures where the rocks of the NMC disappear under the Gariep. The same range of ages has been found in rocks underlying the Karoo in the south and as far north as Aus in Namibia as well as in Natal. If all these ages are plotted on a map, they do not, as could have been expected, conform to any obvious pattern. The areas of granulite facies metamorphism do not reflect a generally younger age than elsewhere, nor are older ages clustered in the proximity of the 1200 Ma-old Koras or near the sediments of the Gariep in the west. Instead, the older and younger ages within the same range are scattered rather indiscriminately throughout the area. It is thus rather puzzling that, at the time of extrusion on surface of the Koras lavas in the east, granites were being emplaced nor far away and into the entire NMC even for some time afterwards.

*

4 0 0

TABLE 10.W Sequence of events of sedimentation, volcanism, magmatism, deformation and metamorphism, Namaqualand Metamorphic Complex Age in Ma

Sedimentation and volcanism

Granitic intrusions

500

1000

Mafic intrusions

Nama sedimentation

Kuboos-Bremen line of plutons (pegmatites)

Gariep and Damara sediments

Lekkersing Granite Richtersveld complex pegmatites

Koras and Sinclair sediments and lavas

burg Granites Charnockitic adamellites

1

Metamorphism Pan-African metamorphism

dyke swarm, west noritoid bodies, Okiep mafic and ultramafic bodies along shear zones (?) dyke swarm S. of Kakamas

high-pressure metamorphism in east; extensive 1000 Ma metamorphic event, zoning in Namibia (?) Mz metamorphism

2000

mafic and ultramafic intrusions

Matsap-Kaaien quartzites Orange River Group

Deformation

1

-(F4)monoclinal folding (NNE) and shearing

(F3)open folding (EW) and shearing (WNW) concomitant with updift ( ? ) recumbent folds

(Fd

1

metamorphism

TRANSVAAL SUPERGROUP 2500 Marydale(?) and Seekoebaard

,

Skalkseput Granite ( ? ) Porphyroblastic granite gneiss (Nahaheep)

?

3000

Draghoender Granite (?) Metasediments

and lavas NMC (deposited on a granitic basement)

(?) recumbent folding

norites

morphism (?) zoning in W. Namaqualand

(F2)

701

There are a few older ages determined on rocks from widely separated areas. The Rb-Sr and K-Ar ages of a quartz-sericite schist from the Groblershoop For Ition were found to be 1750 Ma, while two Pb-Pb ages of rocks north of Pofadder were determined as k1575 and 1720 Ma. The latter two rocks most probably represent the eastern extension of lavas of the Orangc River Group. Another Pb-Pb age of k1755 Ma was determined on a h rvnblende gneiss occurring near Luderitz in Namibia. The most comprehensive geochronological study of a limited area was conducted by Clifford et al. (1975). Briefly their results are as follows: ~~

1000 Ma

F4 open folding

1070 5 20 1166 f 26

steep structures F3 open folding

1213 k 2 2 1400 ? 1850

F, recumbent folding

emplacement of pegmatites emplacement of noritoid bodies

emplacement of Concordia Granite F1

?

F1

?

Mz regional metamorphism Mi ? emplacement of Vioolsdrif Granite MI ? ~~

Emplacement of Concordia took place from late F, times t o the end of F 3 .

A monazite deposit in southwestern Namaqualand, dated at 1180 Ma, is situated along a structure typically that of the F3 deformational event (Joubert, 197413) and this age is in good agreement with the age suggested by Clifford et al. (1975) for this event. There are, however, indications in the Sinclair area of Namibia that shearing there could have taken place as long as 1360 Ma ago, but since rocks correlated with the Sinclair Formation are folded by east-west open folds (McDaid, 1976), the associated deformation continued for some time afterwards. There is general agreement on these latter events, but still a great deal of uncertainty as t o when the main metamorphism and recumbent folding took place. Vajner (1974), who correlated the structures found along the Kheis belt with those of the western part of Namaqualand, suggested that the tectonic development commenced in the Archaean and described the main phase of deformation (F,) as having occurred prior to the plutonic intrusions older than 2500 Ma. The later deformational events (F3 and F4) affected the Matsap Formation and are thus younger than 2050 Ma. The metamorphism during the Namaqualand tectogenesis when the NMC was structurally consolidated and large-scale updoming occurred, took place some 1250 t o 900 Ma ago. Cornell (1975) distinguished two phases of deformation in the Kheis belt; an earlier phase when large volumes of granitic magma were intruded and severe compressional folding occurred. This event is dated at 1380 Ma and

702 was followed by mild deformation associated with isostatic uplift and some late-stage intrusive activity about 1100 t o 800 Ma ago. As mentioned before, sillimanite schists associated with other lithologies typical of the Bushmanland Sequence occur well west of the “sillimanitein isograd” of Beukes (1973) and within the lavas of the Vioolsdrif area dated at 2000 t o 1800 Ma. This appears t o indicate that the NMC underwent the main metamorphism prior t o the extrusion of the lavas. In the east the metamorphosed rocks of the Kheis belt are unconformably overlain by the Koras Formation which is hardly deformed or metamorphosed, so that it can be concluded that that part of the area had already reached its present relative elevation some 1200 Ma ago. Similarly, in the west the rocks of the Gariep were being deposited some 900 Ma ago, possibly when the central parts of the NMC were still being uplifted. The isostatic adjustments were accompanied by large-scale shearing, but these movements were largely terminated by the time of the deposition of the sediments of the Nama Group. Since then minor movement occurred along existing lines of weakness and n o r t h s o u t h faulting took place, mainly along the escarpment of western Namaqualand. MINERAL RESOURCES

The list of minerals of economic importance occurring in Namaqualand is a long one, ranging in alphabetical order from alum to zircon, but apart from and excluding those in rocks of Late Precambrian age, the copper mines at Okiep, sillimanite, corundum and barytes in Pofadder area and the numerous minerals yielded by the pegmatites, nothing else was being mined until recently. The discovery of the base metal deposits of Copperton, west of Prieska, resulted in a flurry of exploration in Namaqualand and Bushmanland. Although baryte was being mined on a limited scale for a number of years at Gamsberg, it was only after the discovery of the Aggeneys deposits in 1971-1972 that the presence of base metal sulphides was recognized there. The minor copper showings at the Haib, east of Vioolsdrif, were recognized as a porphyry copper deposit approximately at the same time. Many other small copper deposits, however, have been investigated over the past 50 years. The base metal deposits of Namaqualand may be divided into five groups: (1) The Okiep copper bodies of the “noritoid suite” (mainly norites, diorites, anorthosites), generally situated along the so-called steep structures (Benedict et al., 1964), have been mined since 1852 and are the most important copper producers of Namaqualand at present. The ore reserves calculated at the end of 1972 comprised 24,494,000 tons of sulphide ore of an average grade of 1.58%Cu and 941,000 tons oxide ore of an average grade of 1.13%Cu (Hammerbeck, 1976). (2) The Aggeneys (Cu-Zn-Pb-Ag) and Gamsberg (Zn) deposits are asso-

703 ciated with the supracrustal rocks, mainly quartzites and aluminous schists, of the Bushmanland Sequence. There is still some controversy regarding the origin of the ores, but most geologists favour a syngenetic origin under volcano-exhalative conditions. The ore reserves at Aggeneys of three separate deposits are calculated at 194.5 million tons of 0.55% Cu, 208.4 million tons of 2.81% Pb and 1.82% Zn. The amount of Ag ranges from 17 to 58 grams/ tonne. At Gamsberg the ore reserves are conservatively estimated at 93 million tons of 7.3%Zn and 0.5% Pb (Hammerbeck, 1976). (3) Copper-zinc deposits are associated with metamorphosed volcanic rocks such as those at Copperton and Areachap. At Copperton the main ore body contains 47 million tons of ore averaging 1.7% Cu and 3.8% Zn (Middleton, 1976). The ore body contains about 55% sulphides by volume, mainly pyrite, chalcopyrite, sphalerite, pyrrhotite and minor amounts of menopyrite, galena, magnetite, neodigenite and molybdenite. (4) Although listed only as a minor occurrence of copper (Hammerbeck, 1976), it has now been recognized that the very large low-grade ore deposit at the Haib, east of Vioolsdrif and immediately north of the Orange River, has many features in common with porphyry copper deposits. This deposit and others further to the west have not been described in the literature and at present prices for base metals, are not being exploited. The Haib deposit is unusual in that the rocks belong to the Orange River Group, dated at 2000 to 1800 Ma. (5) Minor copper-nickel mineralization has been found associated with the mafic and ultramafic intrusives situated along the major lineaments, but none of these has proved t o be economically viable. SUMMARY OF EVENTS

The rocks included in the Namaqualand Metamorphic Complex can be divided into several regions, namely the main Namaqualand-Bushmanland region, the Kheis belt and its junction with the Kaapvaal Craton, the Vioolsdrif area and the extension into Namibia. The rocks in Namaqualand consist of a thin but consistent sequence of supracrustal rocks overlying porphyroblastic gneisses. On lithological grounds, the supracrustal rocks have variously been correlated with Archaean formations, with the 2000-1800 Ma-old igneous rocks of Vioolsdrif or with the sediments of the Transvaal Supergroup, but since these NMC rocks have undergone several phases of deformation and metamorphism, the correlations with other less disturbed formations are conjectural. The igneous rocks of Vioolsdrif consist of a series of andesitic t o acidic lavas of calc-alkaline affinity, intruded by gabbros, diorites, tonalites, granodiorites and granites. The rocks along the Kheis belt in the east differ from those of Namaqualand in having a preponderance of mafic rocks and in the absence of the common association of quartzite and aluminous schist. The rocks of the Kheis belt are similar t o those of the extension into Namibia

704

where quartzites are almost completely absent. The rocks of the Kaapvaal Craton become deformed along the western edge of the craton and this aspect is not always recognized; correlations across this boundary have still t o be settled. The rocks of the NMC have been intruded by numerous bodies of ultramafic and mafic rocks, granites, dykes and pegmatites. They are overlain by patches of the 1200 Ma-old formations, by highly deformed Late Precambrian sediments along the west coast and t o the north in Namibia, and extensively by flatlying Nama and Karoo sediments. Four structural events have been recognized in Namaqualand, two of which are of considerable consequence. I t is, however, improbable that folds displaying the same features and trends in widely separated parts of the area were developed synchronously. The main deformation, which resulted in large recumbent folds, was accompanied and outlasted by the main phase of metamorphism. The imprints of several phases of metamorphism have been recognized in the rocks of the NMC. The peak of metamorphism occurred prior to the period of open folding and its associated shearing which was accompanied and followed by a phase of low-pressure, high-temperature metamorphism. The age of the main phase of metamorphism has been recorded as +1213 Ma, but there are indications that it could be considerably older. This phase of metamorphism gave rise to zoning in western Namaqualand with staurolite in pelitic rocks near the coast, kyanite further inland and sillimanite in the higher grade areas t o the east. Granulite facies metamorphism in the Springbok area is indicated by the presence of hypersthene in the mafic rocks with the occurrence of two-pyroxene granulites further inland. It was found that temperatures and pressures exceeded 800°C and 6 kb, respectively, in the high-grade area and that prograde metamorphism proceeded via kyanite. As the mineral parageneses found along the main part of the Kheis belt and along the extension into Namibia indicate only low-pressure, high-temperature metamorphism, these parts of the NMC only underwent the later metamorphism or the P-T conditions were different at the time of the earlier imprint. At about 1200 Ma ago the NMC acquired a degree of rigidity and shearing of the complex commenced. Lavas were extruded and the sediments of the Koras Formation and the Sinclair Group were deposited in the east and the north. In these areas only faulting occurred subsequently, but further west uplift and folding continued. Large open folds, with east-west trends in Namaqualand and east of Luderitz, and northwesterly trends along the eastern margin of the NMC, developed simultaneously with uplift. Largescale dextral movements along the eastern edge of the NMC are evident while west-northwesterly shear zones, also with dextral displacement, transected the greater part of the complex. In the coastal areas sinistral shearing, mainly in a north-south direction, occurred.

705

Granites were emplaced into the entire NMC during the period 1000 f 200 Ma and they were accompanied by extensive thermal metamorphism. Minerals which could not have survived the higher temperatures of the earlier metamorphism were formed and rocks of suitable composition were transformed into cordierite-anthophyllite and sillimanite-corundum bodies. Elevation along a median zone of the Kheis and Namibian belts exposed rocks of charnockitic composition and gave rise to the metamorphic banding there. A mafic dyke swarm was emplaced into the NMC in the east and numerous mafic and some ultramafic bodies were intruded along the major shear zones. An age of +lo70 Ma has been suggested for the time emplacement of the copper-bearing hypersthenite, norite and diorite bodies of Okiep. Pegmatites were intruded throughout the NMC, occurring in profusion along a belt skirting the main Namaqualand-Bushmanland area and separating it from the Kheis and Namibian belts. Shearing movements continued in the western parts and the extensive dyke swarm, stretching from Bitterfontein in the south to central Namibia, was intruded. These movements prepared the troughs for Gariep and Damara sedimentation and continued afterwards to deform these formations. Emplacement of the Richtersveld Complex took place, possibly simultaneously with the Lekkersing Granite which invaded the basal part of the Gariep Formation. In the east and following on the emplacement of the pegmatites, convergent wrenching resulted in minor thrusting and further movement along the Doornberg Fault where the occurrence of glaucophane has been recorded. Uplift of the NMC and peneplanation ceased prior t o the deposition of the generally flat-lying +600 Ma-old Nama rocks and was followed by extensive faulting in western Namaqualand. The Kuboos-Bremen line of plutons were emplaced some 550 Ma ago, and Pan-African metamorphism and deformation of the Gariep, Damara and western parts of the NMC occurred.

C. Eastern or Natal sector of the Namaqua-Natal mobile belt in Southern Africa * INTRODUCTION

The eastern or Natal sector of the Namaqua-Natal mobile belt is essentially a complex of granitic and migmatitic gneisses with a general E to ENE structural trend and a prevailing medium to high grade of regional metamorphism. These features are indicative of erosion t o a deep crustal level, perhaps of the order of 20-25 km. Although the regional outcrop is restricted to a narrow N-S-trending

* By P.E. Matthews.

gronite

gneiss

Amphi b o l i t c s 1 MADlOlMA

NAPPE

Amphibolites ( m a g n e t i t e qtz.1 ( dolomite]

complex

(kyanttr gneisrl 1.NKOMO NAPPE Gneissic o m p h i b o l l l e

I

1

I

I

I

0

10

Gronite

gneiss

1

20 M i l e s

Fig. 10.10. Regional geology of the northern frontal zone of the Namaqua-Natal

Amphiboliles (magnetite qtz I Gneissic omphibolites (magnetite q t z 1 Idolomite I

1 kyanile gneisrl

..,....

Mobile Belt in Natal, southern Africa.

Metogobbro S e r p e n t I ni t e Major thrust-taults

v o n "Doer D l o t e

707 erosional inlier subparallel to the Natal coastline (inset map Fig. 10.10), it nevertheless provides a transverse section about 220 km in length across the structural trend of the mobile belt, and includes the northern structural front but does not extend t o the southern limit of the belt. Radiometric dating of samples from ten widely-spaced localities has provided mineral and whole-rock Rb/Sr and U/Pb zircon ages in the range 9001194 Ma (Nicolaysen and Burger, 1965; Burger and Coertze, 1973; Geol. Surv. pers. commun., 1978). Contrasting lithological, structural and metamorphic features allow subdivision of this sector of the belt into four major E-W-trending zones, which are : (1)A northern frontal zone, some 15 t o 30 km in width, which is characterized by extensive thrust-sheets of predominantly amphibolitic gneisses that have been transported northwards onto the southern flank of the Archaean Kaapvaal craton. (2) A migmatite and granite-gneiss zone, about 60 km in width, which is made up of several alternating linear belts of migmatites and granitoid gneisses. (3) A granitic zone, about 120 km wide, that contains extensive (batholithic) developments of homogeneous, generally megacrystic granitoid rocks within a framework of granitic and migmatitic gneisses. (4) A southern granulite zone, that is limited to the southern part of the basement inlier and contains metasedimentary granulites and intrusive charnockitic rocks. NORTHERN FRONTAL ZONE

Some of the major geological features of this zone, and the southern margin of the Kaapvaal craton are shown in Fig. 10.10. This map and the schematic downplunge profile of the frontal zone shown in Fig. 10.11, incorporate new data obtained during a current regional survey that was initiated by the writer in 1975 as part of the South African contribution t o the International Geodynamics Project *.

Major geological features of the northern foreland On a regional scale there is a remarkable contrast between the intensely deformed metamorphic and migmatitic formations of the frontal zone of the mobile belt and the principal lithological units of the cratonic region t o the north. In brief, the northern foreland is made up of a crystalline, mainly

* Sponsored by the CSIR, this project has been planned for the period 1975-1979. The writer is indebted to the CSIR for generous financial assistance, and to a number of postgraduate students for their collaboration, in particular A. Schulze-Hulbe, s. Rigotti, T. Smalley, R. Harmer and G. Charlesworth.

708 granitoid basement that is overlain unconformably by a thick (-6000 m), relatively undeformed and unmetamorphosed volcano-sedimentary sequence, that constitutes the Insuzi Subgroup of the Pongola Group (Du Toit, 1931; Matthews, 1959). Rb/Sr whole-rock and U/Pb zircon datings from Insuzi lavas have indicated an age in the range 3090 t o 3150 Ma for this intracratonic basin sequence (Burger and Coertze, 1973). Within the southern part of the craton, the Insuzi Subgroup is composed of substantial units of orthoquartzite (800-1800 m thick) that alternate with subordinate units of sheared basaltic, amygdaloidal lavas and phyllitic sediments. The preInsuzi basement is an extensive development of a remarkably uniform megacrystic, potash-rich granite with intrusive contacts against narrow, widely spaced synformal belts of greenstones and contact amphibolites of the Nondweni Formation (Fig. 10.10). Structural framework of the frontal zone This region of northward directed overthrust structures is made up of two major E-W trending structural zones with contrasting lithological and metamorphic features. (1)The northern zone is the Natal Thrust-belt, a narrow (2-5 km wide) southerly dipping, imbricate complex of low-grade metamorphic formations, that defines the structural and metamorphic front of the mobile belt (Du Toit, 1931; Matthews, 1959). (2) The southern zone is the Natal Nappe Complex (Matthews, 1972), a group of four extensive, essentially flat-lying thrust-sheets composed mainly of amphibolitic gneisses with some infolded elements of a granitic infrastructure. These thrust-sheets pass southwards into a steep t o vertical belt of high deformation, that probably represents the downwarped root-zone of the nappe complex, and defines the southern limit of the frontal zone (Figs. 10.10 and 10.11; McCarthy, 1961; Gevers, 1964; Colombin, 1968; Guy, 1974; Schulze-Hulbe, 1977). Nu tal thrust-belt This belt is made up of two major, southerly dipping overthrust units. The basal unit is an extensive tectonic wedge which is restricted to the western sector of the thrust-belt and contains imbricated slices of two unconformable sequences. The older Mfongosi Formation is essentially an alternating sequence of meta-argillites and schistose metalavas with some relict pillow structures and amygdaloidal textures. Minor components are thin units of quartzite, limestone and banded ironstone. The younger Ntingwe Formation is the basal remnant of a transgressive marine sequence, that includes a basal conglomerate followed by grits, shales and dolomitic limestones. The upper structural unit of the thrust-belt is an extensive zone of large imbricate slices of intensely folded units of the Mfongosi Formation. In the western sector, the major thrusts are inclined at low angles (10 to 20") to the south, and the basal thrust is more or less concordant with the

Fig. 10.11.Schematic down-plunge structural profile across the northern frontal zone of the eastern or Natal sector of the Namaqualand-Natal Mobile Belt, The regional structural plunge is usestward, and the plane of projection is close to the western boundary of the map in Fig. 10.10,(Note the reference line A on the profile and the map). The major thrust-nappes are labelled as follows: Nh = Nkorno; Md = Madidima; Mn = Mandleni; and Tu = Tugela.

710 bedding of undisturbed Ntingwe strata that rest unconformably on older foreland formations, and are exposed at intervals for about 50 km along the southern margin of the Kaapvaal craton. When traced eastwards, the southerly inclination of the thrust-belt and the underlying Ntingwe strata increases progressively t o dips of 45 t o 6 0 ° , a relationship that clearly indicates a late differential downwarping of the thrust-front. In addition, the lower imbricate zone wedges out, so that the bounding thrusts converge onto a basal thrust within a thin zone of phyllonitic schists, that probably represent the attenuated basal sequence of the Mfongosi Formation. These features suggest that the basal thrust functioned essentially as a detachment or d6collement slide that passed stratigraphically upwards and northwards from the base of the Mfongosi Formation into the basal part of the overlying Ntingwe Formation. The northward tectonic displacement of these formations from the southern margin of the Kaapvaal craton was accompanied by intense folding and imbrication within the frontal thrust-belt, which was overridden from the south by the basal thrust-sheet of the nappe complex.

Natal nappe complex Due t o a variable but regional westward structural plunge, the four essentially flat-lying major thrust-sheets of this complex outcrop in an ascending structural sequence from east t o west (Figs. 10.10 and 10.11). An aspect of exceptional interest is the occurrence of extensive but discontinuous sheets of talc schist with pods of serpentinite along the basal slides of all the major thrust-sheets. In places, the serpentinite also occurs as extensive lenticular masses up t o 300 m in thickness, within a thin envelope of talc schist. In general, the upper thrust-sheets are made up of folded segments of the Tugela Group, a major metavolcanic-sedimentary sequence, while the lower thrust-sheets also contain tectonic slices and infolded elements of an infrastructural complex of granitic gneisses that probably represent parts of a reworked crystalline basement to, at least, the northern facies of the Tugela Group. In places, the granitic complex has undergone remobilisation so that the marginal zones may contain locally derived inclusions, and the associated Tugela formations usually contain extensive networks of deformed granitic, aplitic and pegmatitic sheets and veins. The Tugela Group is made up chiefly of migmatitic hornblende gneisses and amphibolites with occasional relict pillow structures. Subordinate components are pelitic and semi-pelitic schists and gneisses, and a few thin, generally discontinuous units of magnetite quartzite and lithographic limestone. The upper and most extensive thrust-sheet (the Tugela nappe in Figs. 10.10 and 10.11) is composed almost entirely of Tugela amphibolites and a number of metabasic-ultrabasic intrusions with occasional serpentinites containing podiform segregations of chromite (Du Toit, 1931;

711 Lambert, 1962). An interpretation of this association as the upper part of an ophiolitic complex and the tectonic implications (Matthews, 1972) are mentioned later. Metamorphism and tectonics Although there is an apparent increase in metamorphism southward from the thrust-front (Du Toit, 1931; Matthews, 1959, 1972; Cain, 1975) recent mapping has shown that the original zonation has been modified by northward overthrusting which has juxtaposed upper amphibolite migmatitic gneisses of the Tugela Group against greenschists of the Mfongosi Formation along the northern boundary of the nappe complex. Moreover, the rocks of this complex are polymetamorphic and have been affected by at least three major phases of deformation. The earliest phase is represented by large- to small-scale, sub-isoclinal folds, with associated intrafolial folds and a dominant phase of foliation which is more or less parallel t o the lithological layering. Within the Tugela Group, this deformation probably occurred under upper amphibolite conditions of metamorphism which are recorded by the sporadic occurrence of sillimanite + garnet cordierite in rare metapelitic schists, and clinopyroxene in some amphibolites. It was accompanied by extensive migmatisation, local remobilisation of the infrastructural granitic complex, and widespread emplacement of granitic sheets and veins. The second phase of deformation (Dz) was evidently a regional episode of lateral ductile shearing with a major component of northward rotation that reoriented the D1 fold-system to a low southerly inclination. It also produced extensional effects including attenuation and disruption of the more competent lithological units, particularly the granitic sheets. A diagnostic feature is the lack of a systematic relationship between the Dz extensional structures and the geometry of the distinctive superimposed D3 folds. The Dz deformational phase was probably accompanied by a slight change in metamorphic conditions t o a higher pressure regime, with the development of kyanite and garnet in the metapelitic units, and garnet in the amphibolites. Prior to the northward emplacement of the nappe complex which represents the third major deformational phase (D3),a series of extensive sub-horizontal to southerly dipping ultrabasic (serpentinite) sheets were emplaced more or less parallel to major fold-limbs. These sheets subsequently acted as detachment zones along the basal slides of the major thrust-sheets. A characteristic product of the nappe-forming D3 deformation was widespread disharmonic folding with a consistent sense of asymmetry indicative of regional overthrusting towards the north. This deformation took place under amphibolite facies conditions, with some retrogressive effects along the boundary with the frontal thrust-belt. The total tectonic displacement involved in the northward emplacement

*

712 of the nappe complex and the development of the frontal imbricate zone is estimated to have been not less than 100 km. This estimate is derived from a down-plunge structural profile of the frontal region (cf. Fig. 10.11) on the assumption that the various thrust-sheets were derived from lateral segments of the same depositional basin. MIGMATITE AND GRANITE-GNEISS ZONE

This zone is about 60 km in width and is made up of four major and some minor, ENE trending, linear belts of highly deformed, vertical to steeply dipping migmatitic gneisses with intervening belts of well-foliated granitoid rocks including extensive units of porphyroblastic and augen gneisses (Kent, 1938; Rhodes, 1966; Leith, 1966; Williams-Jones, 1968; Robinson, 1969; Roberts, 1971; Leith and Rhodes, 1971; Guy, 1974; Cain, 1975). Structural and lithological relationships suggest that the granitoid belts represent large anticlinorial elements of a crystalline basement, while the adjacent migmatite belts are down-folded remnants of an original supracrustal sequence that constitutes the Mapumulo Group (Matthews, 1972; Cain, 1975). In contrast to the predominantly amphibolitic gneisses of the Tugela Group within the frontal region, the Mapumulo Group is an assemblage of heterogeneous migmatitic gneisses that occur in two main associations, which are: (1)a group of predominantly quartzo-feldspathic gneisses; and (2) a group of uniform and banded hornblende-biotite gneisses with some interlayered amphibolites, which together probably represent a meta-greywacke-volcanic sequence. Infolding of basement and cover evidently took place under high t o medium grade metamorphic conditions which are recorded by occurrences of sillimanite + garnet k cordierite gneisses, and two pyroxene granulites within each of the migmatite belts. Widespread migmatization of the Mapumulo formations was accompanied by partial remobilization of the granitic basement, particularly near contacts with the migmatitic supracrustal rocks. In addition, anatectic melts were generated and emplaced within the migmatite belts as elongate, syntectonic granite plutons (Leith and Rhodes, 1971). GRANITIC ZONE

With a width of about 120 km and a general E-W structural trend, this zone is characterized by extensive units of remarkably uniform, massive to foliated granitoid rocks, that contain dispersed inclusions and rafts of various metamorphic and migmatitic rocks, particularly near concordant to subconcordant contacts with a complex framework of granitic and migmatitic gneisses (Kent, 1938; Cherry, 1947; Hargraves, 1952; Matthews, 1952; Amett, 1953; Lambert, 1954; Scheepers, 1954; Thompson, 1955; Gold, 1957; Hill, 1958; Jennings, 1963; Davies, 1964; Craig, 1965; McIver, 1966;

713 Geol. Surv. unpubl. maps). Transgressive contacts have been reported from the southern part of this zone, where extensive masses of marble with thin intercalations of siliceous schist and quartzite, are completely surrounded by intrusive granites (Du Toit, 1918, 1919, 1920,1946; Simpson and Tregidga, 1956; Otto, 1977). In general, the granitoid units of this zone appear t o be individual components of plutonic complexes of batholithic dimensions. Early members are represented by deformed, elongate masses of dioritic and tonalitic gneisses within the migmatite belt along the northern margin of the granitic zone, which is well-exposed in the Mgeni Valley, northwest of Durban (Matthews, 1952; Davies, 1964). The border zones of the main granitoid complexes are characterized by extensive units of porphyroblastic and gametiferous granite--gneisses that contain locally derived inclusions and rafts of migmatites including paragneisses with sillimanite + garnet f cordierite, and occasional two pyroxene granulites. These granitic components have some of the field characteristics of S-type granitoids in the classification proposed by White and Chappel (1977). The main components appear to attain batholithic proportions, and are massive to foliated, generally megacrystic hornblende-biotite granodiorites and biotite adamellites. They usually contain widely dispersed inclusions of amphibolite and occasional xenoliths of charnockitic rocks. These features are characteristic of I-type granitoids of deep-seated derivation (White and Chappell, 1977). Late stage components are widely scattered, small plutons of microcline granite with associated systems of granitic sheets and pegmatites. Chemical analyses of the various components of the batholithic masses (Grout, 1935; Hargraves, 1952; Simpson and Tregidga, 1956; McIver, 1966; Matthews, unpub. data) reveal normative proportions of Q-Or-Ab that plot within the well-established granitic field for this system (Tuttle and Bowen, 1958; Winkler, 1974). SOUTHERN GRANULITE ZONE

Within the southernmost part of the basement inlier, granitoid rocks of the granitic zone have transgressive, intrusive contacts against an E-Wtrending zone of metasedimentary granulites and an extensive development of charnockitic rocks (Du Toit, 1919, 1920,1946; Gevers, 1941; Gevers and Dunne, 1942; McIver, 1966; McIver and Gevers, 1968). The metasedimentary sequence consists mainly of banded hypersthene-garnet granulites with thin bands of quartzite, which shows near isoclinal folding overturned towards the north. Infolded with these granulites is an early intrusive mass of porphyritic diorite. This folded sequence is intruded on the south by an orthopyroxene granodiorite which contains granulite and diorite inclusions,

714 and is a major component of a batholithic mass of charnockitic rocks (McIver, 1966; McIver and Gevers, 1968). The various intrusive components define a differentiated suite of orthopyroxene-bearing rocks ranging from dioritic through quartz-dioritic and granodioritic t o granitic. The charnockitic character of these rocks has been attributed t o crystallization of a “normal” granitic suite under oxygen-deficient conditions (McIver, 1966). A Rb/Sr dating of biotite from a granitic phase of this charnockitic mass gave an age of 1011 Ma (Nicolaysen and Burger, 1965). DISCUSSION AND CONCLUSIONS

The limited number of radiometric dates reported so far from the Natal sector of the mobile belt are from granitoid rocks, and fall within the range 900 to 1200 Ma. This can be regarded as the possible time-span of the main tectono-thermal event within this sector of the belt. As the supracrustal Mapumulo Group within the central zone of the belt, and the Tugela Group within the northern frontal zone do not appear to have been affected by more than one orogenic cycle, they can be regarded as integral depositional components of the mobile belt. This interpretation is supported by the absence of any obvious stratigraphical counterparts within the northern cratonic region. Although intense regional deformation has obscured or obliterated original stratigraphic and structural discordances, the existence of a re-worked granitic basement t o the Mapumulo sequence, and at least, t o the northern facies of the Tugela sequence has been inferred from structural relationships. Polyphase deformation of basement and cover evidently took place under medium- to high-grade regional metamorphic conditions, and was accompanied by widespread migmatization and the emplacement of much granitic material, particularly within a broad zone about 100 km from the northern front of the mobile belt where syntectonic granitic intrusions attained batholithic proportions. Attention has been drawn t o major contrasts in the structural style and lithology of the main part of the mobile belt and the northern frontal zone where northward overthrusting with an aggregate lateral displacement of at least 100 km, led to the development of the Natal nappe complex and the associated dkcollement and imbricate structures of the Natal thrust-belt along the front of the belt. An important aspect of the structural succession within the frontal region is an upward increase in the relative proportions of metavolcanic t o metasedimentary formations, so that the upper thrust-sheet of the nappe complex is composed almost entirely of Tugela metabasites and amphibolites with some deformed pillow structures, in association with a number of extensive, layered metabasic-ultrabasic intrusions with occasional serpentinites containing chromite segregations. An interpretation of this association as the upper part of an ophiolitic

715 complex has led to the suggestion that the frontal zone contains tectonic slices of transformed oceanic crust and sediments which have been obducted from a marginal basin onto the southern flank of the Kaapvaal craton (Matthews, 1972). In terms of this tectonic model, the Natal granite-gneiss complex and the infolded meta-greywacke-volcanic Mapumulo sequence (Figs. 10.10 and 10.11) represent the northern margin of a continental plate that was deformed in an environment above a subduction zone, as the leading edge of the plate moved into the collision zone with the Kaapvaal Craton. In conclusion, it is of interest t o consider briefly an alternative interpretation of the northern frontal region of the mobile belt. Although the volcanogenic Tugela and Mfongosi sequences of the frontal zone cannot be correlated directly with the Archaean, Nondweni greenstone formations within the northern cratonic region, nevertheless, the Tugela and Mfongosi formations have certain lithological affinities with typical greenstone sequences (cf. Key et al., 1976). These features include an association of metabasaltic pillow lavas with occasional banded ironstones, subordinate developments of graphitic and aluminous schists and gneisses, and minor intercalations of limestone and quartzite. Another important aspect is the infolding of the northern facies of the Tugela Group with an infrastructural granitic complex from which biotite and zircon radiometric ages of 900 and 1100 Ma, have been obtained (Nicolaysen and Burger, 1965). In brief, these features seem to be consistent with the view that the Tugela and Mfongosi formations are components of a greenstone-type supracrustal sequence of possible Late Precambrian age, that was deposited during an early volcanic episode along the northern margin of the mobile belt. In terms of this interpretation, the tectonic development of the Natal nappe complex and the associated frontal dkcollement and imbricate structures of the Natal thrust-belt, could be attributed t o northward, lateral gravitational spreading from a region of major tectonic uplift that was related to the metamorphic and plutonic development of the central zone of the mobile belt.

D. The western part of the Zambezi Metamorphic Belt

*

In the extreme western part of Zimbabwe several isolated inliers of the Precambrian Zambezi Metamorphic Belt outcrop surrounded by unmetamorphosed and relatively undisturbed Sijarira, Karoo and Kalahari cover. The largest inlier southeast of Wankie which has been mapped in detail by the Geological Survey (Watson, 1962; Lockett, in press) includes the villages of Dett and Kamativi (Fig. 10.12). It is bounded t o the north by Karoo sediments and blanketed in the south by aeolian Kalahari sands. Northeastwards the

* By N.H.

Lockett.

Fig. 10.12. Geological map of the western part of the Zambezi Metamorphic Belt.

717 schists and gneisses pass under the cover rocks for a distance of some 150 km before being again exposed south and east of Kariba where they have been mapped by Broderick (1976). THE DETT-KAMATIVI INLIER

The inlier comprises four north-east trending belts of metamorphosed and highly deformed supracrustal rocks; the Kamativi, Tshontanda, Inyantue and Malaputese Formations, within large expanses of intrusive granitic gneiss. The Malaputese Formation consists of metasedimentary and mafic rocks exposed over a large area in the southeast of the inlier. The other three belts, composed largely of metapelites, lie within a curved linear zone along the northern margin of the inlier. Contacts of the granitic gneisses with the supracrustal rocks are often highly deformed, but are transgressive and recognizably intrusive. Upper amphibolite t o granulite facies metamorphism of the supracrustal belts is the earliest recognizable thermal event. A later, retrograde, greenschist t o lower amphibolite grade of metamorphism accompanied intrusion of the granitic rocks and emplacement of numerous pegmatite swarms in the linear belts. Three fold phases are evident from regional and small-scale structures in the Malaputese Formatica. An initial intense F1 deformation accompanied early high-grade metamorphism of the supracrustals, producing northeasttrending, isoclinal folds, and local large-scale recumbent structures in the southeast of the inlier. Granite intrusion probably preceded an F2 fold episode during which more open, northwest-trending stmctures were formed. Tight F3 folds mostly coaxial with F1 structures deformed both granitic rocks and the supracrustal belts. A variety of regional fold-interference patterns characterizes Malaputese rocks, although very tight F3 folding coupled with retrograde metamorphism has destroyed most of the early structures in the linear belts.

The Malaputese Formation Two stratigraphic domains were established in this unit. In the western part of the exposure arcuate domes of pink, granitic paragneiss are mantled by tightly folded mafic granulites, metapelites and quartzite. In the eastern domain a thick succession of hornblendeandesine amphibolites overlies metapelites in a series of quartzite-enclosed synforms. The metapelites comprise garnetiferous sillimanite-cordierite-andalusitebearing biotite gneisses and schists, and their retrograde derivatives, quartzchloritesericite schists. The mafic rocks, which contain local siliceous, calcareous and aluminousmagnesian intercalations, are interpreted as metamorphosed extrusive basalts

718 (Bahnemann and Lockett, 1979) incorporating thinly bedded volcanogenic sediments and tuffs. The linear belts The Kamativi and Tshontanda Formations are composed of garnetiferous mica schists, some andalusite-bearing, intimately veined by and infolded with garnet-tourmaline-mica pegmatites. These range in size from mere wisps and ptygmas to large sheet-like bodies over thirty metres wide, and were emplaced mostly during the main granitic episode. Gametiferous sillimaniteand cordierite-bearing gneisses, regarded as the high-grade parents from which the schists were derived, are locally preserved in parts of the Tshontanda Belt. The V-shaped Inyantue Belt merges with schists of the Kamativi Belt in the north and extends beneath a cover of Kalahari sand in the south. It is composed of garnetiferous sillimanite-cordierite gneisses and migmatites, with intercalations of calcsilicate rocks, graphitic schists, biotite-hypersthene granulites, quartzite and talc-chlorite schists. Two narrow, widely separated smaller belts of similar rocks were also mapped as part of the Inyantue Formation. The granitic rocks The granites are usually weakly foliated, nebulitic gneisses containing between 5 and 10% of biotite. Potash feldspar, often occurring as augen and porphyroblasts, constitutes between 40 and 60% of total feldspar in most of the gneisses. Along linear zones where the granitic rocks are highly deformed they become strongly foliated with blasto-mylonitic fabrics. Inclusions of supracrustal rocks, ranging in size from a few centimetres to bodies over one kilometre in length, are common in the granites. Near the larger bodies the gneisses are commonly biotite-rich and migmatitic, with numerous inclusions and schlieren. Granite intrusions post-dated the early tectono-thermal events in the supracrustals and probably caused the extensive retrograde metamorphism of metapelites in the Malaputese, Kamativi and Tshontanda Belts. ECONOMIC GEOLOGY

The supracrustal rocks of the inlier contain economic deposits of tin, wolframite, copper, mica and lead, and a number of faults along the northem marginal zone of the inlier are mineralized with fluorite.

719 Tin

Large shallow-dipping sheets of lithium- and tin-bearing pegmatites have been mined at three centres along the Kamativi Belt namely at Kamativi, Lutope and Kapata. The largest and most economically important are the Lutope and Kamativi deposits which occur within marked bulges in the belt of schists. Swarms of smaller, steeply-dipping tin pegmatites occur sporadically throughout the belt, becoming larger and more abundant near the main tin producing centres. These economic pegmatites postdate and are apparently unrelated to the garnet-tourmaline-mica-bearing type. The tin deposits of the Kamativi Belt form the subject of a thesis by Ewart (1960), while the Kamativi Mine pegmatites are described by Fick (1960) and Rijks and Van der Veen (1972). Wolframite

Stockworks of wolframite-tourmaline-quartz veins, some carrying copper sulphides, are located in narrower portions of the Kamativi and Tshontanda Belts at the RHA Mine, Tung Mine, Anglo claims, western Lutope area and Kalinda. Cassiterite-wolframite-bearing quartz veins at Kalinda occur together with vertical tin pegmatites, which as led to the belief that the vein-forming fluids were a late hydrothermal phase of the tin pegmatite magmas. Copper

Economically significant showings of copper are restricted to the mafic sequence of the Malaputese Formation and have been exploited at the Gwaai River Mine and a number of smaller prospects. The ore horizons are thinly bedded quartzites, calcsilicate rocks and a suite of grey cordierite-rich rocks of variable composition. The stratiform deposits of disseminated sulphides are thought to be of pre-metamorphic age and volcanogenic in origin (Bahnemann and Lockett, 1979). Mica

Scrap quality muscovite together with a limited quantity of book mica have been quarried from a number of late stage, cross-cutting tourmaline pegmatites in the Tshontanda Belt. The muscovite occurs chiefly in marginal quartz-mica zones and adjacent t o the quartz cores of large, zoned pegmatite bodies.

720

Lead Hydrothermal quartz veins mineralized with argentiferous galena occur as breccia and shear zone infillings within the Inyantue Formation at the Elbas Mine. The host rocks are veined biotite schists and migmatites containing a high proportion of calcsilicate bands. Several smaller copper-lead showings occur in the Inyantue and Malaputese Formations.

Fluorspar

A number of quartz-infilled fault planes contain small deposits of fluorspar. Sulphides of zinc and lead, silver and tin ores are notably lacking. The deposits are generally present in granitic rocks adjacent to outcrops of Karoo sediments, and the mineralization is thought t o be of post-Karoo age. GEOCHRONOLOGY AND CORRELATION

Radiometric ages determined prior to 1968 are summarized by Vail and Dodson (1969). Lepidolite from Lutope gave an Rb-Sr age of 2100 ? 20 Ma. Model lead ages of 1 3 1 0 f 39 Ma and 1250 f 39 Ma were determined on galena from the Elbas Mine. Galena dated at 1360 f 38 Ma probably also originated from the Elbas Mine, although the reported locality is the Kambanje Mine. On the basis of these lead ages Cahen and Snelling (1966) described the western metamorphic belt of Zimbabwe as a probable extension of the Irumide belt in Zambia, which is of Kibaran age. Priem et al. (1971) have determined an average Rb-Sr age of 2150 k 100 Ma for three samples of granite from the Kamativi area. Biotite from these granites yields an Rb-Sr age of 940 k 30 Ma and a K-Ar age of 988 20 Ma. This later thermal event can probably be related t o the intrusion of the tin pegmatites, which have an Rb-Sr age of 990 15 Ma. To summarize these results, it would appear that the supracrustal rocks were affected by a high-grade metamorphic event before 2000 m.y. ago, the probable age of the intrusive granites and gneisses. The tin, wolframite and lead mineralization is associated with a thermal event which occurred about 1000 Ma ago. Lithological similarities of Kamativi and Malaputese Formations to Piriwiri and Lomagundi Groups respectively were noted by Watson in his description of the Kamativi-Lubimbi area. While the linear supracrustal belts might possibly continue along strike t o correlate with Piriwiri rocks of the Kariba area, the nearest Lomagundi sediments lie some 200 km t o the east of the Dett area and any correlation with these latter and the Malaputese is, t o say the least, highly speculative.

*

*

7 21

E. Precambrian geology of Zaire and Zambia

*

Zai're and Zambia t o a large extent form a distinct geologic province, but they cannot be considered completely in isolation, and this description will therefore embrace adjacent parts of various surrounding countries where applicable. The Precambrian geology of the Zaire-Zambia area is dominated by two major units. The so-called Basement Complex, distributed mainly around the edges of the area as more-or-less isolated shields, contains a significant proportion of ancient rocks. Within this ring and in part separated by shield areas are the major occurrences of the Middle t o Upper Proterozoic Kibara-Katanga formations. Although substantial areas are covered by younger beds, the outline and structure of the Precambrian formations have been fairly well established (Fig. 10.13). SHIELDS

The shields that encircle this area vary widely in their present appearance, and in their structural and geochronological history and relationships. They generally contain recognizable Archaean formations, some younger formations, mainly Lower Proterozoic, and abundant gneisses. Some have been preserved relatively unaltered, whereas others were caught up in and affected to a greater or lesser degree by later tectonic or tectono-thermal events, so that the original nature and age of the rocks are obscured. The Angola, Kasai, Gabon, Northeast Zaire, Tanganyika, and Rhodesian shields are in part of the Archaean age and consist largely of granitic gneiss with recognizable Archaean greenstone (schist) belts of volcanic and sedimentary rocks; most shields seem to have been broadly stable since the Archaean. The Kasai shield in southwest Zaire and northeast Angola contains gneisses older than 3280 Ma, an Archaean gabbro-harnockitic complex, and the Archaean to Lower Proterozoic metasedimentary Lukoshi and Luiza complexes, and the probably somewhat younger Lulua sedimentary-volcanic complex. The Angola shield, separated from the Kasai shield by a strip of thin Karoo formations that follows the trend of the West Congo syncline, contains a number of small Archaean greenstone belts and other formations of Lower t o Middle Proterozoic age within a gneiss terrane. The radiometric ages reported are dominantly around 1300 t o 1700 Ma, suggesting that the Angolan shield had a more complex history than the Kasai shield, and was involved in Middle Proterozoic tectonic events. The gneissic Gabon shield, the Massif du Chaillu, lying to the northwest of Zaire, is similar to the Angolan shield but exhibits radiometric ages from 1700 Ma into the Archaean. In the north, from Cameroon t o the Central African Republic and northern

* By F. Mendelsohn.

7 22 Zaire, lies an east-west belt of Pan-African age (500 Ma) gneisses, which undoubtedly includes older formations, some of sedimentary and eruptive origin. In northeast Zaire, a small shield area contains the West Nile Complex of ancient gneisses together with amphibole gneiss, schist, quartzite, and limestone, and the younger Kibalian formations. The West Nile Complex, probably 3500 Ma old (Cahen and Lepersonne, 1967, p. 160) stretches east and north into Uganda and the Sudan; the Bomu formations t o the west are similar to and probably correlated with the West Nile Complex. The Kibalian is at least 2050 Ma old and seems to be Archaean in habit, occurring like the greenstone belts as elongated patches in granite, but Cahen and Snelling (1966) classify it as Lower Proterozoic, as does the latest geologic map of Zaire, and correlate it with the Buganda-Tor0 of Uganda. I t consists of amphibolites that probably represent metamorphosed lava, overlain by various schists containing iron formations. In the northeast, this shield has been affected by the Pan-African episode. In the east, beyond the narrow Kibaran zone, lies the East African or Tanganyika Shield, which extends southward to southern Tanzania, where it butts against the Ubendian zone; beyond lie only scattered patches of probably Archaean rocks in eastern Zambia, in part forming what has been described as an extension of the Tanganyika Shield. This shield contains metamorphosed sedimentary Nyanzian rocks and the younger but still Archaean Kavirondan arenaceous sediments, as well as patches of younger rocks, and is bounded on the east and southeast by the Pan-African Moqambique belt; radiometric ages range from Archaean to 1000 Ma. In the Copperbelt area and northwestern province of Zambia, the Lufubu mica schists, quartzites, gneisses, and other rocks are Archaean t o perhaps Lower Proterozoic in age, but occurrences are scattered and do not form part of recognizable shield. In the south, the ring of shields is completed by the Rhodesian Shield, the best known and most “typical” of the group. The Lower Proterozoic is not well represented, formations of this age within the shields being in the main fragmentary and even of doubtful correlation; apart from the Ruzizian, there are few whose ages have been established. The 1800 Ma Ruzizian fold belt extends from eastern Zaire and Burundi, where it transects the Kibaran, southeast across Lake Tanganyika into southern Tanzania, where it is known as the Ubendian, through the far northeastern corner of Zambia, and into Malawi; the general stratigraphy of the Ruzizian as given by Cahen and Lepersonne (1967) is outlined below: Upper Ruzizian

( f

Lower Ruzizian

Schists, phyllites, amphibolites, quartzites Phyllitic conglomerate or biotite schist Mica schists and phyllite, often graphitic, with crystalline limestone Quartzites, arkoses, dark phyllitic schists Schists, often graphitic, quartzites.

723 This broad fold belt of mildly metamorphosed sedimentary rocks forms the southwestern boundary of the main Tanganyika shield, and is partly overlain and obscured in Zaire by the Kibaran formations, though here and in adjoining countries a number of patches emerge from beneath the Kibaran. A number of granitic intrusions postdate the Ruzizian orogeny by up to 200 Ma. The Buganda-Tor0 formations of western Uganda are Lower Proterozoic. The mineral deposits belong t o the broad group found in shield areas of the world, mainly gold (both as veins and associated with iron formation), massive sulphide c o p p e r z i n c and coppernickel, and enriched iron formation, though in many parts knowledge is poor and exploration to date rather meagre. In Angola one iron mine and a small gold mine have operated in an Archaean greenstone belt in the south. Gabon, Cameroon, and Central African Republic have not exploited nor intensively explored the ancient shield rocks. In northeast Zaire the old Kilo-Mot0 goldfields produced significant amounts of gold from mines in Kibalian greenstones and iron formations, and workable iron formations exist, but development has been hindered by the remoteness of the region. Similar deposits have been mined in greenstones of the East African Shield in Kenya and Tanzania; the Macalder mine in Nyanzian rocks in Kenya, though enriched in gold near the surface is a copper (zinc, gold, silver) deposit. In Uganda the Kilembe coppercobalt deposit is in Lower Proterozoic (2200 Ma) Tor0 metasediments, probably pre-Ruzizian (Bird, 1968). In Zambia the ancient rocks have been rather unproductive, though some base metal deposits have been found in recent years. In Zimbabwe gold mines in Archaean greenstones and iron formation and copper-nickel and chromite deposits in greenstones are well known, as are numerous small copper deposits in Lower Proterozoic rocks. KIBARAN

The Kibaran is one of the more important formations of Central Africa, though it has not received a great deal of attention in recent years. It is closely associated with the Katangan in both time and space, being the latter’s immediate precursor and commonly lying close to it wherever they occur (Fig. 10.13). Over substantial areas the two were laid down in the same broad depository, with the Kibaran forming the framework of the Katangan depository, in part providing the source of the sediment (Cahen, in Bartholomk, 1974), and still outlining some of the major Katangan occurrences. There is a broadly similar spatial relationship between formations of similar ages in Southern Africa (Kheis/Khoabendus-Damara/Mdmesbury), and similar relationships can be detected in Brazil (Araxa-Bambui; Dechow et al., 1976) and Australia (Willyama-Adelaidean). Since the two together form a substantial and valuable mineral province in Zambia and Zaire, and

724 mineral deposits are known in some of these correlatives, the further investigation of the Kibaran in conjunction with the Katangan could be rewarding both scientifically and economically. The Kibaran consists dominantly of finer clastic sedimentary rocks, and in the type section in southern Shaba Province of Zaire (Figs. 10.13 and 10.15) where it has long been known and investigated (see Cahen and Lepersonne, 1967), the stratigraphy can be summarized as follows:

4 3 2 1

Rock types

Probable thickness

Limestone: dolomitic, siliceous, stromatolitic Schist, graphitic; phyllite; quartzite Quartzite; conglomerate, phyllite, basalt Phyllite, quartz phyllite, sericite schist, chlorite schist; local conglomerate, limestone, rhyolite

500 m 3000 m 3000 m 4000 m

This belt extends north-northeast into Rwanda and Burundi, where it is known as the Burundian, with a succession broadly similar t o the type Kibaran, and still farther northeast becomes the Karagwe-Ankolean of Uganda. Southwestwards, the Kibaran extends t o the border between Zaii-e and Zambia, and possibly into eastern Angola, before all the Precambrian rocks are covered by younger formations. According to Cahen (in Clifford and Gass, 1970), Kibaran sedimentation started after the Ubendian and related orogenies, probably around 1900 Ma and certainly after 2100 Ma and before 1800 Ma, and ceased with the main Kibaran orogeny at 1250-1300 Ma. It formed an important east-northeast structural trend, along which tectonic or igneous activity continued intermittently till the Pan-African event at about 500 Ma, and vertical movement of the Kibaran helped control Katangan sedimentation. Low-grade regional metamorphism is common throughout this belt, the argillaceous formations being represented by phyllite, shale, slate, and argillite; in many areas the metamorphism increases towards the base, where muscovite-biotite schists, amphibolite-grade rocks, and even migmatitic gneiss can be found. The formations have been affected by thrusting and by folding with a north-northeast trend, swinging to northwest in the north; in Shaba folding dies out westwards towards the Kasai shield. The Karagwe-Ankolean rocks are characterized by arenas where erosion has removed the crests of domal areas formed by the culminations of relatively gentle cross-folds. In Zambia, the main Kibaran occurrence is the northeasterly trending zone extending from the Zimbabwean border in the south to Malawi in the northeast; it includes the Muva and Kalonga metasediments and associated

725

Shields

0

I

--.500

1000

Kilometres

Fig. 10.13. Geological map of Zaire and Zambia and adjacent areas of Central Africa.

migmatites and gneisses of the Irumide fold belt, the Mulungushi, Rufunsa, and Sasare quartz ite-m etavolc an ic-m arble-graph ite schist-p elite-migmatite sequences, and other metamorphic rocks in the south (Drysdall et al., in De Wetter et al., 1972). The Muva of the Zambian Copperbelt, consisting of quartz-biotite schist, characteristic orthoquartzite, and other quartzite and conglomerate, has a somewhat varied and erratic trend though it is

726 locally conformable with the overlying Katangan. The higher grades of metamorphism in Zambia are accompanied by folding and faulting of varying complexity; in the Irumide fold zone two fold/fault zones are present. The flat-lying Plateau Series of northeast Zambia, north of the Irumide belt, consists of a coarse clastic lower unit and a finer-grained arenaceous to argillaceous upper unit; in one place the lower unit is folded and metamorphosed and the upper unit is little disturbed and unmetamorphosed, suggesting a time interval between their deposition. Indications of a postUbendian but >1800 Ma age for the lower unit (Cahen and Snelling, 1966), and of >1390 f 70 Ma for unmapped Plateau rocks (Drysdall et al., in De Vletter et al., 1972), indicate a Kibaran age for this sequence. The formations of the West Congo geosyncline, stretching for 1200 km from northern Angola through Zaire, Congo Republic, and into Gabon, have been the subject of considerable debate, and many opinions have been offered on their age, identity, and correlation. There has been general agreement that the formations from the lower tillite up are West Congolian (Katangan), but there have been wide divergences of opinion regarding the formations below the tillite. In Angola, Stanton et al. (1963) include all the clastic sedimentary formations above the basement complex as the Sansikwa series, forming the basal part of the West Congo system. Farther north there is general agreement that the Sansikwa or Monts Bamba, equivalent to the upper part of the Angolan Sansikwa, forms part of the West Congo. In Zaire, Cahen and Snelling, (1966) classify the quartzites, black schists, mica schists, and acid volcanics forming the upper part of the formations below the Sansikwa as the Mayumbian, equivalent t o the Kibaran, and the lower, conglomerates, quartzites, amphibolites, and mica schists, as the (pre-Ubendian?) Zadinian, the unconformities above the two units being related t o their respective orogenies; in the Congo Republic and Gabon several variants have been suggested (described in Cahen and Lepersonne, 1967), but a detailed discussion here is not warranted. In Brazil, in what was probably an adjacent depositional basin in pre-drift Katangan times, the Bambui represents the Upper Katangan (Isotta et al., 1969), and there was non-deposition of Lower Katangan formations (Dechow et al., 1976). Considering the broad conformity of the West Congo and Mayumbe (and Zadinian) formations throughout the zone, it seems likely that sedimentation took place throughout in a single depository that persisted for a long time, but may have here followed a course rather different t o that in the east, but perhaps approaching that in Brazil. This suggests a Kibaran age for the Mayumbian, with the Zadinian rather earlier; the radiometric dates of 1950-1100 Ma available in the Zadinian and Mayumbian (Cahen and Snelling, 1966) are consistent with this interpretation. However, the possibility that the Mayumbian is wholly or partly Katangan cannot be discounted. Farther north, several patches of probable Kibaran-age rocks, the Franceville, Lower Dja or Sembe-Ouesso, and Liki-Bembe formations, occur around

727 the periphery of the Katanga basin. These are sedimentary rocks, mainly sandstones and shales, with some limestone, dolerite, and basic volcanics; the possibility tht bhese rocks too are in part Katangan must also be considered. In general, the Kibaran has been relatively stable and has not been strongly dtformed nor highly metamorphosed, though in places the effects of metamorphism, folding and faulting are notable, commonly related t o syn- and post-tectonic granitic intrusions, and it has undergone uplift in the main zone during Katangan sedimentation and tectonics (Cahen, in Clifford and Gass, 1970). The main Kibaran zone forms an important tin-tungsten-tantalumniobium province. The tin-tantalum-niobium pegmatites tend t o occur in Kibaran and Ruzizian formations associated with 1000 Ma granite, commonly in domal areas; tin and tungsten occur in quartz veins rather than pegmatites and they also tend t o be separated from the pegmatite zones. Stratiform tungsten concentrations are known in Middle Burundian and Karagwe-Ankolean carbonaceous shales, in which tungsten has been locally remobilized into small veins, and similar concentrations are understood t o occur in Zaire. In Zambia, these minerals are unknown, but small copper deposits have been found in Muva schist in the Copperbelt, and it has been suggested that several large stratiform deposits in western Zambia are in probably Kibaran formations (Benham et al., 1976). In Gabon, the Mounana uranium deposit lies in Francevillian quartzite. KATANGAN

Distribution The Katangan dominates the Precambrian geology of this part of Africa in three major occurrences, each of different character (Fig. 10.13). Covering much of Zaire is the central Congo basin, consisting of fairly flat-lying formations that crop out peripherally around the central area of younger beds; continuity beneath the younger formations has been established, in part by drilling (Cahen and Lepersonne, 1967). In the west, trending southeast, is the West Congo “geosyncline”, which is probably connected t o the Congo basin beneath the cover. In the southeast, separated from the Congo basin by the main Kibaran massif, is the type area of the Katangan in the Copperbelt of southern Shaba province (Zaire) and Zambia, from where Katangan formations extend northeast into Shaba to form the “Golfe du Katanga” (Francois, in Lombard and Nicolini, 1963), and southwest through Zambia, southeast Angola, and northwest Botswana, t o finally become the Damaran of South West Africa (Namibia).

728

Stratigraphy The following stratigraphic subdivision of the type Katangan is synthesized from the variety of existing subdivisions, and adapted so as to serve as a general framework to which other occurrences can be most readily fitted. Some of the other names used and some marker units are also shown.

’ Upper Kundelungu Calcaire Rose Petit Conglomerat (tillite) Upper Katangan (Kundu1ungu)c Lower Kundelungu \

Kakontwe limestone Grand Conglomkrat (tillite)

In addition there is a broad three-fold lithologic subdivision of the Katangan in all these areas, though the thickness and stratigraphic position of the divisions vary somewhat, and they tend t o cross chronostratigraphic horizons, suggesting some differences in the timing of sedimentation in the different areas and/or broad facies changes. At the base is the lower clastic division, arenaceous and conglomeratic, but also argillaceous in places; the middle division is characterized by carbonates, generally dolomitic, that are accompanied by argillaceous (carbonaceous) and glacial formations; the upper clastic division is arenaceous/argillaceous. Volcanics and intrusive rocks are rare, but where present occur in the lower part of the succession. Stratigraphic sections through the Katangan formations in the various areas are shown in Fig. 10.14. The detailed correlation is based mainly on the two tillites or tilloids, which though not everywhere true tillites, have undoubted glacial affinities (Binda and van Eden, 1972), represent a widespread glacial period in Central Africa (Haughton, 1961), and are acceptable as chronostratigraphic markers (Van Eden and Binda, in De Vletter et al., 1972). Over most of the area the limits of the Katangan are clearly defined, but in places there are doubts about the limits and correlation, and in places only a portion of the sequence is preserved. The West Congo correlation has been discussed, and although this is not yet fully established, the Sansikwa has been included as Lower Katangan. Farther north, tillite correlated with the Katangan lower tillite overlies the Liki-Bembe formations; it is not clear whether in this area the formations immediately beneath the tillite are

729 Katangan, or whether the Lower Katangan was not deposited here and the tillite overlies the Kibaran. In the north, the Lindi-Ubangi successions of Upper Katangan formations rest on more ancient schists and granites; under the Congo basin, red sandstones have been correlated with the upper SchistoGreseux of the Lindian. In the southeast part of the basin the Bushimay represents Lower Katangan (Cahen, in Bartholomk, 1974). In the Golfe du Katanga in northeast Shaba, Lower Katangan formations emerge from beneath the extensive Upper Katangan (Kundelungu) and extend t o Lake Tanganyika as the Marungu formations. In northwest Zambia highly metamorphosed rocks in the Solwezi area overlying basement rocks are either Kibaran or Lower Katangan but probably older than the Lower Katangan of the Copperbelt (Benham et al., 1976). Limited direct data are available t o fix the time of deposition of the Katangan: Cahen (in Bartholomk, 1974) quotes ages of 840 and 888 Ma for the formation of microcline in veins cutting Lower Katangan rocks in Zambia and Shaba, and 938 5 15 Ma for lava a t the top of the Bushimay, which fix a minimum age for the Lower Katangan; the oldest age established in the Bushimay, from syngenetic lead, is 1055 Ma, and in the Musoshi area 1040 Ma. Deposition was clearly initiated after the Kibaran orogeny (12501300 Ma), and Cahen suggests that Bushimay sedimentation may have started before 1200 Ma, but it seems likely that it was around 1000 Ma before Katanga sedimentation was general, and locally, as where Lower Katanga beds are absent, even later. Deposition of the Katangan formations continued with at least local interruptions until the major Lufilian (i.e. Katangan) orogeny at 640 Ma.

Structure The Katangan formations peripheral t o the gently depressed central Congo basin with its cover of younger rocks are generally flat-lying or dip at low angles, though locally they are folded and faulted (Cahen and Lepersonne, 1967). The West Congolian has a strong southeasterly trend (about S55"E) and folds related t o the main orogeny are somewhat variable but have a general south-southeasterly trend (Stanton, 1963, p. 123; Cahen and Lepersonne, 1967). The Bushimay trends southeast and open folds also trend southeast, though some tight folds close t o the Kibaran parallel its northeast trend. In the Congo basin the folding and metamorphism are slight, but both increase somewhat in the West Congolian, where they also increase downwards through the succession. In the Shaha-Zambia area, the dominant Kibaran-Katangan contacts and therefore the limits of the Katangan trend northeast and north-northeast (Fig. 10.15). The important Lufilian arc, within which occur the ore deposits of the Shaba-Zambia copperbelt, has long been recognized and has been

730 documented by Van Doorninck (1928) and Garlick (in Mendelsohn, 1961); Garlick suggests that the nature of the arc is related t o the two bounding Kibaran massifs, which is analogous to the structure of part of the West Congolian between older massifs (Stanton, 1963, p. 123). The Lufilian arc has been accepted as the dominant Katangan and Late Precambrian t o early Paleozoic structure of Zambia (Cahen and Snelling, 1966; Clifford, in Clifford and Gass, 1970; etc.), but both Garlick and Brock (in Mendelsohn, 1961) show important northeasterly and north-northeasterly trends around the Copperbelt, parallel to nearby major Kibaran boundaries. More recent work, such as that by De Swardt and Drysdall (1964), shows that the arcuate trend is largely confined to the Copperbelt area, and is separated from the Moqambique (Pan-African) belt by the northeast-trending Irumide (Kibaran) belt, though the latter has suffered some overprinting (Drysdall et al., in De Vletter et al., 1972). In the southern part of the Lufilian arc, the basement domes and anticlines around which the Katangan formations are draped represent a lower structural level and exhibit trends that parallel the nearest major northeast and north-northeast Kibaran trends of the Golfe du Katanga as well as those of the Lufilian arc; farther south structural trends in the Katangan are random t o varied (Fig. 10.15; Garlick, figs. 27, 28, in Mendelsohn, 1961). It seems that the broad structural significance of the Lufilian arc has been somewhat overemphasized, and it is proposed that the primary structural trend is northeast t o north-northeast, and is related t o and controlled by the Kibaran; age dates in the 440-800 Ma range along the Kibaran extension of this trend to the north (Hurley and Rand, 1969) support this interpretation. It would seem that this Katanga-Kibaran trend is the locus of a less intense, possibly even discontinuous Pan-African belt that parallels the main (Moqambique) belt to the east and forms an extension of the Damaran zone. In Zambia, the extensive folding of the Katangan during the Lufilian (Katangan) orogeny was accompanied by metamorphism that was generally low grade but reached medium to high grade in places. Northwards into southern Shaba the Lufilian arc is marked by disruption and gravity sliding as well as folding elongated along the arc, but only mild metamorphism; in northern Shaba the Kundelungu (Upper Katangan) is flat-lying and it seems that the two Kibaran massifs protected this wedge from the tectonism. The Lower Katangan (Marungu) formations emerging from beneath this Kundelungu to form the northern end of the Golfe du Katanga near Lake Tanganyika (Fig. 10.15) are somewhat disturbed and show low-grade metamorphism. Attention was first drawn to the disparate nature of structures associated with the Pan-African thermotectonic event by Kennedy (1964). What Kennedy described as the normal type of orogeny, involving geosynclinal (or intracratonic) sedimentation and progressive regional metamorphism, as for the Katangan, has produced an effect quite different from that in the

Fig. 10-14. Stratigraphic sections through Katanga formations, compiled from various sources in the list of references. The Grand Conglomerat and Petit Conglomdrat glacial formations are shown as chronostratigraphic markers and are used as the basis for detailed correlation.

0Katangon

Kibaran

Ruzizian

Trend lines

Fig. 10.15. Map of portions of Zambia and of the Shaba Province, Zaire, showing geologic units and structural trends. The Lufilian arc forming the southern part of the Golf du Katanga, the random trends in the Katangan in the south, the northeast and east-northeast trends of the Kibaran formations, and the cmfining of the Katangan by the Kibaran are evident. The Central African Copperbelt stretches from Kolwezi through Shituru, Lubumbashi. and on to Ndola in Zambia.

736 characteristic gneiss belts, such as the MoCambique belt, with a high thermal effect and an age of around 500 Ma. It has been suggested (Mendelsohn, in Lister, 1973) that the two are linked by a plate-tectonic model, with the gneiss belts forming in linear zones of high heat flow and “intra-cratonic jostling” after plate collisions and as a prelude t o subsequent separations, many of which took place along Pan-African gneiss belts (e.g., Fyfe and Leonardos, 1973). Clifford (1967) found peaks in radiometric ages in Southern Africa at about 520 and 630 Ma, and suggested that these be called the Damaran and Katangan episodes respectively; tectonic ages of 600-650 Ma have been recorded by Cahen and Snelling (1966), and Cahen (in Clifford and Gass, 1970) suggests that there were separate phases of the Katangan orogeny at 710 Ma, 670 Ma, and 620-670 Ma. It is concluded that the period around 640 Ma represents the peak of the Katangan orogeny in most areas, and that the circa 500 Ma age is that of the Pan-African thermotectonic event. The Pan-African event gave rise t o the widespread gneiss belts of Africa (e.g., Clifford, in Clifford and Gass, 1970) that are largely separate from the Katangan orogenic belts, but in places this event followed along the trend of the Katangan orogeny, partly or wholly overprinting the older ages and events, as in and near the Lufilian arc and the West Congo geosyncline. The widespread nature of these events is emphasized by Hurley (1974), who proposes that the world-wide Pangeaic Orogenic System, incorporating and extending the Pan-African, is an assemblage of genetically related orogenies of considerable importance that occurred between 650 Ma and 250 Ma, and by Sawkins (1976), who concludes that there was a world-wide relative tectonic quiescence between 900 Ma and 700 Ma, contrasting with a widespread re-setting of radiometric clocks between 700 Ma and 500 Ma. Sedimentation of the Katangan was in the main of the intra-cratonic basin type, in tectonically controlled depressions largely confined within Kibaran rocks that formed resistant bounding massifs. The overall geologic environment was relatively quiescent, with mainly shallow-water sedimentation into slowly subsiding basins of oval t o elongate shape, under generally transgressive conditions in the lower part. Volcanic activity during deposition was minimal, and later granites are rare t o absent, as more and more “intrusive” bodies have been found t o be pre-Katangan but with indications of some remobilization; basic intrusives have been found in the Lower Katangan (Fig. 10.14). The widespread continental t o shallow-water reworked glacial beds and the presence of Red Beds in the Upper Katangan suggest some emergence after about 750 Ma. Cahen (in Clifford and Gass, 1970) suggests that vertical movement of the Kibaride belt controlled Katangan sedimentation and that there were several phases of movement and sedimentation during Upper Katangan times.

737 Mineral deposits The copper (cobalt, uranium, lead, zinc, vanadium, cadmium, silver) deposits of the Shaba/Zambia Copperbelt form one of the world’s major mineral fields, and although a number of occurrences and even a few mines are known in Katangan formations in other areas, very few deposits of comparable size are known outside the Copperbelt. By far the greatest production is copper, but several of the other elements make important contributions to the total metal production. The stratiform copper deposits lie in the Lower Katangan clastic sequence, the Lower Roan formations, though there are notable differences between the Shaba type deposits, occurring over most of the Shaba part of the belt, and the Zambian type, occurring in the southeastern part of Shaba and the Zambian Copperbelt. The Zambian type deposits lie in the middle part of the lower clastic division, the Lower Roan, though locally these lie close to the basal unconformity. The main type of deposit here is the “Ore Shale”, consisting of a silt-size calcareous argillite, locally carbonaceous, with substantial amounts of impure dolomite, representing algal mats (Mendelsohn, in Walter, 1976); locally ore spreads into adjacent arenites. Associated are stromatolitic bioherms or patch reefs (Clemmey, in Bartholomk, 1974), the bioherms being barren of copper and surrounded by richer shale. The deposits lie within a zone about 2 km wide and 100 km long, and pass down the (depositional) slope into a broader pyritic argillaceous zone. In the general vicinity of many of these ore deposits are the small rich “Footwall” deposits, occurring below the ore shale and immediately above the basal contact, nestling against basement highs or in local depressions. In the east are the arenaceous bodies of the Mufulira type, consisting of quartzite, carbonaceous greywacke, and, locally, impure dolomite. Superposition and repetition of orebodies is common, as at Mufulira. In Shaba there are two major mineralized formations: a lower siliceous dolomite, which is at least chemically similar to the impure dolomite of the Zambian Copperbelt, and an upper dolomitic shale that is virtually identical to the Ore Shale. Between the two is a dolomitic rock containing biohermal and biostromal stromatolitic formations and generally unmineralized except for included shale laminae. In the Kolwezi area other higher shaly beds are mineralized. Because of the extensive movement on bedding faults, direct correlation between the two areas has not been possible, and in the past it was considered that the more dolomitic mineralized zone of Shaba was stratigraphically equivalent to the dolomitic Upper Roan of Zambia; in recent years a similar stratigraphic succession t o that of the Zambian deposits has been found (Bartholomk et al., in Amstutz and Bernard, 1973) and a preference has been expressed for the idea that the ore-bearing formations throughout the mineral province are stratigraphically equivalent (Francois, in Bartholomk, 1974). The main copper minerals are chalcopyrite and bornite,

738 and there is a well-defined mineral zoning, in places demonstrably parallel to the depositional shore line, of chalcocite, bornite, chalcopyrite (linnaeite) pyrite; secondary alteration is variable, and many of the northern deposits have consisted largely of carbonate and other “oxide” minerals. Stratiform concentrations of copper have been found at various levels through the Katangan in the general area, such as the top of the Upper Roan, the Mwashia (Lefebwe, in Bartholomb, 1974), and the Kundelungu (Ellis and McGregor, 1967), and though in some cases minable, they are generally isolated and do not approach the Copperbelt deposits in size or value. The major Lumwana deposits in kyanite-grade schists west of the main mineralized area, have not been directly correlated, and have been suggested as being Lower Roan or pre-Katanga (Kibaran) in age. Cobalt is associated with many of the copper deposits, to the extent that the area is a significant cobalt province and a major cobalt producer. Uranium accompanies the copper in many deposits in slight though detectable concentration, and a small stratiform uranium deposit was found in a barren near-shore sandy facies of the ore shale at Nkana (Garlick, in Mendelsohn, 1961). In Shaba, the major uranium deposit at Shinkolobwe, and others nearby, occur as veinlike transgressive bodies within the normal ore-bearing formations, accompanied by nickel, cobalt, and other sulphides. Among the other elements found as separate stratiform bodies or accompanying copper concentrations, though not so far in economically recoverable amounts, are nickel, molybdenum, and tungsten. In carbonates of the Upper Roan at Broken Hill are transgressive pipe-like bodies of massive sulphides containing lead, zinc, and vanadium. The famous Kipushi (Prince Leopold) deposit, from which copper, zinc, cadmium, germanium and lead are still being recovered, is a transgressive body occurring in the Kakontwe limestone formations along a major fault. In the Bushimay, stratiform lead deposits and minor amounts of copper, zinc, and iron sulphides occur in the lower formations, stratigraphically equivalent to the Lower Roan. The Lindian-Ubangian is virtually unmineralized. In the West Congo, copper, lead, zinc, and other mineral occurrences of both stratiform and veinlike character have been found in the MatadiKinshasa area, mainly in the carbonates of the Upper Shiloanga and SchistoCalcaire. Farther north, there are stratiform occurrences of copper, lead, zinc, and pyrite in clastic and carbonate rocks at various levels in the succession. In Northern Angola small stratiform deposits in the West Congolian have been worked at Bembe and Mavoia. In the eastern part of Angola, small stratiform copper deposits have been found in the Malombe formations, Lower Katangan and probably correlated with the ore formations of the Copperbelt; still farther t o the southwest in Botswana along the same trough, are deposits of the ore shale type in the Ghanzi formations. Also in Botswana, the Matsitama basin of metamorphosed sedimentary and volcanic rocks,

739 which could be Katangan age or much older, contains small stratiform copper deposits of the Copperbelt type. The reason for the intense concentration of ore deposits along the Lufilian arc, contrasted with their relative sparsity elsewhere, particularly outside the Katanga-Damara trough, is not clearly understood. The stratiform deposits formed early in the history of the rocks, during sedimentation and diagenesis, and during the Katangan orogeny and Pan-African event these were modified and some other deposits were formed, but there was no general redistribution of metallic or other elements. Thus the presence of abundant Lower Katangan formations and ore deposits along the Lufilian arc suggests that the arc could perhaps be related to an ancient basin-edge structure; as suggested earlier, the arc is a second order structure, but is nevertheless highly significant as far as the ore deposits, and perhaps exploration, are concerned.

F. The Zambezi Metamorphic Belt in Zimbabwe

*

The Rhodesian craton is surrounded in the south, east and north by zones of high-grade metamorphic rocks that have undergone intense and repeated structural deformation. The Zambezi or Northern Metamorphic Belt may be subdivided into a northeast trending zone which is described by Lockett (this chapter, part D) in the northwest of Rhodesia and which is unaffected by the -500 Ma Pan-African Event. This zone is largely covered by considerably younger unmetamorphosed Karoo and Sijarira sediments and only emerges to the south of Kariba where it intersects the generally east-trending Zambezi Belt that is bounded by the fault-defined Zambezi Escarpment in the north and by the Rhodesian craton in the south. This portion of the Zambezi Belt can be traced for more than 450 km from southern Zambia to its junction with the north-trending Moqambique Belt in the northeast of Zimbabwe. The Moqambique Belt is coeval with the Zambezi Belt in so far as both have been metamorphosed by the -500 Ma Event. Gneisses are also exposed as inliers in the Chewore area and it should be realized that a large proportion of the belt is covered by Karoo sediments north of the escarpment. The Zambezi Belt is not a mobile belt in the same sense as the Limpopo Belt. .It is far more complex, involving more lithological units that have experienced a long history of metamorphism and deformation up until the latest Katangan and Miami metamorphic events ?800 Ma and 400 to 650 Ma ago, respectively. The apparent easterly Zambezi trend was probably an initial Archaean trend, but later deformations have resulted in both northeasterly and northwesterly imprints across the belt.

* By T.J. Broderick.

7 40

741 The Rhodesian Basement Complex consists largely of the ancient granite and gneiss terrane which incorporates the Early Precambrian greenstone belts of the cratonic area. Basement Complex rocks are recognized in the east and central portions of the Zambezi Belt where this traverses the northern portions of the Sipolilo and Darwin schist belts and the foliated granites in the Pfungwe Tribal Trust Land. Much of the eastern section of the belt is underlain by migmatitic tonalite gneisses which contain inclusions of possible Basement Complex age. In the west, however, no proven Basement Complex rocks have been recognized although the Urungwe Paragneiss forms the local basement over a wide area north and west of Karoi and is probably equivalent to the Escarpment Gneisses of Workman (1961) and Thole (1974). Numerous inclusions of possible Basement schist were noted in the postPiriwiri granite of the southern Urungwe which itself could represent remobilized Basement (Stagman, 1962). Age dating from the Rhodesian Basement Complex indicates a period of between 2600 and 3500 Ma, during which the initia! metamorphism took place along the easterly Zambezi trend. The Great “Dyke” was intruded approximately 2500 Ma ago after this first metamorphism (Wiles, 1968). Post-Basement sediments, now metamorphosed to gneisses, were subsequently deposited, probably in the form of greywackes derived from greenstone terrane, and have been recognized as Chiroti, Chipisa, Kariba and Chitumbi paragneisses in the west and as the Mazoe Schists east of Mount Darwin (C.B. Anderson, pers. commun., 1972). Loney (1969) determined whole rock Rb-Sr isochron ages for the Kariba and Chipisa paragneisses of between 2300 and 2500 Ma, using alternative decay constants. These paragneiss formations were metamorphosed prior to the deposition of the Piriwiri and Lomagundi Groups which overlie them west of the Shamrocke Mine and trend north to northeastwards across the Zambezi Belt. Age determinations on the Piriwiri-Lomagundi sequence, where it is unaffected by the Miami metamorphism, show that they are undoubtedly of mid-PreCambrian age and may be older than 2000 Ma. Younger ages of -1600 Ma indicate later phases of the Magondi Orogeny (Kirkpatrick, 1976). An age of 1170 k 40 Ma from the post-Piriwiri Badze Granite is taken as the minimum age of intrusion for the granite and sets a definite younger limit to the Piriwiri Group (Clifford et al., 1967). The metamorphic grade of the Piriwiri Group rocks was generally as low as greenschist facies at this time, but in the area east of Kariba and north of Karoi local charnockite and granulite development is recognized, as it is north of Sipolilo (Wiles, 1968) and north of Mount Darwin (Leitner and Phaup, 1974). The granulites are a product of polymetamorphism and were probably exposed in major anticlinal structures during the Katangan Episode although they only occur in pre-Katangan rocks all of which have subsequently undergone widespread granitization and metasomatic alteration, especially in the Urungwe District.

742 Loney (1969) was the first to describe the largely psammitic group of metasediments east of Kariba, which he dated at 800 f 30 Ma or 850 k 30 Ma, using alternative decay constants. This younger sequence, the Makuti Group, overlies the older paragneisses with distinct unconformity and can be traced northwestwards across the Zambezi River, where it is correlated with the Katangan Group of southern Zambia (Broderick, 1976). The Makuti Group is predominantly pelitic towards the base but in the upper formation pink, feldspathic paragneiss is intercalated with quartzite, calc-silicate rock, micaceous and aluminous schists and amphibolite of igneous origin. These rocks and the underlying formations were metamorphosed and repeatedly folded on northwest trending axes during both the Katangan and Miami metamorphic events. The metamorphic grade rises progressively from the greenschist facies near Gota Gota, west of Makuti southeastwards t o the upper amphibolite facies at Vuti Township. There was local granitization and pegmatites from the Miami area give numerous radiometric ages of between 400 and 650 Ma, so effectively setting an age for the Miami Metamorphism. North of Mount Darwin the Rushinga Group has much the same lithology and bears a similar relationship t o the underlying migmatitic gneisses as the Makuti Group (Leitner and Phaup, 1974). However, the two groups have not been directly correlated. The basal metasediments of the Rushinga Group are largely pelitic and the upper formation consists of dolomite, pink metaarkose and metapelites. The rocks reach the upper amphibolite facies of metamorphism and are characterized by complex east-trending dome and basin structures in the Chimanda Tribal Trust Land (C.B. Anderson, pers. commun., 1972). The Chimanda area was first mapped by Talbot (1973) and he recognized the Rushinga Group although he correlated it with the Umkondo Group after Johnson and Vail (1965), who managed t o trace the younger metasediments into the south-trending Mopmbique Belt. The Rushinga lithological sequence continues southwards along the Rhodesian border with Mocambique to the vicinity of the Senga Senga Dome, where on the Mocambique 1 : 250,000 geological map, sheet SUL-E-36/J (1968), it is depicted as overlying sediments typical of the Umkondo Group (V.R. Stocklmayer, pers. commun., 1976). With the uncertainty of correlation it is advisable t o retain separate group names for the latest metasedimentary sequences which have all been involved in the -500 Ma event.

Summary After Archaean deformation resulting in the overall easterly trend of the belt, the sedimentary rocks of the western part were metamorphosed and deformed on a northeasterly trend t o form the paragneissic basement on which the Piriwiri was deposited. During the Magondi Orogeny the Piriwiri rocks south of the Zambezi belt were folded on northeast to north-northeasterly axes, the former being reflected in the Urungwe District.

743 The dominant northwesterly and westerly trends of the overlying Katangan Makuti Group which produced complex interference patterns had little effect on the underlying formations. Likewise, there appear t o be no structural trends caused by the Miami Metamorphic Event in the western part of the belt. Farther east, t o the north of Sipolilo and Mount Darwin, early east-northeasterly trends were strongly overprinted by a polyphase deformation characterized by domes and basins and a strong easterly foliation of the Rushinga Group. Numerous -500 Ma pegmatite ages have been determined but little evidence has yet been found of the -800 Ma Katangan Event. Near the eastern border the trends swing southwards and appear to merge into the south-trending Moqambique Belt.

G. The Pan-Africanevent in Malawi and eastern Zambia * INTRODUCTION

The concept of a widespread thermo-tectonic (orogenic) event affecting the whole of the present African continent some 500 Ma ago was first put forward by Kennedy (1964). This major phase of orogenesis, basement reactivation, and granitization is widely reflected in the isotopic ages of the rocks concerned and Kennedy showed that the continent is made up of three old cratonic nuclei encircled by what he termed “Pan-African” mobile belts. The structural framework established by Kennedy was subsequently refined and modified by Clifford (1970) and others but the basic pattern of three cratons, the West African, Congo, and Kalahari, surrounded by zones affected by orogenies dated at -500 Ma remained unchanged. One of the fundamental structural features of Africa, first delineated by Holmes, is the Mozambique Belt, a roughly north-south zone of medium to high-grade metamorphic rocks that stretches 4000 km from southern Egypt, through the Sudan, Uganda, Kenya, Tanzania, Zambia and Malawi, to Mozambique. When originally defined, although it clearly truncated a number of older tectonic units, the age of this belt was not clear but it is now known t o be characterized by isotopic ages in the range 700 Ma to 400 Ma, i.e. part of the widespread Pan-African event. Although Holmes did not define the Mozambique Belt as more than a geographic entity, he implied that it was an orogenic zone, presumably characterized by the normal geosynclinal cycle of subsidence, deposition, uplift, folding, metamorphism and plutonism. This was subsequently extended by geologists working in Africa t o include the concept of a “Mozambique

* By K . Bloomfield. Published by permission of the Director of the Institute of Geological Sciences, London.

744

Fig. 10.17. Main structural elements of Malawi and eastern Zambia showing the 4 5 published K-Ar and Rb-Sr age determinations (after Carter and Bennett, 1973 and Snelling et al., 1972). M marks the Misuku Belt and N and S, separated by a dotted line, the northern and southern sub-provinces, Malawi Province, Mozambique Belt. Note the Rb-Sr isochrons quoted by Clifford (1974) are not included.

745 Geosyncline”, the axis of which followed the eastern coast of the continent. However, with more detailed work, it soon became apparent that, within Holmes’ Mozambique Belt, geological relationships were extremely complex and that it was both polyphase and subject t o reactivation. Vail(l965) was one of the first to point this out and indicated that the Belt contained facets representing old infracrustal gneisses and granulites, paragneisses possibly formed from sediments deposited in the “Mozambiquian” cycle, and marginal shelf deposits at least partly affected by orogeny. A detailed account of the lithology of the Precambrian rocks of Malawi was given by Bloomfield (1968). When this was written, only two-thirds of the country had been covered by regional mapping and the most that could be concluded was that the greater part was made up of “Basement Complex” paragneisses, schists, and granulites with a general north-south trend, an integral part of the Mozambique Belt. Cannon et al. (1969) used lithological, metamorphic and structural criteria to divide the ‘‘Malawi Province” of the Mozambique Belt into northern and southern sub-provinces. The line of separation was the extension into Malawi of the major ENE-trending Mwembeshi fracture zone (Fig. 10.17). However it was subsequently found that only the first criterion is really valid since areas in northern and central Malawi have been found to have broadly similar structural/metamorphic histories. The present position is summarized by Carter and Bennett (1973). LITHOLOGY

The metamorphic rocks of the southern sub-province of Malawi may be assigned to either the granulite or the amphibolite facies. The former includes banded charnockitic types, khondalites, and enderbites, associated with quartz-feldspar granulites, marbles and high-grade pelites. The more extensive amphibolite facies rocks comprise widespread semipelitic biotitegneisses, cafemic gneisses, psammites, pelites and locally abundant marbles and calc-silicate gneisses. Graphitic types are common and there are some ferruginous gneisses. Unusual aegirine and nepheline-gneisses also occur. No relict sedimentary features have been observed but the greater part of the granulites and gneisses is thought to represent metasediments - shales, marls and sandstones with interbedded limestones and ferruginous and carbonaceous bands. The outstanding geochemical features of the rocks are their relative richness in Mg and Ca and the dominance of Na over K. Si is much less abundant than in many other metamorphic terranes. Pre-kinematic orthogneisses include metadolerites and meta-anorthosites and there are many small masses of synkinematic basic and ultrabasic rocks. A distinctive feature is a province of late-kinematic perthitic syenites and gneisses and several infracrustal ring-complexes of associated syenitic and ultramafic rocks.

746 TABLE 10.V Main lithological differences between the northern and southern sub-provinces, Malawi Province, Mozambique Belt Northern sub-province

Southern sub-province

1. Charnockites rare

1. Banded charnockitic gneisses and granulites common 2. Marbles and calc-silicate gneisses common in both amphibolite and grknulite facies 3. Graphitic gneisses widespread and pyritiferous gneisses locally common 4. Cordierite rare

2. Marbles absent and calc-silicate gneisses rare 3. Graphite and pyrite absent 4. Banded cordierite-gneisses and granulites widespread 5. Kyanite rare 6. Meta-anorthosite absent

5. Kyanite common in pelitic rocks 6. Meta-anorthosites associated with granulites

A psammo-pelitic sequence, the Mchinji Group (in part, the Mwami Group) extends along the Malawi-Zambia border and rests on high-grade gneisses. In the northern sub-province calcareous rocks are rare and the schists and gneisses are mainly pelitic and psammitic. A sequence of psammites and pelites, the Mafingi Group, rests on granulites and gneisses and straddles the Malawi-Zambia border. I t may be equivalent t o the Mchinji Group further south. Sillimanite-bearing cordierite-granulites are common and there are broad zones of micaceous phyllonite resulting from the cataclasis of pre-existing rocks. The major lithological differences between the two sub-provinces are summarized in Table 10.V. STRUCTURAL HISTORY

Northern Malawi, with adjacent northeastern Zambia, is a region of great structural complexity, lying at the junction of three major orogenic belts: the NW-trending Ubendian Belt (2100-1800 Ma); the N- to NE-trending Irumide Belt, deformed during the Kibaran Orogenic Cycle (1600-1100 Ma); and the generally N-trending Mozambique Belt (700-400 Ma), (Fig. 10.17). Southern and central Malawi have generally been regarded as situated entirely within the Mozambique Belt. The true significance of the Pan-African event can only be appreciated in the wider context of the known Precambrian structural history of this complex region.

747

Northern Ma law i The structural history revealed by systematic mapping is shown in the composite Table 1O.VI. Four episodes of essentially plastic deformation and medium- t o highgrade regional metamorphism are assigned to the Ubendian orogenic cycle but there is evidence of pre-Ubendian granulites. The Nyika Granite was emplaced soon after a final Ubendian phase of shearing and retrograde metamorphism and its U-Pb zircon age of 1970 f 30 Ma sets a minimum limit for the Ubendian cycle in this area (Dodson et al., 1975). The Mafingi Group of sandstones grading up into shales was deposited on peneplained Ubendian rocks during a marine transgression. It is an excellent lithostratigraphic marker as both the sediments and the underlying gneisses were subsequently involved in four episodes of mainly brittle deformation under greenschist facies conditions. The oldest phase, F,, is marked by strong phyllonitization that Cannon et al. (1969) deduce to be of Kibaran age but the later phases are assigned t o the Mozambiquian orogenic cycle. A lengthy time interval between F, and F,, was implied since the main compressive stress field changed from east-west t o north-south. No metasediments of definite late Precambrian “Mozambiquian” age have been recognized. There is a similar picture where the Ubendian Belt trends SE into the extreme north of Malawi t o become the Misuku Belt (Ray, 1974) (Fig. 10.17). Here the Misuku Gneisses are divided into the Chambo and Songwe Gneisses, the line of separation being the major NW-trending Mugesse Shear Zone. The Jembia River Granulites, containing cordierite and sillimanite, may or may not be pre-Ubendian; their sharp contact with the Misuku Gneisses is assumed t o be tectonic. Three Ubendian fold phases on NW-trending axes have been identified in the Chambo Gneisses and one in the Songwe Gneisses, in both cases accompanied by amphibolite facies metamorphism (Table 1O.VII). Folding was accompanied by the intrusion of major granite bodies, now strongly foliated. Again the deposition of the Mafingi Group separates the Ubendian movements from the later Irumide (Kibaran) and Mozambiquian folding. The effects of the Irumide event are only evident in the Chambo Gneisses as open folds with N-trending axial planes accompanied by lithostatic metamorphism to produce corona-metadolerites but no marked textural changes in the gneisses. Four phases of deformation in the Chambo Gneisses and three in the Songwe Gneisses are ascribed t o the Mozambiquian cycle. In the latter the oldest phase is effectively dated by the intrusion of the Songwe Syenite Complex, an elongate mass of perthitic syenite with local weak F,, foliation, associated with foyaites and sodalite-bearing rocks. Initial phases of isoclinal and open folding on NW-trending axes associated with amphibolite facies metamorphism (unlike the lower greenschist facies

TABLE 1O.VI Structural history of Northern Malawi (except Misuku belt) (after Cannon e t al., 1 9 6 9 ; Dodson e t al., 1975; Fitches, 1 9 7 1 ; Kemp, 1 9 7 3 ; Thatcher, 1973a, b) _

_

_

~ ~~

Age (Ma)

700-

400

~~

~

~~

Orogenic cycle

Deposition

Mozambiquian

Deformation

Metamorphism

Plutonism

FI"

MI" - greenschist facies MI,, - greenschist facies M,I - greenschist facies

late pegmatites

FIJI

FII 1200-1100

F,

Irumide (Kibaran)

- 1300

- strong

cataclasis -kink bands - open folding of phyllonites - phyllonitization and augengneisses

MI

- greenschist,

locally amphibolite facies

~

_ _ __ _ _ . _ _ ~

Mafingi Group psammites and _ _ _ _ _ _ _ _._- _ pelites ________ ______-

__ F4

- - _ _ _ _ _ _ - _ _ - _ _ _ _ _ _ _ _ - - -- - - .-_ - open folding,

- - - -- - - - - - - - -- - - - - - - -- .-

M4

- amphibolite facies

M3

- amphibolite facies

-major banding and foliation parallel to axial planes - mineral banding

M,

- amphibolitegranulite facies

MI

- amphibolite-

earliest fabric' poorly developed

-

partly conjugate, followed by local shearing -variable folding, some cataclasis

granites, syenites, and nephelinesyenites epidotization ________-____-_-

Nyika Granite ( 1 9 7 0 k 30 Ma)

L

2100-1800

Ubendian

pelites, semipelites and psammites

F3

b. F,

?

?3000--2500

Pre-Ubendian

? sediments

and volcanics

granulite facies ? granulite facies

granites and mafic igneous rocks

TABLE 1O.VII Structural history of the Misuku belt, northern Malawi (after Ray, 1 9 7 4 ) Orogenic cycle

Chambo Gneisses Deformation

Songwe Gneisses Metamorphism

Deformation

Plutonism Metamorphism doleri tes

. . . . . . . . . . . . . .late shearing along SE-trending zones . . . . . . . . . . . . . . . . . . . . 'F8

F,,

- disharmonic

open folds

- open concentric folds

. . . . . . . . . . . . . . . . . .shearing . . . . . . . . . . . . . . . . . . . . . . . . . . 700-

400

F, - tight folds

Mozambiquian

with pyllonitization and shearing

I

l5 I F,

-1100

......

Irumide (Kibaran)

.....

-1300

M4 - greenschist facies

- open folds

- isoclinal folds F4 - open folds

...........................

FIII

-tight shear MIXI - greenschist facies folds

FII

-open to tight similar folds

M3 -lower amphibolite facies

MI,

-lower early syntecamphibolite tonic intrusion facies of Songwe Syenites (c. 685 Ma)

M, - lithostatic lamprophyres

..................................................................... Deposition of Mafingi Group -psammites and pelites

. . . . . . . . . . . . . . . . . . .. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . F3 - development of boudins

2 1 00-1800

Ubendian

F,

,F1

- isoclinal large scale structures -?

dolerites

MI -amphibolite facies

FI

-mineral banding/ foliation

MI

-lower syntectonic amphibolite basic and facies granitic rocks

. . . . . . . . t e c t o n j c contact.. ........................................................................ Jembia River Granulites deposition o f pelites and semi-pelites

..........................

4 A

a

750 further south) are succeeded by tight shear folding accompanied by intense cataclasis to form phyllonites and the Mugesse Shear Zone. This was followed by a period of open folding and by late NW shearing. It is clear that the Mozambiquian movements caused realignments of many Irumide structures and generated folds controlled by pre-existing Ubendian structures. In the Ubende area of Tanzania, the northwestern extension of the Misuku Belt, the most intensive shearing was associated with Ubendian tectonism. In northern Malawi, however, there is clear evidence that the main shearing took place during the Mozambiquian cycle and so this NW-trending zone has probably been a major line of weakness over a long period of time.

Zambia Northeastern Zambia is effectively divided geographically into two separate Precambrian terrains by the Luangwa Valley, a narrow trough infilled with Karoo sediments. The eastern area is an extension of the Precambrian of northern and central Malawi and has probably undergone a similar structural history (E.A. O’Connor, pers. commun., 1977). Amphibolite t o granulite facies metasediments were probably deformed during the Ubendian cycle and locally downgraded by cataclasis during the Mozambiquian. Polyphase deformation is common but available isotopic ages, mainly from pegmatites, range from 470 Ma to 380 Ma. Table 1O.VIII shows the structural history of an area in eastern Zambia, just west of the Malawi border, and is typical. The low-grade psammites and pelites of the Mwami Group in eastern Zambia may be correlated, in part, with the Mafingi Group but, as in northern Malawi, there is no evidence of later Precambrian sedimentation (Drysdall et al., 1972). The main NE-trending Irumide Belt lies to the west of the Luangwa Valley, flanking the cratonic Bangweulu Block. This 120 km-wide zone is made up of basal granitic gneisses succeeded by at least 10,000 m of banded metasiltstones and metaquartzites. Here the Kibaran orogenic cycle, known locally as the “Irumide Orogeny”, has been divided into five phases (Klinck, 1980). The granitic gneisses are syn- and late-kinematic with respect t o F1. A “Turnbide” phase (F,) resulted in recumbent isoclinal structures with N to NW-trending axes; the “Irumide” phase (s.s.), F3, is characterized by upright, open t o tight folds on NE to ENE axes. Traced southeast across the belt, away from the craton, the metamorphic grade increases from low greenschist to amphibolite facies. Isotopic ages obtained from this part of the belt range from 905 Ma to 495 Ma, none truly “Irumide~’,and have been taken t o indicate thermal overprinting of Kibaran events. Recent work by Ramsay and Ridgway (1976) suggests that the Irumide Belt may not be recognizable as a metamorphic entity and that it could be the result of a single tectonothermal episode that ended about 500 Ma ago.

TABLE 1O.VIII Structural history of the Chipata area, eastern Zambia (after Vavrda and Vrana, 1972) ~

Orogenic cycle

Deposition

T

? Mozambiquian

F4

F,

? Irumide

(Kibaran) ......................

.......

.......

._._._ .........

~~

Metamorphism

Deformation

-.,

-

N-S cataclasis and mylonitiza-

M 5 - greenschist facies

tion - NW-trending strain-slip cleavage (sericitization and chloritization)

M4 - greenschist facies

- open

M3 - greenschist-amphi facies

folding and penetrative deformation

......................

Mwami Group ............

Plutonism

......

__

...... _ _ _. .._ _ ........... - tight

.....................

............... ~

............

isoclinal folding and basin and dome interference structures

? Ubendian

biotite-granite

olite

.

_ _ .._............. ._.....................................

_ - ......................................

M2 - amphibolite-granulite

facies

--.......

.......

Musoka Granite

regional migmatization (F1 - N-S

folding

752

With A.C.J. Hickman (pers. commun., 1976) they place the western front of the Mozambique Belt at the eastern margin of the Bangweulu Block. Hickman also concludes that the Pan-African event in eastern Zambia is made up of a cycle of separate polyphase fold belts which include the Irumides and stretch back in time to the waning stages of the Kibaran orogeny.

Central and southern Malawi Although this area has been mapped in some detail and the various lithological units are well-established, its structural history is still uncertain.

TABLE 1O.IX Possible comparative structural history of northern and central Malawi (after Thatcher, 197313) Orogenic cycle

N. Malawi

C. Malawi

Mozambiquian

shearing, cataclasis, and amphibolite to greenschist facies metamorphism

?

Irumide (Kibaran)

brittle deformation and greenschist facies metamorphism

brittle deformation and greenschist facies metamorphism

deposition of Mafingi Group

deposition of Upper Mchinji Group

intrusion of Nyika Granite

intrusion of Dzalanyama Granite deposition of Lower Mchinji Group

i Ubendian

(

late syntectonic granite

late syntectonic perthitesyenite complexes

periods of tight folding in amphibolite and cordierite-granulite facies followed by more open folding

periods of tight folding and metamorphism in amphibolite and pyroxenegranulite facies

syntectonic granites

753 TABLE 10.X Outline of the structural history of southern Malawi (after Clifford, 1 9 7 4 , p. 1 3 ) Age (Ma)

Orogenic cycle

493 f 16

?

550

718

Nature of event

I

pegmatite formation and emplacement of post-tectonic granites and syenites

Mozambiquian

cataclasis and metamorphism

i

* 25

o p e n folding, amphibolite facies metamorphism and emplacement of syntectonic syenite complexes -.... -----.-._ __----. _ _ _ _ _ _ _ _ _ _ _ _ _- .- _ _ _ _ _- -.. ._____._. ? ? deposition of Upper Precambrian sediments, subsequently removed by erosion --.--___-__ - ---_ - - - - - - - - - - _ _ _ _ _- -- _- _ _ _ __ > 1 1 2 6 t 29

Irumide (Kibaran)

tight folding and granulite facies metamorphism

In central Malawi it seems possible to recognize pre-Mafingi, Mafingi, and post-Mafin@ elements (cf. Tables 1O.VI and 1O.VII) and Table 1O.IX gives a possible comparison of this area with the better-established sequence of events further north. I t should, however, be noted that, at present, there are few reliable age determinations. Central and southern Malawi appear to have essentially the same basic structural histories. In the south a period of tight folding accompanied by granulite facies metamorphism was followed by more open complex folding under amphibolite facies conditions. The granulites commonly occupy the cores of major structural culminations. Large bodies of syntectonic perthitic syenite lie within culminations and interference zones and there is a province of late- t o post-tectonic alkaline granites. K-Ar age determinations of the rocks of central and southern Malawi lie within the range 700 Ma t o 350 Ma but many of these are apparent ages and the preliminary results of more comprehensive Rb-Sr work are given in Table 10.X. The results show that the granulite facies metamorphism is of broadly Kibaran age and that more open folding, accompanied by amphibolite facies metamorphism, is early “Mozambiquian”. The post-tectonic granites were emplaced towards the end of the Mozambiquian cycle. Clifford (1974) considered that the interface between the older “dry” granulite facies rocks and the younger “wet” amphibolite facies types could represent the upper P-T limits of new amphibolite facies metamorphism. I t is possible that much of southern and central Malawi consists of remnants of a formerly much more extensive zone of granulites preserved within what are now mainly amphibolite facies rocks. However, it is by no means certain that all amphibolite

754 facies rocks result from downgrading. Nowhere is there firm evidence for the deposition of upper Precambrian sediments in the Kibaran-Mozambiquian interlude. THE PAN-AFRICAN EVENT

All recent work points t o the existence of a polyphase tectono-thermal event affecting the whole of Malawi and eastern Zambia during the period from about 900 Ma t o 350 Ma ago but nowhere is there.firm evidence of sedimentation during this time-span. The main effects were confined t o re-activation of pre-existing Precambrian rocks and associated plutonism. The metasediments involved in the Pan-African event were deposited preKibaran and, in part, possibly pre-Ubendian. The marked lithological differences between the northern and southern sub-provinces of Malawi and adjacent areas could result from original sedimentary differentiation. In the north the Pan-African event was reflected by several phases of brittle deformation and strong local cataclasis under amphibolite t o greenschist facies conditions but, in southern Malawi, more plastic deformation accompanied by amphibolite facies metamorphism took place. Alkaline metasomatism and plutonism was very widespread during the earlier stages (700 Ma t o 600 Ma), forming perthitic syenites and nepheline-syenites in both the northern and southern sub-provinces and, in the closing stages, this changed t o calc-alkaline activity. There is evidence that the effects of the Pan-African event persist as far west as the edge of the Bangweulu Block in Zambia and it may include separate polyphase fold belts that either merge into each other or interfere in only a limited way. The available Rb-Sr and K-Ar age determinations from this area included in the histogram of Fig. 10.17 confirm that the main Pan-African event took place between 700 Ma and 350 Ma ago. The 950 Ma t o 750 Ma ages could be interpreted either as overprinting of Kibaran ages or as a separate thermal episode.

H. The Damara Mobile Belt in South West Africa/Namibia

*

INTRODUCTION

The Damara Mobile Belt is the most southerly of the great Pan-African belts (Kennedy, 1964), which are defined by Late Precambrian sedimentary and volcanic sequences, and are characterized by regional metamorphism and related granitic intrusion of Late Precambrian t o Early Palaeozoic age (550 100 Ma). The Damara sequence was deposited on a foundation of granitoid basement complex gneisses, granites and infolded supracrustal formations,

*

* By R. Mason.

755 of pre-Irumide (1500 Ma-pre-2000 Ma), and Irumide (900-1400 Ma) ages. Recognition of the Damara Belt as an entity was made by Henno Martin (1965), whose description and interpretation of its geology remains a fine and perceptive synthesis of data available at that time. Since 1965 there have been several major research contributions on the Damara Belt, in addition t o the continued contribution made by the Geological Survey. The present account is an attempt t o place the results of the recent research into the perspective created by Martin, and it records personal observations and interpretations on the Damara Belt which have developed from extensive regional studies in Namibia. Seen in the context of present-day crustal configuration, the Damara Belt is disposed relative t o two major features - the faulted, aseismic Atlantic seaboard of S.W.A./Namibia, and the southwestern extension of the East African Rift System. The central Damara Belt has developed on the extensions of the Zambezi/Okavango Rift System, and its evolution has been controlled and dominated by northeasterly trending, taphrogenic faults. The northwestern arm of the Damara Belt (Fig. 10.18) in the Kaokoveld and southwestern Angola has developed under the influence of north/northwesterly taphrogenic faults, which eventually guided the partition of Africa and South America at the onset of Mesozoic sea-floor spreading, and development of the present Atlantic Ocean. We will see from this account that these two taphrogenic fault systems were developed at least 1400 Ma ago, and that they controlled the subsequent evolution of both the Irumide, and the Damara cycles I . They also influenced the development of Karoo sedimentation and volcanicity and post-Karoo alkaline intrusive complexes. Furthermore the inherent weaknesses along the rifted margins of the Atlantic seaboard guided the partition of Africa and South America, at the onset of Mesozoic ocean-floor spreading and the opening of the South Atlantic. Present-day landforms in S.W.A./Namibia reflect this influence and the broken edge of the African Plateau forms a spectacular escarpment, subparallel t o the coast between the Orange River and Walvis Bay and about 100 km inland. Significantly the edge of the plateau surface cuts back inland along the southern margins of the central Damara Belt, and becomes rather indeterminate northwards, until it is encountered again around the northern margins of the Damara Belt, where it re-establishes an escarpment parallel t o the coastline in the Kaokoveld. The coastal plain, which is occupied by the Namib desert, has been spectacularly warped and faulted during the subsidence of the Atlantic margins of Africa following sea-floor spreading in the Mesozoic. Presumably the edge of the continental slope off the Namibian coastline, once marked the crestal position of a great anticlinal warp, which

Cycle here is used in the sense of Sutton’s (1967) “chelogenic” cycles.

7 56 disintegrated by a process of rifting, spreading and subsidence during Mesozoic times. THE BASEMENT COMPLEX

The Basement Complex crops out in several major inliers along the northern and southern margins of the Damara Belt, and in smaller inliers in the central parts of the belt between Okahandja and Walvis Bay (Fig. 10.18). It is predominantly granitoid, with infolded remnants of metavolcanic and metasedimentary cover rocks. The lithologies and metamorphic/structural styles of these cover rocks are remarkably similar, and their common identity suggests their evolution in similar environments under similar conditions. Available geological and geochronological data indicate that these Basement formations and the granitoid rocks which surround them, were evolved prior to the commencement of the Irumide Cycle (1400 Ma) and probably prior to 2000 Ma. A major basement ridge separates the Irumide formations (Sinclair Group) of the Sinclair-Helmeringhausen area from the Atlantic coastline, and several small basement inliers occur within the area underlain by rocks of the Sinclair Group. This basement ridge, which is well exposed between Luderitz and Aus, exposes deeper crustal sections than most of the Damara inliers to the north, and granulite facies rocks are developed in part (Jackson, 1976). I t is not surprising to find such exposure of deeper crustal sections in a present-day continent margin situation which has been a site of repeated tectonic instability associated with rifting. Apart from being the foundation of the Damara Belt, the Basement Complex appears t o have had an inherent configuration which strongly influenced the development of both subsequent Irumide and Damara cycles. The major Basement inliers, which are briefly described here, have acted as positive features during the development of the Damara Belt, and have thus influenced sedimentation and the development of sedimentary systems in the Damara depositary. In addition, the behaviour of the Basement Complex in response to tectonism during the Damara Cycle, has resulted in the development of a variety of structural styles within the Damara Belt. In some cases the Basement domes and ridges appear to have remained relatively inert, and the Damara cover sequence drapes around these “highs” in large, open structures (e.g., the northern margins of the Damara Belt). Along the southern margins of the Damara Belt, brittle deformation has resulted in slices of basement rocks being thrust into the Irumide and Damara cover rocks, resulting in the formation of nappe-type structures. In the central parts of the belt, reactivation of the basement has led to partial melting and the generation of mantled gneiss domes in extreme cases.

7 57

The southern margin inliers The Rehoboth inlier is the largest of the Basement inliers along the southern margins of the Damara Belt, and it forms part of a major northeasterly trending, pre-Damara ridge which effectively bounds the Damara Belt. De Kock (1934) and Gevers (1934) recognized the existence of a complex granitoid basement with infolded, metamorphosed cover rocks, but it was De Waal (1966) and Halbich (1970) who first mapped out the cover formations and intrusive complexes of various ages in the western parts of the inlier. Radiometric age determinations have recently indicated that many of the granites in the Rehoboth inlier are in fact of Irumide age (1400-900 Ma), and that Irumide cover formations are probably more widespread than was previously recognized (Burger and Coertze, 1973 and 1975; Hugo and Schalk, 1975). The gneisses and associated metavolcanic and metasedimentary rocks of the Basement Complex were included in a Marienhof Formation by De Kock (1934). De Waal (1966) continued the usage of “Marienhof Formation” but Halbich (1970) proposed the term Abbabis Formation (previously Abbabis System, Gevers, 1934), for the supracmstal rocks, and recently Malling (1975) has referred t o these as the Elim Formation. The metamorphosed volcanic and sedimentary rocks are represented by chloritic, sericitic and amphibolitic schists, metaquartzites and related quartz-sericite schists, banded calc-silicate rocks, marbles and magnetite quartzites (partly recrystallized banded iron formations). This association, with a structural style characteristic of rocks which have undergone metamorphism under greenschist to lower amphibolite facies conditions, is very typical of the Basement Complex cover sequences throughout central and northern Namibia and southern Angola. These cover sequences appear to have developed in restricted depositaries which were floored by older granitoid gneisses. Several other inliers are exposed in the southern shelf area (between the Khomas Trough and the Rehoboth ridge) and around the closure of the Khomas Trough near Steinhausen (Fig. 10.18). These inliers are all underlain by granitoid gneisses, granites and remnants of infolded cover rocks similar to those exposed in the Rehoboth inlier. It is possible that some of the granites in these inliers may also be of Irumide age.

Znliers in the central Damara Belt The Abbabis inlier is the largest of several small inliers of Basement which are found in the cores of major domal structures draped by units of the Damara Sequence in the central parts of the Damara Belt. Gevers (1931)was the first t o recognize the unconformable relationships between the Damara sequence and the Basement Complex in the Abbabis inlier, although it was

758 not until Smith (1961) mapped the area systematically, that the broad relationships of the inlier were established. Smith also first suggested that Basement cores would be found in many of the domal structures in the central Damara Belt, and this has only recently been confirmed by Jacob (1974 and 1978), Bunting (1977) and Sawyer (1978), who have shown that much of what was previously regarded as reconstituted basal Damara (Nosib Group) - “red granite-gneisses”, is actually Abbabis granite gneiss and reactivated equivalents, which form the cores of typical mantled gneiss dome (Eskola, 1949). U-Pb dating of zircons from two samples of Abbabis granitegneiss has yielded a concordia intercept age of 1925 ? 300 Ma (Jacob e t al., 1978), which fits into the emerging picture of an 18Oci-2200 Ma age grouping for the Basement Complex forming the infrastructure elsewhere in the Damara Belt (Burger et al., 1976). The only cover rocks of significance occur as narrow infolded remnants in the Abbabis inlier itself (Smith, 1961), and consist of a variety of acid, intermediate and basic schists, minor calc-silicate rocks with impure marbles, minor meta-conglomerates and meta-arkose, and gneisses displaying amphibolite facies mineral assemblages. These rocks appear t o be mainly metavolcanic with minor sedimentary intercalations, and include thin discontinuous magnetite quartzite (meta-chert) horizons which are closely associated with copper-zinc mineralization a t the Henderson and Ubib prospects.

The northern inliers The northern margins of the Damara Belt are marked by an arcuate chain of major basement ridges and domes, which extends over more than 1000 km from east of Grootfontein, through the Kamanjab inlier into the Kaokoveld inlier, and thence into southern Angola (see Fig. 10.18). These inliers were positive features during the deposition of the Damara sequence and they exerted a strong influence on the disposition of both vertical and lateral facies changes in the Damara rocks. The Grootfontein inlier is only poorly exposed, close t o the edge of the Kalahari sand plains, and outcrops are highly weathered. However, granitoid gneisses can be identified, together with a variety of supracrustal metavolcanic and metasedimentary schists which appear t o occur in narrow, infolded strips within the gneisses. Where exposures are adequate the foliation of the gneisses and the supracrustal rocks appears t o be aligned along ENE axes. The Kamanjab inlier and related Welwitschia inlier is the largest of the basement inliers in the Damara Belt, and its northwesterly trending extensions through the Kaokoveld into southern Angola (Fig. 10.18). The southwestern lobe and southern parts of the inlier show an increase in metamorphic grade relative t o the northern parts of the inlier and there is other evidence that block faulting and shearing has exhumed lower parts of the crust in this area and influenced major outbursts of Nosib igneous activity at the

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763 commencement of the Damara Cycle (Frets, 1969; Miller, 1974; and Martin and Porada, 1977a,b). The eastern half of the Kamanjab inlier consists of a complex development of granitoid gneisses, granites and granodiorites of various types and supracrustal gneisses, with minor amphibolites, metaquartzites and calcsilicate rocks. Studies of a small area of granite near Fransfontein (Clifford et al., 1962 and 1969) indicate that the Fransfontein granite was emplaced about 1650 Ma ago, and Frets (1969), Guj (1970b) and Porada (1974) have all recorded situations where Fransfontein-type granites intrude supracrustal rocks of the Huab and Khoabendus Formations. The western half of the Kamanjab inlier is characterised by the development of thick sequences of metavolcanic and metasedimentary supracrustal rocks of the Huab Formation in the south (Frets, 1969) and the Khoabendus Formation in the north (Porada, 1974). Guj (1970a) draws attention t o evidence which suggests a correlation between these formations, and apart from physical continuity, both formations are intruded by Fransfontein-type granites. Burger and Coertze (1975) record U/Pb zircon ages between 1760 and 1860 Ma for granitoids and acid lavas of the Huab complex west of Fransfontein. The Khoabendus Formation is characterized by regional metamorphism and a tectonic style of greenschist t o amphibolite facies type, and shows the higher grades of metamorphism where it merges southwards into the Huab Formation. The Huab rocks are more complexly folded and attain upper-amphibolite grades of metamorphism. Both Khoabendus and Huab Formations are characterized by thick developments of metamorphosed volcaniclastic sedimentary rocks, ash flows, volcanic rocks of basaltic t o rhyolitic composition with a preponderance of intermediate and acidic pyroclastic rock types, and metaquartzites, marbles, metadolomites and associated calc-silicate rocks, phyllitic rocks (siltstones, shales?) and meta-cherts (some of which are ferruginous). Several generations of basic and acid porphyry intrusions can be distinguished. Until further work is undertaken to date these formations, there is no certainty about their relationship to other pre-Damara formations in Namibia. Some of the supracrustal rocks in the area may even belong t o the Irumide Cycle (9001400 Ma) as shearing has affected the basement during the Irumide and Damara cycles. Very similar supracrustal formations are developed in southern Angola. The Kaohoueld inliers occur in the extreme northwestern parts of Namibia and extend into southern Angola. The basement complex has been severely sheared in ENE-trending and NNW-trending zones, chief of which are the Marienfluss zone (NNW) and the Ruacana zone (ENE) (Fig. 10.18). The Marienfluss zone includes a series of shear belts in the basement on the eastern borders of the Marienfluss (itself a major faulted valley), and a major ductile shear belt to the west of the Marienfluss which extends southwards towards the Etendeka lava plateau (Karoo age), and northwards towards

764 Mocamedes in southern Angola. These shear zones are characterized by steeply dipping shear planes, the development of cataclasites with thin mylonite layers and extensive refoliation parallel t o the shear zones. The Ruacana shear-zone extends from east of Ruacana, along the faulted southern margins of the Kunene Complex (Vermaak, this volume, Chapter 9), and in places extends up t o 10 km in width. It marks the axial zone of a major ENE-trending warp in the basement, which was a positive feature through the Damara Cycle, and which appears linked t o the northeasterly trending Walvis Ridge. Small infolded, and fragmented enclaves of the supracrustal Epupa Formation occur in the area south of the Kunene Basic Complex and around Epupa Falls. The Epupa Formation appears to be similar t o the Haub-Khoabendus Formations of the Kamanjab inlier, and consists of a variety of basic and acid schists, leptites and metasedimentary rocks (meta-arkose mainly) with minor chert-jaspilite horizons. Again there is evidence of several generations of basic dyke intrusion in the inlier. Martin (1965) not only suggested a correlation of the Epupa supracrustals with other basement formations in Namibia, but also suggested that extensive developments of similar rocks in southern Angola should be included in the Epupa Formation. Certainly the physical continuity of the Epupa Formation across the Kunene River is an established fact, and there is little reason to separate the Kaokoveld basement inliers from the extensions into Angola apart from an international boundary. In southwestern Angola, continuations of the basement complex exposed in the Kaokoveld inliers, extend around the Kunene Complex and northwestwards towards Mocamedes. Sheared granitoid gneisses with infolded remnants of the Epupa Formation extend up to 25 km across the Kunene River into Angola and the grade of metamorphism in this zone adjacent to the river is upper amphibolite facies. Just as the southwestern parts of the Kamanjab inlier expose lower parts of the crust, this strip along the southern borders of Angola appears t o represent lower level, more mobilised sections of the basement complex, probably exhumed by faulting and crustal warping. Major NNW trending faults and shear zones cut through the southern tip of the Kunene Complex and extend through Oncocua northwards. The basement granites have been reduced t o cataclasites along this zone. A major belt of metavolcanic and metasedimentary schists, sheared lavas, quartzites and cherts, crops out in mountainous terrain WSW of Oncocua, in a zone some 120 km along a WNW strike and 35 km width. These have been named the Ovipaka schists (Cochran et al., 1975) and they are partly flanked by higher grade migmatitic gneisses and partly intruded by younger granites. The schist belt is truncated in the west by a major north-northwesterly shear belt which appears t o be an extension of the shearing associated with the Marienfluss valley in the Kaokoveld of Namibia. The Ovipaka schists are predominantly of intermediate composition but contain a series of prominent

765 metabasalt zones which invariably include pyritic chert beds. Numerous basic intrusives occur with the Ovipaka sequence, which is characterized by a greenschist facies metamorphic and structural style. The whole belt of schists appears t o form a major synformal structure which refolds early isoclinal folds. Crenulation cleavages cut an earlier schistosity at high angles. In the Virei area to the north, a sequence of more acid schists, metarhyolites, ferruginous cherts, graphitic and aluminous schists and sheared porphyry occur in another major belt infolded in complex gneisses and migmatites, and extensively intruded by younger granites. Available radiometric age dating indicates that most of the younger granites in southern Angola are of Irumide age (900-1400 Ma) but no dating has been undertaken on any basement rocks as yet. The structural evidence and evidence of the generation of several phases of basic dyke intrusion in the basement, some of which are metamorphosed and deformed, suggests a long and complex history prior to the Damara Cycle. Again some of these dykes may be of Irumide age.

Discussion The Basement Complex of S.W.A./Namibia and southwestern Angola bears little resemblance t o the granite-greenstone complexes of the Rhodesian and Kaapvaal cratons (Anhaeusser et al., 1969), apart from the superficial similarity of infolded fragments of predominantly volcanic cover sequences set in granitoid gneiss complexes. There are no major developments of typical basaltic “greenstone” sequences with their common duality of tholeiitic and high-magnesian basalt types and associated ultrabasic rocks. In fact the volcanic rocks and associated sedimentary units of this region are typical of calc-alkaline provinces and are very similar t o the Irumide sequences which developed subsequently in rifted depositaries in S.W .A./ Namibia. The granitic rocks appear t o be predominantly granodioritic to granitic, rather than tonalitic to granodioritic, and their relationship t o the cover rocks shows little of the arcuate patterns so typical of the granitegreenstone terranes. The inadequate geochronological data permit little in the way of meaningful comparisons, but it would appear that the Basement Complex in S.W.A./Namibia evolved at a later stage than the Rhodesian and Kaapvaal cratons, and may only have been initiated about the time of the major granite-forming events at between 2500 and 2800 Ma. THE IRUMIDE CYCLE

The period between 1400 Ma and 900 Ma (Irumide and Kibaran of Cahen and Snelling, 1966), appears t o have been characterized by localized mobility in restricted, fault-bounded zones, which were the sites of sporadic volcanism, sedimentation and granite emplacement in S.W .A./Namibia, northwestern Botswana and southern Angola. There is evidence of widespread

766 isotopic readjustments of Irumide ages in the pre-Irumide basement complex of the region, probably caused by tectonism involving uplift along major crustal warps accompanied by extensive faulting and shearing. The main Irumide zones of volcanism, sedimentation and granite intrusion, coincide with the northeasterly trending basement ridge which defines the southern margins of the Damara Belt (De Villiers and Simpson, 1974), and the southwestern extensions of the Zambesi-Okavango rift system, and with rifting parallel t o the north-northwesterly trending Atlantic seaboard in the S i n c l a i r Helmeringhausen area t o the south of the main Damara Belt (Watters, 1974) (Fig. 10.18). The so-called “Rehoboth magmatic arc” of Watters (1977) encompasses both these zones and Watters has suggested that the “arc” represents a zone of subduction pre-dating the Damara Cycle. The commencement of activity in the Irumide cycle was characterized by sporadic granite emplacement, but volcanic activity and related sedimentation appears to have been initially restricted t o the SinclairHelmeringhausen areas of Namibia (1300-1100 Ma). A series of more widespread granite emplacements occurred during the same period and this partly preceded, and was partly coeval with the onset of major volcanic activity along the southern margins of the Damara Belt in Namibia and northwestern Botswana (Burger and Coertze, 1973, 1975; Hugo and Schalk, 1975; Watters, 1977). Apart from a series of Irumide ages obtained from various granites in southern Angola, little is known of the existence of supracrustal rocks belonging to the Irumide Cycle, although there are indications that these may be developed there.

The SinclairHelmeringhausen area The Sinclair Group was deposited in a fault-bounded depositary between 1400 and 1000 Ma ago and the same north-northwesterly striking zone was the site of major granite emplacements during this period (Watters, 1974, 1977). This fault zone is aligned along the south-southeastward extensions of the faulted coastline between the Walvis Ridge in the north and Swakopmund in the south. Von Brunn (1967), and Watters (1974) have investigated this area in two major regional studies, and Watters (1977) has attempted to place the area into the wider perspective of the major Irumide zones in S.W .A./Namibia which he has collectively called the “Rehoboth Magmatic Arc”. Five volcanic and sedimentary formations have been recognized in the area, and radiometric dating has indicated that these were evolved about 1265 ? 1 0 Ma (Rb/Sr whole-rock isochron) (Kroner, 1975). A basal Nagatis Formation consisting of ashflow tuffs, felsic lavas, basic lavas, agglomerates and minor interlayered arkose, shale and grit bands, is followed by the Kunjas Formation, consisting of a basal conglomerate, arkose and shale. This is followed by a major volcanic sequence, the Barby Formation which con-

767 sists of porphyritic trachy-andesite and trachy-basalt, basalt, rhyolite, ashflow tuffs, tuffs, volcanic conglomerates and quartzites, which are intruded by gabbroic, dioritic and monzonitic intrusives. The Barby Formation has been dated at 1265 f 10 Ma (Kroner, 1975), and is followed by the Guperas Formation consisting of lithic sandstones, conglomerates, rhyolitic lavas, ashflow tuffs, rhylotie plugs, basic lavas and a series of quartz-porphyry dykes and basic sills and dykes. The Guperas Formation is followed by a thick but localized development of red feldspathic sandstones, conglomerates and shales of the Auborus Formation. The volcanic formations appear to be of calc-alkaline affinity (Watters, 1974), a factor which persuaded Watters (1977 and 1978) to suggest that they had evolved over a subduction zone related t o an island arc setting. The Sinclair Group is only locally affected by shearing and igneous intrusions and can be regarded as non-metamorphic in a regional sense. Several major granite emplacement events occurred before and after the deposition of the Sinclair Group, a similar evolutionary pattern t o that proposed in the Nauchas/Rehoboth area.

The NauchadRehoboth area The regional distribution and stratigraphic disposition of the Irumide formations along the southern margins of the Damara Belt in Namibia have been established by Martin, 1965; De Waal, 1966; and Schalk, 1973. Martin (1965) placed these formations in a Dordabis System, but subsequently Schalk (1973) has recognized three formations which are very similar to those of the Sinclair Group (see previously). These include the basal Grauwater Formation, consisting of volcanic rocks with minor interlayered sedimentary horizons, followed by the Doornpoort Formation which is a mixed volcanic and sedimentary sequence and is overlain by a sedimentary sequence, the Klein Aub Formation, These formations are collectively placed in a Rehoboth Group. The Grauwater Formation is preserved in major anticlinal cores, and consists of basic intermediate and acid lavas and pyroclastic rocks, with minor interbedded quartzite, arkose and conglomerate. The Doornpoort Formation consists of a predominantly volcanic basal portion with basic to acid lavas and pyroclastic rocks, followed by conglomerates and distinctive reddish coloured quartzites. The Klein Aub Formation consists of a thick but localized pile of poorly sorted conglomerates followed upwards by volcaniclastic sandstones, quartzites, calcareous quartzites, mudstones, shale and thin impersistent limestones. The Klein Aub Formation is remarkably similar t o the Ghanzi Formation in northwestern Botswana. Ripple marks, mudcasts, small scale cross-bedded units and very finely laminated, varve-like bedding indicate a shallow-water lacustrine-type of environment. It would appear that infilling of the depository was initiated by volcanic activity which was followed by rapid deposition of coarse clastic

768 material from fault scarps bounding it, and finally was infilled by progressively finer material in a lacustrine environment. It is significant that within these lacustrine beds, small, thin but extensive, cupriferous beds are found and these are being mined at Klein Aub. Extensive sedimentary copper mineralization in similar beds has been discovered recently in northwestern Botswana. Although there are striking similarities in the nature and disposition of formations to the Rehoboth Group and those of the Sinclair Group, the rocks in the Nauchas/Rehoboth area have been affected by regional dislocation metamorphism developed under greenschist facies conditions. Furthermore the Grauwater Formation appears to have been evolved after the deposition of the volcanic formations of the Sinclair Group (Burger and Coertze, 1973 and 1975). However, the Nauchas/Rehoboth area is characterized by extensive emplacement of granitic rocks in several major episodes during the development of the Irumide Cycle, with very much the same evolutionary pattern as in the SinclairHelmeringhausen area. In addition, a major layered basic complex (the Alberta Complex - De Wad, 1966) was intruded into the Nauchas area prior to the deposition of the Grauwater Formation. The tectonic style of the formations is characterized by major northeast to southwest-oriented upright folds with axial plane cleavage better developed in the argillaceous rocks and in the finer-grained rocks, than in the more massive volcanic rocks and quartzites. Local variations in the intensity of deformation are typical and appear to be related to proximity or otherwise of major shear zones especially in the areas along the northern edge of the pre-Damara ridge. The Nauchas/Rehoboth area was extensively invaded by granitoid intrusions, prior to, during, and after deposition of the Rehoboth Group.

North western Botswana In northwestern Botswana the Irumide rocks are preserved within a zone of major rifting - the Okavango rift system - and they comprise a volcaniclastic Kwebe Porphyry Formation, overlain by a sedimentary-volcaniclastic Ghanzi Formation (Thomas, 1969a,b). Available radiometric dating on rocks of the Kwebe Formation indicates ages of between 900 Ma and 1050 Ma. The Kwebe Porphyry Formation consists predominantly o i riiassive quartz-feldspar porphyries, with subordinate flow-banded porphyry and porphyry breccias, massive and vesicular diabases, feldspathic sandstones, grits and conglomerates. Some of the sandstones appear t o be pyroclastic in origin, and most of the sedimentary rocks in the formation appear to have been derived locally from the volcanic rocks. The Kwebe Formation appears to occupy the cores of northeasterly trending anticlinal structures (Thomas, 1969a). The Ghanzi Formation consists of strongly cleaved feldspathic sandstone, feldspathic quartzites, calcareous sandstones, mudstones, shales and thin limestones. Thomas ( 1969a) records heavy-mineral laminae, cross-

769 bedding, ripple markings and mudcrack casts in these rocks. Where cleavage is strongly developed, the argillaceous rocks tend to become phyllitic. The two formations are tightly folded in major upright folds and axial planar cleavages are well developed. In some of the tighter folds the cleavage is developed to the extent that fold closures are obscured. The outward appearance of these rocks signifies little metamorphic change from their original state and in fact even biotite is rarely developed or incipient in the pelitic rocks. In the vicinity of the Namibia/Botswana border, the Kwebe Formation appears to be either poorly developed or absent, and the Ghanzi Formation is characterized by an increase in the quantity of quartzitic rocks developed in the sequence. Schalk (1973) proposed the name Doornpoort Formation for the continuations of the Ghanzi Formation into Namibia (in the Gobabis and Rehoboth districts). The remarkable similarities of sedimentary lithologies and structures of the Klein Aub Formation and the Ghanzi Formation have already been mentioned, and the development of sedimentary copper deposits of similar type in both is further evidence of their development in similar environments. Toens (197 5) has reported copper mineralization in similar beds in the Witvlei area. In all these cases the mineralization consists of finely disseminated chalcocite with significant amounts of silver.

Discussion Although the Irumide formations and granites can be related in two major zones parallel t o the Atlantic margins of S.W.A./Namibia, and to the northeasterly trending extensions of the rift system, there is little evidence to support Watters’ (1977) suggestion that the Irumide formations developed in major arc-like subduction zones, apart from the fact that the volcanic formations are of calc-alkaline types. Indeed, most of the evidence points towards their evolution in a rifted intracontinental environment, and the problems of correlating the Irumide formations are a manifestation of the lack of continuity between the relatively small, localized yoked basins, in rift valleytype situations. The Irumide Cycle thus appears t o have marked a long period of sporadic tectonic unrest with rifting, associated volcanic activity and the development of restricted intermontane yoked basins, which accommodated thick, but localized piles of clastic sediments, volcanic rocks and lacustrine deposits. The cycle is remarkably similar in these respects t o the recent development of volcanic and sedimentary sequences in the East African rift system. However, there is no such recent parallel there for the major granite emplacement events which appear to be a persistent feature of the entire Irumide Cycle in the region.

770 THE DAMARA CYCLE

There appears t o have been a pause in the development of major rockforming events after the emplacement of late-Irumide granites along the southern margins of the Damara Belt (950 t 50 Ma). About 800 Ma ago, the basal formations of the Damara Supergroup began t o evolve in a series of disconnected, faulted depositories across the Damara Belt. The Damara Supergroup is divided into four groups. A basal Nosib Group is followed by the Otavi Group (equivalent t o the Swakop Group of the central Damara) and in the north the Mulden Group overlies the Otavi Group (Fig. 10.18). Martin (1965) drew attention t o the possibilities of correlating the Nama sequence, which forms a major undisturbed platform sequence to the south of the main Damara Belt (Germs, 1972), and concluded that the Nama sequence might be equivalent t o the upper parts of the Swakop/Otavi Groups and the Mulden Group. Kroner (1971) took this further by suggesting a correlation between the Nama sequence and the two upper groups of the Damara Supergroup as a whole. Germs (1974) argued against both correlations and suggested that the Nama Group could be correlated with only the Mulden Group of the Damara Supergroup. The Nama Group is certainly of Late Precambrian age (Germs, 1974), and whatever its precise correlation with the Damara Supergroup may be, it must have developed within the framework of the Damara Cycle as a whole. Martin (1965) separated the Damara Belt into two main facies - an Outjo “miogeosynclinal”, and a Swakop “eugeosynclinal~’facies. The miogeosynclinal facies was restricted t o the northern of margins of the Damara Belt disposed about the basement inliers of northern Namibia, and the eugeosynclinal facies extended over the rest of the belt. Facies transitions were recognized between the two major facies, but from Martin’s descriptions it is clear that the inclusion of the majority of the Damara Belt in the Swakop facies is too simplistic. In fact it can be demonstrated that the only eugeosynclinal developments (in the classic sense) in the Damara Belt, are confined t o the Khomas Trough, and the coastal strip in northwestern Namibia extending into southwestern Angola - the Kaokoveld Trough (Guj, 1970a). These troughs are typical of active flysch depositories, with great thicknesses of greywacke and shale, and interlayered volcanic schists, basic lavas and associated graphitic schists and cherts. The shelf sequences show rapid transitions into these troughs and direct correlation of units across the Khomas Trough is not possible. The remainder of the Damara Belt was covered by thinner developments of the shelf deposits typical of the Outjo facies, and the Khomas Trough separates a northern shelf facies developed over the central Damara Belt from a very similar southern shelf facies, developed over the basement ridge which marks the southern limits of the main Damara Belt (Fig. 10.18). The present nomenclature in Namibia retains a “Swakop Group” for the Damara sequence above the Nosib Group in the central and southern parts of

771 the Damara Belt, but the term “Otavi Group” is used in place of Martin’s “Outjo Facies”. However, this usage still maintains a separation of the Otavi shelf (Outjo Facies) from the rest of the Damara Belt, and the “Swakop Group” is still used as something of a “sack” term which does not take into account the variations in tectonic setting, stratigraphy and sedimentation across the remainder of the Damara Belt. The Swakop Group includes Martin’s “Hakos” and “Khomas Series”. Reactivation of the pre-Damara basement during the later stages of sedimentation in the Damara Belt, led to regional metamorphism and deformation under a variety of conditions ranging from dislocation greenschist to more mobile amphibolite and even granulite facies. The style of metamorphism and deformation was related t o the degree and type of mobility in the Basement Complex and the culmination of the Damara Cycle was marked by widespread post-tectonic recrystallization and the intrusion of major posttectonic granite batholiths, granite sheets and pegmatites in the central parts of the Damara Belt.

The Nosib Group The volcanic, volcaniclastic and sedimentary rocks of the Nosib Group are remarkably similar in their lithologies, occurrences and tectonic setting, to the formations developed during the Irumide Cycle. Indeed, deposition of the Nosib Group rocks could almost be regarded as a long drawn-out continuation of the pattern of sedimentation and volcanicity established during the Irumide Cycle. However, the Damara Cycle proceeded beyond the stages of initial volcanism and terrestrial sedimentation associated with restricted fault-bounded basins, and subsidence of the rifted zones was followed by localized marine transgressions with successive overlapping of the lower Otavi and Swakop sequences onto pre-Damara formations. Martin and Porada (1977a,b) have related the commencement of the Damara Cycle t o restricted deposition of Nosib Group sedimentary and volcanic rocks within three major northeasterly trending graben zones. These zones correspond to the transition zone to the edge of the Otavi shelf, a central zone lying t o the north of the Khomas Trough and the southern margins of the Damara Belt, adjacent t o the Basement-Irumide ridge. Another major fault-bounded Nosib depository occurs a t the edges of the Kaokoveld Trough with a northnorthwesterly disposition. There is a growing body of evidence that not only was the Nosib deposited in fault-bounded yoked basins, floored by Basement Complex, but that it was largely restricted t o these zones, which were then filled up by clastic and carbonate units of the lower parts of the Otavi/ Swakop sequence. The Nosib Group is characterized by the products of high-energy mass wasting. Thick (up t o 6000 m), localized clastic wedge sequences of conglomerates, arkoses, feldspathic sandstones, feldspathic quartzites, and inter-

772 layered volcanic material of predominantly rhyolitic t o dacitic composition are developed in the Nosib depositories. In some of the more tectonically active areas, basic volcanic rocks are developed, particularly in the Kaokoveld area and the southern flanks of the Grootfontein inlier at the edges of the Otavi shelf. The Nosib Group is not extensively developed across the Damara Belt, but is confined t o relatively narrow, fault-bounded zones within the belt. In the northern Kaokoveld and the southern parts of the Damara Belt, the Nosib includes substantial thicknesses of finer-grained clastic material (shale, mudstone sequences), and there is evidence of glaciogenic beds in the Nosib sequences associated with the Otavi shelf and the areas to the south of the Khomas Trough. Volcanic beds in the Nosib Group are best developed in the northern parts of the Damara Belt (Martin, 1965; Guj, 1970a; and Miller, 1974), and tend to occur in the upper parts of the Nosib sequence. The main centres of volcanic activity appear to have been the southern parts of the Otavi Mountainland, the southwestern edge of the Kamanjab inlier, and the eastern margins of the Kaokoveld Trough. Basic lavas and associated pyroclastic rocks and cherts form a major unit (the Askevold Formation) above the basal Nosib clastic rocks in the southern parts of the Otavi Mountainland. Similar rocks occur along the margins of the Kaokoveld Trough where felsic volcanic units (the Noeroep Volcanic Formation) are also extensively developed (Guj, 1970a). Miller (1972 and 1974) and Frets (1969) have described thick (up to 6000 m) sequences (the Naauwpoort Formation) of ash flows, ignimbrites and calc-alkaline lavas with associated pyroclastic units of rhyolitic and dacitic composition, from areas close to the southwestern margins of the Kamanjab inlier. These are intruded by porphyries, basic bodies and small alkaline plugs. Three major centres of volcanic activity (the Summas Mountains, the Mitten Dome and the Ais Dome) described by Miller (1972), form part of a ridge which remained a positive feature during the deposition of the Swakop Group, and which coincides with major facies changes in the Swakop Group. Extensions of this ridge t o the northeast coincide with a northeasterly trending zone of nappes and recumbent folds overturned to the southeast (Clifford, 1962). There is evidence for volcanic activity in the Nosib sequence flanking the Khomas Trough t o the north and south, but there do not appear to be major centres of volcanic activity such as those described to the north by Miller (1972). In areas where the Nosib sequence is transitional upwards into the lower parts of the Otavi or Swakop sequences, quartzites, calc-arenites, calcareous shales and thin, impersistent carbonate horizons are typical of the uppermost Nosib units. This is best illustrated in the central Damara Belt between Swakopmund and Karibib (Smith, 1961; Jacob, 1974), and Nash (1971), drew attention to the presence of anhydrite in some upper Nosib rocks of this area, indicating yet again the restricted nature of the depository and the development of hypersaline conditions at this time (Martin, 1978). Else-

773 where, the Nosib clastic sequence is terminated abruptly and there is a pronounced unconformity separating it from the Otavi/Swakop sequences lying above it. Schalk (1973) and Halbich (1970) have demonstrated a lateral facies change in the Nosib of the southern shelf area, where thick (up to 6000 m) developments of terrestrial quartzites (Kamtsas Formation), grade southwestwards into a mixed marine facies (Duruchaus Formation), of calcareous shales (phyllites), thin impure limestones, quartzites and conglomerates. Schalk (1973) describes the Nosib depository here as being very irregular with prominent topographic depressions and “highs” and with sedimentation effected by high-energy fluviatile systems. The upper Nosib sequence developed in both the Kamtsas and the Duruchaus Formations, is characterized by upward transition t o interlayered calcareous phyllites, thin, impure and impersistent limestone beds, calc-arenites, and amphibolite bodies which appear to be of volcanic and sub-volcanic origin.

The Otavi-Swakop Groups The larger of the restricted Nosib depositories were filled up with sequences of the lower Otavi and lower Swakop Groups and these sequences are thus themselves limited to relatively small basins. There is evidence from some parts of the Damara Belt that the absence of lower Otavi-Swakop sequences in areas away from the major Nosib depositories is more a result of non-deposition rather than erosion. In the Otavi shelf the Abenab Subgroup represents the lower Otavi sequence, and its equivalents in the Swakop Group are the Ugab Subgroup (central Swakop shelf) and the Kudis Subgroup (southern shelf), both of which were included in the Lower Hakos Series previously (Martin, 1965). In general the Abenab Subgroup tends to be dominated by thick carbonate sequences with localized calc-arenite and calcareous shale developments, especially in the northern and western Kaokoveld. Stromatolitic beds and reef complexes are commonly developed in the carbonate units. The volcanic activity which characterized deposition of the Nosib Group in the western Kaokoveld, and the areas south of the Kamanjab inlier, continued during deposition of the lower Otavi sequences in these areas, where acid t o intermediate lavas, pyroclastic beds and ash flows occur interbedded with carbonate and clastic units. In the southwestern Kaokoveld large gabbroic sills and plugs (the Khankhaib Complex) appear to be penecontemporaneous with the lower Otavi sequence and form an integral part of the volcanic activity there (Guj, 1970a). Apart from the Abenab sequence in the Otavi Mountainland, which varies up t o 1500 m in thickness, there is a very thick (up to 2000m) development of Abenab rocks in the central Kaokoveld (south of Ohopoho). Elsewhere the Abenab shows sudden variations in thickness from 0-900 m, and is only sporadically developed.

774 The central Swakop shelf appears t o have been a positive feature over large areas during deposition of the Ugab Subgroup, and where deposition did take place, clastic material was an important a contributor to the sequence as was carbonate deposition. The major zones of deposition coincide with the main Nosib depositories. Miller (1972) described a sequence of dolomite with calcareous schists and localized conglomerate, quartzite, arkose and thin ironstone horizons in the areas adjacent t o the Nosib ridge (see previously) south of the Kamanjab inlier. The Ugab Subgroup is represented by the Rossing Formation in the areas immediately north of the Khomas Trough, and Smith (1965) and Jacob (1974) have described the formation in this area (Smith uses the term Hakos Series as does Martin, 1965). Here the upper Nosib Group sequence grades transitionally upwards into the lower parts of the Swakop Group, which are characterized by rapid lateral and vertical variations in clastic/carbonate proportions and the development of impersistent carbonate, and calc-arenite units. The carbonate units are metamorphosed t o marbles and the calc-arenite rocks form calc-silicate gneisses and granulites (Nash, 1971; Jacob, 1974). In the Otjosondu area, which lies close t o the northeastern edge of the Khomas Trough, Roper (1959) described a thin sequence of quartzites and marbles, which can be correlated with the Kudis (formerly Hakos) Subgroup of the southern shelf area. The Kudis Subgroup is predominantly composed of clastic rocks, whereas the Abenab Subgroup of the Otavi shelf is predominantly composed of carbonate rocks. All along the southern margins of the Khomas Trough, the Kudis sequence is accommodated in small basins punctuated by basement “highs”, and it is closely associated with the Nosib depository of the southern shelf region. The sequence consists of metamorphosed feldspathic sandstones, quartzites and local conglomeratic beds, interspersed with calc-arenite units, impure carbonate beds, dolomitic marbles, graphitic schists, impure iron formations and amphibolites of both sedimentary and igneous derivation. Rapid lateral variations in individual lithological units are typical, and sedimentation appears to have developed in restricted lagoonal environments. Stromatolitic beds and reefs are commonly developed close to the basement “highs”. On both sides of the Khomas Trough the Kudis and Ugab sequences grade laterally into increasingly impure carbonate and clastic units until they are indistinguishable from the total greywacke pile of the Trough. The same relationship can be seen at the margins of the Kaokoveld Trough (Guj, 1970a). Infilling of the Nosib depositories with the lower Otavi-Swakop sequences was terminated by a period of uplift and erosion which degraded most of the highland areas and culminated in an unprecedented, and very widespread marine transgression across the whole Damara Belt. Apparently for the first time in the history of the Precambrian of Namibia, the Damara Belt was covered by a shallow, epicontinental sea, and this coincided with the onset of cold climatic conditions. The record of this period is preserved in glacio-

775 marine shales, tilloids, iron formations, graphitic and pebbly schists of the Chuos Formation, which provides one of the few reliable and well distributed time markers within the Damara Belt. What followed the Chuos Formation in the Damara succession was very much more predictable and stable than the pre-Chuos sequences. From the northern margins of the Khomas Trough through t o the northern liqits of the Damara Belt, a shallow epicontinental sea formed the site for the deposition of generally thin, but very extensive, shelf carbonate units resting on the Chuos Formation. However, the northern margins were characterized by thicker carbonate deposition, and these areas define the Otavi Shelf Province. Furthermore, the Otavi shelf was the major site for the deposition of the Mulden Group which consists of a sequence of feldspathic quartzites and sandstones, shales and thin carbonate beds overlying the Otavi Group unconformably. The Tsumeb Subgroup overlies the Chuos Formation in the Otavi shelf region in the north, and its equivalent on the Swakop shelf is the Khomas Subgroup. The Otavi shelf was the site of major carbonate deposition in upper Otavi times and a sequence of up to 3000 m of massive to well-bedded dolomite overlies the Chuos Formation over much of the Otavi shelf. Only in the extreme northwestern parts of the Kaokoveld and into southwestern Angola is this monotonous dolomite sequence relieved by a major facies change into quartzites, sandstones and calc-arenite units with relatively thin carbonate developments. There is marked thinning of the Tsumeb Subgroup in the northern Kaokoveld. The dolomites of the Tsumeb Subgroup contain oolitic, stromatolitic and cherty units and represent a period of very stable shelf carbonate deposition. The southern and western margins of the Otavi shelf are characterized by thinning of the carbonate units and the development of localized mudflow breccias and shaley units interlayered with dolomite units. The Swakop shelf area is characterized by carbonate units (marbles and associated calc-silicate rocks) and interlayered biotite-quartz schists (sub-greywacke and shale) of the Karirib Formation. This formation varies in thickness between 300 and 500 m, and is remarkably persistent and constant over the entire Swakop shelf area, indicating that shelf conditions persisted in this region throughout the deposition of the upper Swakop Group. In the Otjosondu area (Roper, 1959) and in the southern shelf area (Martin, 1965; Habich, 1970) the equivalent Auas Formation (lower Khomas Subgroup) consists of a sequence dominated by quartzites and interlayered marbles, in contrast to the Swakop shelf carbonate sequences of the Karibib Formation. Both the Karibib and Auas Formations are followed upwards by sequences of biotite-quartz schists with minor quartzite and marble beds. The Khomas Subgroup sequences again grade transitionally into the Khomas Trough and merge into massive thicknesses of biotite and quartz-biotite schists, representing metamorphosed greywacke, sub-greywacke and shale.

776

The Khomas Trough The Khomas Trough represents a major and very thick development of flysch-type sedimentary rocks with interlayered volcanic rocks. These appear to have been deposited in a tectonically active trough which developed between the two major rifted Nosib depositories presently flanking the trough to the south and the north. Whether the trough was initiated following deposition of the Nosib or prior t o deposition of the Nosib, is uncertain, but it was certainly operational towards the close of the Nosib times. The shelf facies sequences of the Swakop Group surround the trough, which closes t o the northeast, and they show rapid facies transitions into the trough on all sides. The northern edges of the Trough are marked by rapid facies transitions and are extensively invaded by granites (see later). There is evidence that a major fault scarp existed there prior t o deposition of the Damara rocks and that the zone remained unstable subsequently. Martin and Porada (1977a,b) had referred t o this zone as the “Okahandja lineament”. The trough can be subdivided into two major zones, trending parallel to the northeasterly trending margins of the trough. The northwestern half of the trough is characterized by a thick development of greywacke and subgreywacke which in places is only weakly metamorphosed and shows typical development of laminites and turbidite-type sedimentary sequences, characterized by recognizable graded bedding and soft sediment slump structures. Towards the centre of the trough a zone of graphitic schist units separates this major greywacke sequence from the other half of the trough which is characterized by increasing proportions of shaley material (now schists) with interlayered metavolcanic material concentrated along the socalled “Matchless Amphibolite Belt” (Martin, 1965). Calc-silicate horizons and lenses are commonly developed within parts of the trough and these may represent impure marly sediments. The whole sequence in the Khomas Trough is overturned towards the southeast and it is difficult t o estimate thicknesses in the trough although Halbich (1970) has suggested a thickness of more than 10,000 m. One of the curious features of the Khomas Trough is the linear continuity of formations along the axis of the trough, the lack of continuity of formations across the trough and the almost impossible correlation of any single unit from one side of the trough t o the other. The Matchless Amphibolite Belt is interesting because it is a major zone of undoubted volcanic rocks (basic lavas, basic sills and pyroclastic rocks, varying from basaltic t o rhyolitic compositions) with which are associated cupreous pyrite deposits of the Besshi type. There is no obvious external source for these rocks, or for any of the rocks in the Khomas Trough, and there is evidence that much of the Swakop sequence in the Trough has accumulated from cannibalization of volcanic material in an incipient rift/spreading centre, coupled with material eroded from topographic highs within and adjacent t o the trough.

777

The Mulden Group The Mulden Group is confined t o the Otavi shelf region where it overlies the Tsumeb carbonate sequence unconformably. I t consists of up t o 2000 m of feldspathic quartzite, arkose, greywacke and shale, and is confined t o the northern parts of the Damara Belt. It reaches its thickest development in the Ovamboland Basin (Hedberg, 1979). In the southern parts of the Otavi Mountainland a thick sequence of phyllites occurs above a poorly eroded Tsumeb carbonate sequence. This appears to represent a volcanic episode which may be related to deposition of the Mulden Group. The Mulden sequence in the north may have been deposited contemporaneously with the Nama sequence in the south (see below).

The Nama Group The Nama Group has been investigated by Germs (1972 and 1974) and mention has already been made of the problems of correlating the Nama Group with the Damara, but available evidence clearly indicates that whatever correlation is adopted, the Nama Group was deposited at some stage during the Damara Cycle and it represents a very stable platform sequence, remarkably similar in its elements t o the Damara sequence. K/Ar dating of white micas (Ahrendt et al., 1978) has indicated a metamorphic event affecting the Nama sequence about 530 Ma, and there is thus no doubt about the Nama sequence being part of the Damara Cycle. A basal Kuibis Subgroup consisting predominantly of clastic sedimentary formations, is followed upwards by the Schwarzrand Subgroup which consists of carbonate units with interlayered clastic units. This is overlain by the predominantly clastic Fish River Subgroup. Germs (1974) points out that the provenance of the Kuibis Subgroup was to the east of the Nama basin, whilst the upper parts of the Nama sequence were derived from the north and west, i.e. the Damara Belt. Glaciogenic sedimentary rocks are developed at several stratigraphic levels within the Nama sequence, and the glaciogenic beds developed in the middle of the Schwarzrand Subgroup have been equated with the Chuos Formation of the Damara (Kroner, 1971). Hartnady (1978a), mapping in the Naukluft nappe zone, has produced evidence which suggests that the nappes were already developing at the time of commencement of the Nama deposition, which indicates that the entire Nama sequence may be a correlative of the Mulden Group in the north, as suggested by Germs (1974) and Martin (1965).

778 GRANITE GENERATION, REGIONAL METAMORPHISM AND DEFORMATION

Granite generation during the Damara Cycle appears t o have been confined to the western parts of the central shelf area and along the zone north of the Khomas Trough and followed a pattern of syntectonic mantled-gneiss dome-type reactivation, which culminated in spectacular, post-tectonic emplacement of granites and pegmatitic granites as sheet-like bodies in four major northeasterly trending zones. The most southerly zone (marking the northern edge of the Khomas Trough) was the site of the catastrophic emplacement of the Donkerhoek batholith, which disrupted and invaded the northern edges of the Khomas Trough for nearly 200 km along strike. The only other major post-tectonic batholith t o invade the Damara Belt is the Sorris-Sorris granite, which occurs close t o the intersection of the main Damara Belt with its northwesterly extensions in Kaokoveld (Fig. 10.18). Regional metamorphism and much of the deformation in the Damara Belt is clearly related to two thermal events - the mobilization of the basement and the lower parts of the Damara cover sequence and related generation of syntectonic granites, and the emplacement of post-tectonic granite bodies towards the close of the Damara Cycle. These events appear to have commenced sometime between 750 Ma and 650 Ma and continued progressively and intermittently up t o 450 Ma (Clifford, 1967; Kroner et al., 1978) when the last pegmatites and pegmatitic granites were intruded in the central Damara Belt. The deformation, regional metamorphism and the generation of granites during the Damara Cycle, appears to be related to zones of weakness established prior t o the deposition of the Nosib Group. It is significant though, that no major thermal reactivation of the basement occurred on the southern side of the Khomas Trough, and granites do not appear to have been generated anywhere along the southern margins of the Damara Belt during the Damara Cycle in any major bodies. The Damara granites can be divided into three major types which are a Red Gneissic Granite type, the Salem Granite suite and the post-tectonic granites. The Red Gneissic Granites are confined t o areas where reactivated Basement cores have partially melted the lower parts of the Nosib sequence to form red gneisses and mobilized granitic material in arkosic metasedimentary units. However, it would appear that much of what has been mapped as Red Gneissic Granite is in fact remobilized Basement material. These granites are thus confined t o the west-central parts of the Damara Belt where partial melting of the Basement has occurred. Most of the areas underlain by Red Gneissic Granite contain relict cores of Basement rocks. The Salem Granite suite is predominantly granodioritic in composition, and it developed syntectonically and was still locally mobile after the deformation and regional metamorphism of the Damara sequence was completed. The generation and emplacement of the Salem suite appears t o have taken

779 place over an extended period between 750 Ma and 550 Ma, and field relationships suggest that the Red Gneissic Granites also developed during the early part of this time span. There is a need for more detailed and definitive radiometric dating on the Damara granites as a whole. The Salem Granites are confined to the zone north of the Khomas Trough between Omaruru and Swakopmund, and this type of granite is also developed in the Kaokoveld Trough zone. The Salem Granite suite exhibits a great variety of structurally conformable, to discordant, and obviously intrusive, relationships with the Damara sequence. Smith (1961), Martin (1965), Miller (1972), Jacob (1974, 1978), all considered the Salem Granites t o have evolved by partial melting of Khomas schists in the upper parts of the Swakop sequence, because in many areas the Salem Granite suite seems t o be confined t o synformal areas partly occupied by these schists. However, the regional setting of the Salem Granites, their relationship to Basement and the Damara cover, and the structural metamorphic style developed in the areas intruded by the granites, are such that they must have been generated from mobile Basement sources rather than the Khomas schists, a point which is now generally accepted. The third major group of Damara granites are the major post-tectonic batholiths - the Donkerhoek and Sorris Sorris batholiths and associated pegmatitic granite sheets and pegmatites which have developed in four major northeasterly-trending zones. Miller (1972) has shown that the SorrisSorris batholith is in fact a composite granitoid mass consisting mainly of Salem-type granites intruded by smaller plutons of post-tectonic granite. These post-tectonic granites are predominantly adamellitic and granitic in composition and appear t o have been intruded in the period between 550 Ma and 450 Ma. Whereas a sequence of regional metamorphism and deformation characterized the development of the Red Gneissic Granites and the Salem suite, the intrusion of the post-tectonic granites has produced a major regional contact metamorphism which may obscure the earlier regional imprint in the central Damara Belt. The effects of metamorphism and deformation are very variable throughout the Damara Belt but they follow similar patterns typical of other Precambrian mobile belts (Anhaeusser et al., 1969; Windley, 1977). The metamorphic and structural styles of the various parts of the Damara Belt, are intimately related to the tectonic elements which controlled the initiation and development of sedimentation and volcanism during the Damara Cycle. In the more stable shelf areas, reactivation of the basement has tended to result in mild flexuring and low-grade dislocation metamorphism. More extreme dislocation metamorphism has occurred as a result of basement slices being thrust into the Damara cover rocks along the southern margins of the Damara Belt, and along the eastern edges of the Kaokoveld Trough, and major regional shear zones have developed in both these situations. In the areas where the Damara granites were generated (Fig. 10.18) mantled-gneiss dome-type structures are common and early tight folds are refolded into

780 major dome and basin structures, and metamorphism under amphibolite and even granulite facies conditions is the norm. It is clear that there were great variations in the physical conditions under which metamorphism occurred in these different environments. A pattern of early mineral growth related to the development of schistosity and early intra-folial folds is evident over most of the Damara Belt. This was followed by continued mineral growth, porphyroblast development and subsequent deformation during widespread cleavage development related to later phases of folding. Post-tectonic mineral growth over pre-existing fabrics is particularly widespread in the central Damara Belt, and is emphasized in zones where post-tectonic granites have left a superimposed contact metamorphic imprint. This pattern is common to both the Damara cover rocks and their enclosed ore deposits, and can be recognised in the cupreous pyrite deposits of the Matchless amphibolite belt and in many small sulphide deposits which have been involved in the Damara Cycle. Deformation throughout the belt can be regarded as a variety of responses to various types of reactivation of the granitoid Basement Complex. In the more plastically mobile areas, the Damara cover sequence appears to have behaved passively and has been moulded and intruded by the partially melted Basement, whereas in the less mobile areas, brittle deformation has often led to detachment of the cover sequence from the Basement and the development of large scale disharmonic folding, sliding and nappe formation. This variation is accentuated by local variations in the thickness of the Damara rocks, and bv the n o s s ductility contrasts between the Nosib clastic sequences and the overlying carbonate formations of the Otavi and Swakop sequences, and ductility contrasts within the latter themselves. Some of the most spectacular folding in the Damara Belt occurs in the Brandberg West area of Damaraland, where the Swakop sequence is very thin and consists of alternating carbonate and clastic units (Jeppe, 1952). Here the cover rocks appear to have become detached from the Basement and have crumpled into a remarkable series of tight, long limbed folds which form a separate structural province within the Damara Belt. One of the most striking features of the Damara Belt is the change of structural style from the central Swakop shelf to the Khomas Trough (Gevers, 1963; Martin, 1965; Blaine, 1978). The dome and basin pattern of folding of the shelf area is emphasised by the development of mantled gneiss domes, and the ductility contrasts between the carbonate units of the Swakop Group and the interlayered clastic units. This gives way southwards across the Donkerhoek batholith t o the strong northeasterly trending linearity of the Khomas Trough, with its constant northwesterly dips. The schists and more massive sedimentary and volcanic units lie stacked across the Trough like a pack of playing cards, and fold closures are difficult to detect in the monotonous sequences of schists. Hiilbich (1970) has described the structural elements in the Trough and suggested southward directed com-

781 pression of the Khomas schists against a rigid southern foreland area. The southern margins of the Damara Belt are marked by major thrusts, nappe complexes and shc i belts (Martin, 1965; Schalk, 1973; Martin and Porada 1977a,b; Hartnady, 1978b). Similar structure occur on the eastern margins of the Kaokoveld Trough (Guj, 1970a), where major nappe complexes have been thrust eastwards against the Basement Complex from the edges of the Kaokoveld Trough, which are further defined by major ductile shear belts extending into southern Angola, and the intrusion of granite sheets which themselves have been severely deformed in the shear belts. ORE DEPOSITS IN THE DAMARA BELT

Ore deposits in the Damara Mobile Belt reflect the pattern of crustal evolution outlined so far in this account, and in fact the development and distribution of these ore deposits not only provide further clues to the identity of the various Damara environments, but also impose some constraints on the tectonic mechanisms which may be envisaged as a framework for the evolution of the Damara Belt. The cover rocks of the Basement Complex contain a number of mineralized volcanic units in the Rehoboth, Abbabis and Kamanjab inliers and in southern Angola. Mineralization is usually associated with exhalite beds (ferruginous metaquartzites, magnetite quartzites and meta-cherts) and so far exploration has revealed pyritic copper-zincsilver mineralization in these situations. Good examples of this type occur in the Abbabis inlier at the old Henderson and Ubib prospects, where the mineralization is intimately associated with drag folds in thin magnetite-quartzite horizons preserved in fragments of the Basement cover sequence. The Irumide cover sequences associated with the Rehoboth-Ghanzi ridge, flanking the southern margins of the Damara Belt, contain a number of sedimentary copper deposits in Namibia and northwestern Botswana. In addition, some of the basic amygdaloidal lavas in the Dordabis area contain disseminated copper mineralization. Klein Aub mine is situated in a small basin of sedimentary rocks flanking the southern margins of the Rehoboth Basement inlier. The copper mineralization occurs in seven distinct bands in a sequence of finely laminated, well-bedded argillaceous and arenaceous units (see p. 767) and consists of finely disseminated chalcocite with associated silver. The Witvlei copper deposits and extensive copper mineralization near Lake Ngami in northwestern Botswana (Toens, 1975; Cole and Le ROUX, 1978) are very similar in type and age t o the Klein Aub deposits. In general the volcanic units of the Irumide Cycle appear to have been evolved in sub-aerial, predominantly terrestrial environments. The apparent lack of sub-aqueous features, coupled with the predominance of porphyry units and the lack of exhalite beds in the region under consideration, are factors which combine t o limit the potential of these volcanic units as base

782 metal targets. Likewise the Irumide granites appear to lack the extremely fractionated representatives of the granite suite which would hold potential for the concentration of incompatible metals (Sn, W, Mo, U etc.). Smith (1965) and Martin (1965, 1978) both realized the links between a developing Damara “geosyncline” and base metal mineralization, and remarked on the concentration of copper, lead and zinc mineralization in the upper parts of the Nosib sequence and the lower parts of the Otavi (Abenab Subgroup) and Swakop (Ugab and Kudis Subgroups) sequences. In fact, there are numerous small, sedimentary copper deposits and carbonate leadzinc-vanadium deposits which occur in these stratigraphic positions. The Oamites copper mine, situated in the southern shelf region t o the south of Windhoek, is situated on a small, low-grade (1.2% Cu) copper deposit which occurs in fluviatile conglomerates and feldspathic quartzites close to the Nosib-Swakop Group boundary (Lee and Glenister, 1976). There are several other copper deposits and occurrences of this type in the southern shelf region, the Nosib depositories of the central Damara, and in the Otavi shelf region. The Abenab and Berg Aukas lead-zinc-vanadium deposits occur in recrystallized dolomites of the Abenab Subgroup (Otavi Group) on the flanks of the Grootfontein Basement inlier. They are typical karst-type deposits and contain rich concentrations of secondary lead (cerrussite), zinc (willemite) and vanadium (descloizite and vanadinite). The vanadium was introduced into the karst system at a later stage than the lead-zinc mineralization. Karst processes are still active in the Otavi Mountainland, and circulating groundwaters redeposit crusts of secondary lead, zinc and vanadium minerals in these deposits. The Elbe and Rudenau coppei-zinc deposits occur in the Nosib sequence of the shelf transition zone along the edges of the Khomas Trough, and appear to be small volcano-sedimentary deposits associated with metamorphosed exhalite beds. It is interesting t o note that the northern margins of the Khomas Trough (the edges of the central shelf) are geochemically “active” compared to the Trough itself, and there is a pronounced build-up of heavy metals in this zone, quite apart from the deposits at Elbe and Rudenau. The restricted nature of the tectonically active Nosib-lower Otavi/ Swakop depositories provided ideal situations for the generation of base metal deposits of both sedimentary and volcano-sedimentary types. These restricted but active depositories, with their flanking Basement “highs” and the irregular nature of the Basement floor, permitted the development of near-shore lagoonal facies in clastic units, stromatolitic reef complexes in carbonate units, and a whole range of impure carbonaceous and fermginous sedimentary horizons, which collectively form an ideal environment for the deposition of a wide range of sedimentary copper, lead and zinc mineralizations. The potential for further discoveries in the Otavi shelf region, parts of

783 the central shelf (where Nosib rocks are developed), and the southern shelf region is considerable. ‘ The Chuos Formation marks another metal build-up in the Damara belt, and although this reprksents a very small vertical stratigraphic interval, it is so widespread across the Damara belt that its regional significance cannot be overemphasised. Indeed, in many areas recognition of the Chuos Formation is facilitated by the development of iron- and manganese-rich beds, which in some cases (in the Kaokoveld, Otjosondu and at several places on the southern shelf region), form substantial beds of ironstone and manganiferous ironstone (Martin, 1965, 1978). Martin (1978) draws attention to the distribution of the major iron and manganese accumulations around pre-Damara basement highs. The Chuos tilloids and associated graphitic schists are invariably pyritic to a greater or lesser extent, and traces of copper, lead and zinc occur at numerous localities. The Tsumeb orebody (Sohnge, 1964) is one of the world’s great polymetallic orebodies (lead, zinc, copper, silver, cadmium and germanium), and it occurs in an irregular pipe-like brecciated body in the upper Otavi (Tsumeb Subgroup) dolomites of the Otavi shelf. Although the origin of the metals is obscure, there are many features which suggest that the pipe-like structure originated as a large pothole, and Sohnge (1974), has suggested a possible palaeophysiographic model for the Tsumeb orebody and similar orebodies developed in the Otavi valley around Kombat. The unconformity separating the Tsumeb dolomite sequence from the overlying Mulden clastic sequence (and the Otavi valley phyllites), is considered to represent a major erosional break, during which karst formation occurred with extensive development of karst fissures, cavities and potholes, and related breccias. The “pseudo-aplite” at Tsumeb and Kombat appears to represent sand and gravel washed into the karsts during the period in which they were mineralized and subsequently cemented by carbonate and silica. Similar rocks are found associated with mineralized karst structures at Berg Aukas and Abenab (in the Abenab Subgroup), and indeed “pseudo-aplite” and the related brecciation are very much prospecting guides in exploration for lead-zinc mineralization in the Otavi Mountainland. Two small lead deposits occur in the Karibib marbles of the Khomas Subgroup in the Central Shelf region between Karibib and Swakopmund, the Namib lead mine (operational) and the Usakos lead mine (disused). Both deposits occur in a very prominent and widely developed white marble unit which contains sulphide-rich beds and is anomalously rich in lead and zinc along a strike of over 100 km. At Namib lead mine the ore has been metamorphosed, tectonized and remobilized into a series of disjointed, small, but rich pods of sphalerite, galena (with associated silver) and pyrrhotite. Apart from the fact that these deposits are stratabound, it is significant that they occur in favourable environments close t o the Basement “highs” which punctuate the central shelf region and which gave rise to many of the

784 mantled gneiss dome structures during the regional metamorphism and granite generation of the Damara cycle. Cupreous pyrite deposits occur in association with the Matchless amphibolite belt in the Khomas Trough (Martin, 1965, 1978), and include numerous small pyritic deposits < 5 million tonnes in the Gorob, Matchless and Otjihase areas. The Otjihase orebody (Goldberg, 1976) is by far the largest single orebody along the belt and has probable reserves of over 10 million tonnes. The major metal, apart from iron, is copper in the form of chalcopyrite, and this contains minor but recoverable silver. Sphalerite is usually present in minor amounts but may be locally concentrated. Where the grade of metamorphism is higher, pyrrhotite replaces the dominant sulphide species pyrite. Without exception the known orebodies are contained in lensoid ferruginous chert beds (magnetite quartzites), or are very closely associated with these beds in adjacent magnetite-bearing chloritic and sericitic schists. The copper-rich ore shoots tend to be preserved in tight fold closures. All the known orebodies occur on the south side (structurally lower) of the amphibolite belt and tend t o be spatially separated from the main amphibolites by between 100 and 500 m of sericitic, chloritic and biotitic schists. There is evidence of inversion of the orebodies in most instances and the amphibolite belt as a whole appears to be overturned t o the southeast. The amphibolite belt itself is a continuous zone of metavolcanic units which form an almost straight, northeasterly trending belt through the Khomas Trough. Basic units define the belt and consist of metabasalts associated basic tuffs and layered sills, and these are flanked to the south by metavolcanic schists of varying composition with associated exhalite beds (metacherts). The metabasalts include amygdaloidal and pillowed types, and preservation of original volcanic textures and structures is particularly good in the Gorob area. The layered sills consist of gabbro-pyroxenite units, in which the original gabbroid textures are well preserved, but the pyroxenitic units tend to become more sheared and in some cases they degenerate into talc-tremolite schists. Finnemore (1978) has presented chemical evidence that the amphibolites of the Matchless belt are predominantly of basaltic parentage and that they can be classed as oceanic tholeiites. Mineralization associated with the Damara granites in the central Damara Belt includes tin mineralization, a variety of pegmatite mineralization, and uranium mineralization, all of which appear t o be associated with extremely fractionated post-tectonic granite, pegmatitic granite and pegmatite sheets and dykes. Small deposits of beryl, pollucite, columbite-tantaIite, lepidolite, petalite, feldspars and micas occur in zoned pegmatites, which are concentrated in a series of northeasterly trending linear zones between the northern edge of the Khomas Trough and the Sorris-Sorris granite. The so-called “tin belts’’ are related t o four of these zones and their tectonic significance is further underlined by the widespread occurrence of fluorite in pegmatites and greisens.

7 85

Tin mineralization occurs in unzoned pegmatites and pegmatitic granite bodies as weak disseminations and in thin greisen veins, pockets and stockworks, and is occasionally found as skarn-type mineralization where the pegmatitic granite has intruded marbles of the Swakop Group. The Uis tin mine is situated on large unzoned, weakly mineralized pegmatites of the most northerly tin belt. The Brandberg West mine is a tin-tungsten producer situated to the north of the main tin belts. The deposits consist of stockworks of quartz-feldspar veins which are apparently remote from any granitic source, and although the veins themselves contain up to 1%combined tin and tungsten (cassiterite and scheelite), the stockworks are poorly developed with veins widely separated, and the veins are generally narrow (2-50 cm). It is interesting to note that the other tungsten producer in S.W.A./Namibia, the Kranzberg mine near Omaruru, is a wolframite deposit associated with greisens formed at a late stage in the emplacement of the post-Karoo Erongo granite complex. The deposit falls on the most southerly of the tin belts which has been subsequently intruded by the post-Karoo Spitzkoppe granites and the Erongo complex. Obviously the weakness established along this zone during the evolution of the Damara belt persisted as a locus of igneous activity after the close of the Damara Cycle. Uranium mineralization associated with post-tectonic pegmatitic granites is widespread in the area north of the Khomas Trough between Swakopmund and the Abbabis inlier. The Rossing mine is already a major producer of uranium and there are numerous small deposits and prospects in the Rossing area near Swakopmund. There is a close relationship between the mineralization and intrusion of the pegmatitic granite into the Rossing Formation (Ugab Subgroup of the Swakop Group) around major mantled gneiss domes, and this was first recognized by Smith (1965) and described more fully by Berning et al. (1976), Jacob (1978) and Von Backstrom and Jacob (1979). Thus the uranium mineralization appears to have two major controls on its distribution in central Damara Belt, namely the development of the lowest parts of the Swakop Group -the Rossing Formation, and the presence of post-tectonic pegmatitic granites (the so-called “alaskites” at Rossing mine). The significance of the Rossing Formation and the beds immediately below (the Khan Formation at the top of the Nosib Group) appears to be that it provided a suitable environment for precipitation of uranium from extremely fractionated uraniferous granitic melts. In addition it would appear that the extensive carbonate units developed in the upper parts of the Rossing Formation and in the Khomas Subgroup, acted as barriers to the late pegmatitic granite bodies which tend t o congregate below the carbonate units, The Rossing Formation itself contains abnormal concentrations of pyrite pyrrhotite and graphite, and it is significant that where the mineralized granites have intruded these rocks, uranium has crystallised in the form of uraninite. Jacob (1978) has presented a more detailed account of the

786 origin and concentration of uranium in the pegmatitic granites of the Rossing area and he suggests that release of sulphur during metamorphism created the necessary reducing conditions for precipitation of uraninite. The Rossing granites have been dated by the Rb-Sr method at 468 5 8 Ma (Kroner and Hawkesworth, 1977). The uranium may be remobilized from the basal Nosib beds or from the unconformity at the base of the Nosib, or it may have had a Basement source. Whatever the case, the location of uranium in this area now appears to be related to a chain of events which commenced prior to deposition of the Damara and which had its foundation in the initial tectonic configuration of central Namibia. Reactivation of the Basement Complex led t o the regional metamorphism, deformation and generation of a variety of granitic rocks, culminating in the intrusion of post-tectonic granites at the close of the Damara Cycle. This evolutionary pattern is confined to a relatively small part of the Damara Belt, close to the intersection of the Khomas Trough and related northeasterly trending lineaments (the “tin belts”), with the South Atlantic rift system (Fig. 10.18). THE TECTONIC FRAMEWORK O F THE DAMARA MOBILE BELT

In the introduction to this account, the gross setting of the Damara Mobile Belt was mentioned with respect to extensions of the East African rift system, and the faulted, aseismic Atlantic seaboard of Namibia. The latter is related to the great Atlantic rift system developed as a prelude t o the separation of Africa and South America and the formation of the present Atlantic Ocean. Figure 10.18 illustrates these elements and outlines a tectonic subdivision of the Damara Belt. The evolution of the Damara Mobile Belt appears to have been dominated by sporadic mobility in narrow, linear rifted zones parallel to the northeasterly trending East African rift system and north-northwesterly trending Atlantic rift system, up to the initiation of the Late Precambrian Damara sequence itself. The incursion of epicontinental seas and the deposition of shelf-type sedimentary rocks in the upper Damara (the Otavi, and Swakop Groups), and the development of the Khomas Trough and the Kaokoveld Trough, during the Damara Cycle, were events which had no parallel in pre-Damara times. The disposition of Basement inliers in the Damara Belt is of considerable tectonic significance, and Martin (1965) appears to have been alone at that time in recognizing this, and realizing that the Basement configuration had played a major role in the evolution of the Damara Belt. Pre-Damara relief certainly played a major role in determining the disposition of sedimentary facies within the larger framework of the Damara depository. Kroner (1977) and Martin (1977) have emphasized the Basement inliers of the Damara Belt as evidence that the Damara Belt evolved as an “ensialic” feature. There is no doubt from the evidence available that the Basement cover sequences, the

787 cover sequences of the Irumide Cycle, and the Damara sequence itself, all evolved on sialic crust in epicontinental situations. And yet several authors have recently suggested a plate tectonic continental collision model for the Damara Belt (Blaine, 1978; Sawyer, 1978; Kasch, 1979), and Watters (1978) suggested that the Lrumide Cycle in Namibia developed along subduction zones adjacent to a continental margin -his “Rehoboth magmatic arc”. There is little doubt that the local geology of the Khomas Trough, and particularly its tectonised southern margins, and its tectono-thermally mobilised northwestern margins, suggest some type of initial rifting and possibly limited spreading, followed by closure, compression and the tectonic and thermal reactivation which can be seen in the record at present. However, a glance at Fig. 10.18 will place this in perspective, for the Khomas Trough and its margins form only a part of the Damara Belt as a whole, and the shelf areas cover a much larger portion of the belt where no spectacular reactivation of the Basement has occurred, and where there has only been limited tectono-thermal mobility during the Damara Cycle. The partially preserved Kaokoveld Trough exhibits many of the features of the Khomas Trough, especially with regard to the tectonization of its eastern margins and the development of major nappe complexes there. The Khomas Trough peters out in the vicinity of Steinhausen, and the Trough is surrounded by a series of small Basement inliers which punctuate a thinly developed shelf facies linking the central shelf with the southern shelf (Fig. 10.18). Indeed the Khomas Trough, with its fault-bounded rifted margins terminating towards the continental interior, is a fine example of an aulacogen in its own right. I t is difficult to visualize any major plate movement having been involved in its evolution, and the idea of anything more than a limited “rebound” type of collision, as the Khomas Trough closed up towards the end of the Damara Cycle, is not substantiated by any of the available evidence. The limited distribution of Damara granites (Fig. 10.18) and especially the post-tectonic granites, suggests that whatever mechanism was involved in their generation, it was of limited effect, and of local significance in the context of the Damara Belt as a whole. But why was the generation of Damara granites concentrated in the central Damara and t o the northwest of the Khomas Trough? A review of the evidence provides some insight into this question and to the wider issues concerning the evolution of the Damara Belt as a whole. The southern margins of the Damara Belt are marked by a major Basement ridge, the development of a Damara shelf facies over this ridge (the southern shelf), the development of major shear belts and faults bounding and transecting the Basement ridge, and by the development of two major zones of thrusting and nappe structures which flank the southern margins of the Khomas Trough and the southern margins of the southern shelf. The Khomas Trough is a linear zone of schisted, flysch-type sedimentary rocks with interlayered volcanic units, consistently overturned towards the south-

788 ern margins of the Damara Belt, in sympathy with the marginal thrusts and nappes. The Nama platform sequence developed during the Damara Cycle prior to the intrusion of the post-tectonic granites and was deposited in a stable shelf environment. It is relatively undisturbed except in the Naukluft area where Korn and Martin (1959) first recognized major nappe structures. Finally the Damara granites are virtually confined t o the northwest of the Khomas Trough between Okahandja and Swakopmund, and t o the Kaokoveld Trough. It would appear that the region south of the Khomas Trough remained relatively stable during much of the Damara Cycle, and that compression from a northerly direction produced the linear, overturned structures of the Trough and the thrusts and nappe structures along the southern margins of the Trough, as suggested by Hiilbich (1970). The extensive shearing parallel to the Khomas Trough in the southern shelf region is probably related to this movement coupled with reactivation of major northeasterly trending rift faults. The thermal reactivation of the crust to the northwest of the Khomas Trough and the generation of granites there, suggests that the same mechanism which produced the deformation in the Khomas Trough and the areas south of the Trough, could be responsible for partial melting of the crust and the generation of the Damara granites. The flysch fill of the Khomas Trough does not appear to have had an external source, and the evidence suggests that it was derived from interal degradation of intermediate and basic volcanic rocks developed along an aborted spreading centre. The Matchless amphibolite belt and its associated cupreous pyrite deposits provide additional support for this argument. The amount of spreading which occurred must have been very limited, judging from the strike length of the Khomas Trough, the surrounding shelf environments, and its present width of recognisable flysch and metavolcanic material. There could have been no major subduction zone developed, and the generation and localisation of the Damara granites could easily have resulted from limited reactivation of a southward moving “flap” of crust closing the Khomas Trough, compressing the rocks in the Trough, and overriding the Trough to a limited extent. As for the remainder of the Damara Belt, the “aulacogen model” presented by Martin and Porada (1977), more than adequately explains the observable field relationships in the Belt. REFERENCES

Part A Bahnemann, K.P., 1970. A note on the anorthositic gneiss in the Messina District, Northern Transvaal. Geol. SOC.S. Afr. Spec. Publ., 1: 715-720. Bahnemann, K.P., 1971. In: Excursion guidebook. Symposium o n the granites, gneisses and related rocks of Rhodesia (Compiled by E.R. Morrison and J.F.Wilson). Geol. Soc. S. Afr., pp. 18-22.

789 Bahnemann, K.P., 1972. A Review of the Structure, the Stratigraphy and Metamorphism of the Basement-Rocks in the Messina District, Northern Transvaal. Ph.D. Thesis, University of Pretoria (unpubl.). Bahnemann, K.P., 1973. The origin of the Singelele granite gneiss near Messina, Northern Transvaal. Geol. Soc. S. Afr. Spec. Publ., 3: 235-244. Barton, J.M. and Ryan, B., 1977. A review of the geochronologic framework of the Limpopo Mobile Belt. Proceedings of a seminar on the Limpopo Mobile Belt. Geol. Surv. Botswana Bull. 9. Barton, J.M., Fripp, R.E.P., Horrocks, P. and McLean, N., 1977a. The geology, age and tectonic setting of the Messina Layered Intrusion, Limpopo Mobile Belt. Proceedings of a seminar on the Limpopo Mobile Belt. Geol. Surv. Botswana Bull. 9. Barton, J.M., Fripp, R.E.P. and Ryan, B., 1977b. Rb/Sr ages and geological setting of ancient dykes in the Sand River area, Limpopo Mobile Belt, South Africa. Nature, 267 : 487-490. Bennett, J.D., 1970. Craton-mobile belt relations with particular reference to the Mosetse-Matsitama area, northeastern Botswana. Geol. Mag., 107(2): 113-123. Bennett, J.D., 1971. The tectonometamorphic complex of eastern Botswana. Publ. Geological map 1 : 500,000 and unpubl. rep. GeoI. Surv. Botswana. Burger, A.J. and Coertze, F.J., 1973. Radiometric age measurements on rocks from southern Africa to the end of 1971. S. Afr. Geol. Surv. Bull. 58. Coomer, P.G., Coward, M.P. and Lintern, B.C., 1977. Stratigraphy, structure and geochronology of ore leads in the Matsitama Schist Belt of northern Botswana. Precambrian Res., 5: 23-41. Coward, M.P., James, P.R. and Wright, L., 1976. Northern margin of the Limpopo Mobile Belt, southern Africa. Geol. Soc. Am. Bull., 87: 601-611. COX, K.G., Johnson, R.L., Monkman, L.J., Stillman, C.J., Vail, J.R. and Wood, D.N., 1965. The geology of the Nuanetsi Igneous Province. Philos. Trans. R. Soc. London, Ser. A, 257: 71-218. Crockett, R.N., 1965. The Geology of the country around Mahalapye and Machaneng. Rec. Geol. Surv. Bech., 1961-1962: 77-99. Crockett, R.N., 1968. Brief explanation of the geology of the Shashe area, Rep. Geol. Surv. Botswana. Davies, R.D., Allsopp, H.L., Erlank, A.J. and Manton, W.I., 1970. Sr-isotopic studies on various layered mafic intrusions in southern Africa. Geol. SOC.S. Afr. Spec. Publ., 1: 576-593. Du Toit, M.C. and Van Reenen, D.D., 1977. The southern margin of the Limpopo Mobile Belt, Northern Transvaal, with particular reference to metamorphism and structure. Proceedings of a seminar on the Limpopo Mobile Belt. Geol. Surv. Botswana Bull. 9. Ermanovics, I.F., 1977. Post-tectonic remobilization along the southern margin of the Limpopo Mobile Belt. Proceedings of a seminar on the Limpopo Mobile Belt. Geol. Surv. Botswana Bull. 9. Gerrard, I., 1965. The geology of the West Tuli area. Rec. Geol. Surv. Bech,, 1961-1962: 5-23. Gifford, A.C., 1961. The Geology of Eastern Marangudzi, Southern Rhodesia. Ph.D. Thesis, University of London (unpubl.). Gordon, P.S.L., 1973. The Selebi-Pikwe nickel-copper deposits, Botswana. Geol. SOC.S. Afr. Spec. Publ., 3: 167-187. Hickman, M.H., 1976. Geochronological Investigations in the Limpopo Belt and Part of the adjacent Rhodesian Craton. Ph.D. Thesis, University of Leeds (unpuhl.). Hickman, M.H., 1978. Isotopic evidence for crustal reworking in the Rhodesian Archaean Craton, Southern Africa. Geology, 6 : 214-216. Hickman, M.H. and Wakefield, J., 1975. Tectonic implications of new geochronologic

790 data from the Limpopo Belt at Pikwe, Botswana, southern Africa. Geol. SOC. Am. Bull., 86: 1468-1472. Hor, A.K.,Hutt, D.K., Smith, J.V., Wakefield, J. and Windley, B.F., 1975. Petrochemistry and mineralogy of early Precambrian anorthositic rocks of the Limpopo Belt, southern Africa. Lithos, 8: 297-310. Jacobsen, J.B.E., 1967. The Geology of the Artonville Mine, Messina District. Ph.D. Thesis. University of Witwatersrand (unpubl.). Jacobsen, J.B.E. and McCarthy, T.S., 1976. An unusual hydrothermal copper deposit a t Messina, South Africa. Econ. Geol., 7 1 : 117-130. Jacobsen, J.B.E., McCarthy, T.S. and Laing, G.J., 1976. The copper-bearing breccia pipes of the Messina District, South Africa. Miner. Deposita, 11: 33-45. James, P.R., 1975. A Deformation Study across the Northern Margin of the Limpopo Belt, Rhodesia. Ph.D. Thesis. University of Leeds (unpubl.). Jennings, C.M.H., 1963. The geology of the Makhaware hills area. Rec. Geol. Surv. Bech., 1959-1960: 22-33. Key, R M., 1977. The geological history of the Limpopo Mobile Belt, based on the field mapping of the Botswana Geological Survey. Proceedings of a seminar on the Limpopo Mobile Belt. Geol. Surv. Botswana, Bull. 9. Key, R.M. and Hutton, S.M., 1976. The tectonic generation of the Limpopo Mobile Belt, and a definition of its western extremity. Precambrian Res., 3 : 79-90. Key, R.M., Litherland, M. and Hepworth, J.V., 1976. The evolution of the Archaean crust of northeast Botswana. Precambrian Res., 8: 375-413. Lear, P A . , 1971. A Geological Investigation of Some Precambrian Ultramafic Rocks in eastern Botswana. M.Sc. Thesis, University of Manchester (unpubl.). Light, M.P.R., Broderick, T.J. and Watkeys, M.K., 1977. Preliminary report on the central zone of the Limpopo Mobile Belt, Rhodesia. Proceedings of a seminar on the Limpopo Mobile Belt. Geol. Surv. Botswana. Bull. 9. Mason, R., 1965. The geology of the Macloutsie area. Rec. Geol. Surv. Bech., 19611962: 25-41. Mason, R., 1967. The Geology of the Country around Baines Drift, Eastern Botswana. M.Sc. Thesis, University of Newcastle-upon-Tyne (unpubl.). Mason, R., 1970. The Geology of the Country between Francistown and Madinare, Northeastern Botswana. Ph.D. Thesis, University of Witwatersrand (unpubl.). Mason, R., 1973. The Limpopo Mobile Belt, southern Africa. Philos. Trans. R. SOC. London, Ser. A, 273: 463-485. McCarthy, T.S. and Jacobsen, J.B.E., 1976. The mineralising fluids at the Artonvilla copper deposit: An example of a silica-deficient, alkaline hydrothermal system. Econ. Geol., 71: 131-138. Odell, J., 1975. Explanation of the geological map of the area around Bangala Dam. Rhod. Geol. Surv. Short Rep. 42. Reeves, C.V. and Hutchins, D.G., 1975. Crustal structures in central southern Africa. Nature, 254: 408-410. Robertson, I.D.M., 1968. Granulite metamorphism of the Basement Complex in the Limpopo Metamorphic Zone. Geol. SOC.S. Afr. Annex. 71 : 125-133. Robertson, I.D.M., 1973a. Metamorphism and Intrusion in Part of the Southern Edge of the Rhodesian Craton and the North Margin of the Limpopo Mobile Belt. Ph.D. Thesis, University of London (unpubl.). Robertson, I.D.M., 197313. Potash granites of the southern edge of the Rhodesian Craton and the northern granulite zone of the Limpopo Mobile Belt. Geol. SOC.S. Afr. Spec. Publ., 3: 265-276. Robertson, L.D.M., 1973c. The geology of the country around Mount Towla. Rhod. Geol. Surv. Bull. 68.

791 Robertson, I.D.M., 1974. Explanation of t h e geological m a p of t h e country south of Chibi. Rhod. Geol. Surv. Short Rep. 41. Robertson, I.D.M. and Van Breemen, O., 1970. The Southern Satellite Dykes of t h e Great Dyke, Rhodesia. Geol. SOC.S. Afr. Spec. Publ., 1: 621-4544. Sohnge, P.G., 1946. T h e geology of t h e Messina copper mines and surrounding country. S. Afr. Geol. Surv. Mem., 40. Sohnge, P.G., le Roex, H.D. and Nel, H.J., 1948. The geology of t h e country around Messina. Explanation of sheet 46. Geol. Surv. Dep. Min. S. Afr. Swift, W.H., 1962. The geology of t h e Middle Sabi Valley. South. Rhod. Geol. Surv. Bull., 52. Swift, W.H., White, W.C., Wiles, J.W. and Worst, B.G., 1953. The geology of the Lower Sabi Coalfield. South. Rhod. Geol. Surv. Bull. 40. Thomas, C.M., 1970. Brief explanation of the geology of the Selebi area. Geol. Surv. Botswana. Vail, J.R. and Dodson, M.H., 1969. Geochronology of Rhodesia. Trans. Geol. SOC. S. Afr., 73: 79-113. Van Breemen, O., 1968. Geochronology of t h e Limpopo Orogenic Belt, Southern Africa. Ph.D. Thesis. University of Leeds (unpubl.). Van Breemen, 0. and Dodson, M.H., 1972. Metamorphic chronology of the Limpopo Belt, southern Africa. Geol. SOC.Am. Bull., 83: 2005-2018. Van Breemen, O., Dodson, M.H. and Vail, J.R., 1966. Isotopic age measurements o n the Limpopo Orogenic Belt, southern Africa. Earth Planet. Sci. Lett., 1:401-406. Van Eeden, O.R., Visser, H.N., van Zyl, J.S., Coertze, F.J. and Wessels, J.T., 1955. The geology of t h e eastern Soutpansberg and t h e Lowveld t o t h e north. Explanation of sheet 42. Geol. Surv., Dep. Min., S. Afr. Van Graan, S.J., 1964. Geology of the Messina copper deposits. In: S.H. Haughton (Editor), Some Ore Deposits in Southern Africa. Vol. 11. Geol. SOC.S. Afr., pp. 303314. Van Reenen, D.D. and Du Toit, M.C., 1977. Mineral reactions and t h e timing of metamorphic events in the Limpopo Metamorphic Complex, south of the Soutpansberg. Proceedings of a seminar o n t h e Limpopo Belt. Geol. Surv. Botswana, Bull. 9. Van Zyl, J.S., 1950. Aspects of the geology of t h e northern Soutpansberg area. Ann. Univ. Stellenbosch. 26. Sect. A: 1-95. Wakefield, J., 1974. T h e Geology of t h e Pike Ni-Cu Province, Eastern Botswana, Ph.D. Thesis. University of Leeds (unpubl.). Watkeys, 1976. The geology of t h e area west of Beitbridge. Ann. Rhod. Geol. Surv., 2: 1-14. Wilson, J.F., 1964. T h e geology of t h e country around F o r t Victoria. South. Rhod. Geol. Surv. Bull., 58. Wilson, J.F. and Harrison, N.M., 1973. Recent K-Ar age determinations o n some Rhodesian granites. Geol. SOC.S. Afr. Spec. Publ., 3 : 69-78. Windley, B.F., 1973. Archaean anorthosites: a review with t h e Fiskenaesset Complex, West Greenland, as a model for interpretation. Geol. SOC.S. Afr. Spec. Publ., 3 : 319332. Worst, B.G., 1962. The geology of t h e Buhwa iron ore deposits and adjoining country: Belingwe District. South. Rhod. Geol. Surv. Bull. 53.

Part B Benedict, P.C., Wiid, D. d e N., Cornelissen, A.K. a n d staff, 1964. Progress report o n t h e geology of t h e O’okiep Copper District. In: S.H. Haughton (Editor), T h e Geology of

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793 Joubert, P., 1974b. The gneisses of Namaqualand and their deformation. Trans. Geol. SOC.S. Afr., 77: 339-345. Joubert, P., 1976. The realtionship between the Namaqualand Metamorphic Complex and the Kheis Group. S. Afr. J. Sci., 72: 312-313. Kroner, A., 1974. Late Precambrian formations in the western Richtersveld, northern Cape Province. Bull. Precambrian Res. Unit, Univ. Cape Town, 13: 115 pp. Kroner, 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. McDaid, J.M., 1976. Preliminary report on the geology of the northern part of Diamond Area No. 1, South West Africa. Annu. Rep. Precambrian Res. Unit, Univ. Cape Town, 13: 89-96. Martin, H., 1965. The Precambrian geology of South West Africa and Namaqualand. Precambrian Res. Unit, Univ. Cape Town, 1 5 9 pp. Middleton, R.C., 1976. The geology of Prieska Copper Mines Limited. Econ Geol., 71: 32 8-3 50. Moore, J.M., 1976. The geology of Namiesberg, Northern Cape. Bull. Precambrian Res. Unit, Univ. Cape Town, 20: 69 pp. Nicolaysen, L.O., 1962. Stratigraphic interpretations of age measurements in Southern Africa. In: A.E.J. Engel, H.L. James and B.F. Leonard (Editors), Petrologic Studies: Buddington Volume. Geol. SOC.Am., pp. 569-598. Nicolaysen, L.O. and Burger, A.J., 1965. Note on an extensive zone of 1000 million-year old metamorphic and igneous rocks in Southern Africa. Sci. Terre, 1 0 : 497-516. Reid, D.L., 1976. Geochronology of the Haib-Vioolsdrif igneous province. Annu. Rep. Precambrian Res. Unit, Univ. Cape Town, 1 3 : 4 9 - 6 0 . Rogers, A.W. and du Toit, A.L., 1908. Report on the geology of parts of Prieska, Hay, Britstown, Carnarvon and Victoria West. Annu. Rep. Geol. Comm. C.G.H., 13: 8-127. Schultz, R., 1976. The origin and mode of emplacement of hypersthene-bearing granitoids and associated rocks in parts of the Upington Geotraverse. Annu. Rep. Precambrian Res. Unit, Univ. Cape Town, 13: 19-25. Schwarz, E.H.L., 1910. The correlation of the Kheis Series. Trans. Geol. SOC.S. Afr., 13: 140-146. Smit, P.J., 1973. The Waterberg System in the Northern Cape. Abstr. of paper read at 15th Congress, (Bloemfontein). Geol. SOC.S. Afr.: 47-48. Sohnge, P.G. and De Villiers, J., 1948. The Kuboos Pluton and its associated line of intrusives. Trans. Geol. SOC.S. Afr., 51: 1-36. Toogood, D.J., 1976. Structural and metamorphic evolution of a gneiss terrain in the Namaqua belt near Onseepkans, South West Africa. Bull. Precambrian Res. Unit, Univ. Cape Town, 1 9 : 1 8 9 pp. Vajner, V., 1974. The tectonic development of the Namaqua mobile belt and its foreland in parts of the Northern Cape. Bull. Precambrian Res. Unit, Univ. Cape Town, 14: 201 pp. Von Backstrom, J.W., 1960. Die geologie van die gebied om Nieuwoudtville. Explan. Sheet Geol. Surv. S. Afr., 241 pp. Von Backstrom, J.W., 1964. The geology of the area around Keimoes, with special reference to phacoliths of charnockitic adamellite-porphyry. Mem. Geol. Surv. s. Afr., 53: 206 pp. Von Backstrom, J.W. and De Villiers, J., 1972. The geology along the Orange River between Onseepkans and the Richtermeld. Explan. Sheets 2817 D, 2818 C and D and 2819 C. Geol. Surv. S. Afr., 1 0 1 pp. Watters, B.R., 1974. Stratigraphy, igneous petrology and evolution of the Sinclair Group in southern South West Africa. Bull. Precambrian Res. Unit, Univ. Cape Town, 1 6 : 218 pp.

794

Part C Arnett, O.J., 1 9 5 3 . The Geology of t h e Groenberg District, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Burger, A.J. and Coertze, F.J., 1 9 7 3 . Radiometric age measurements o n rocks from southern Africa to t h e end of 1971. Bull. Geol. Suw. S. Afr., 58 pp. Cain, A.C., 1 9 7 5 . A preliminary review of t h e stratigraphic relationships and distribution of metamorphism in the northern part of t h e Natal-Namaquarides, South Africa. Geol. Rundsch., 6 4 : 192-215. Cherry, D.W., 1 9 4 7 . T h e Geology of t h e Umgeni Dam and surrounding Country, Valley of a Thousand Hills, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Colombin, L.G.N., 1 9 6 8 . T h e Petrology of t h e Basement Rocks in t h e Lilani Area, Natal. M.Sc. Thesis, Univ. of Witwatersrand. Craig, A.J.J.L., 1 9 6 5 . T h e Geology of t h e Umzinto Area, Natal. M. Sc. Thesis, Univ. of Natal (unpubl.). &vies, R.D., 1 9 6 4 . T h e Geology of t h e Area northwest of Table Mountain, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Du Toit, A.L., 1918. T h e geology of t h e crystalline limestone deposits near Port Shepstone and Hermansberg, Natal. Geol. S u n . S. Afr. Mem., 11: 104-141. Du Toit, A.L., 1 9 1 9 . T h e geology of t h e Marble Delta. Q. J. Geol. SOC. London, 27: 53-72. Du Toit, A.L., 1 9 2 0 . T h e geology of Pondoland. Geol. Sum. S. Afr. Explan. sheet 27. Du Toit, A.L., 1931. T h e geology of the country surrounding Nkandhla, Natal. Geol. Surv. S. Afr., Explan sheet 109. Du Toit, A.L., 1 9 4 6 . T h e geology of parts of Pondoland, East Griqualand and Natal. Geol. Surv. S. Afr., Explan. sheet 1 1 9 . Gevers, T.W., 1 9 4 1 . Carbon dioxide and exhalations in northern Pondoland and Alfred County, Natal. Trans. Geol. SOC.S. Afr., 4 5 : 183-214. Gevers, T.W. and Dunne, J.C., 1 9 4 2 . Charnockitic rocks near Port Edward in Alfred County, Natal. Trans. Geol. SOC.S. Afr., 4 5 : 183-214. Gevers, T.W., 1964. Thermal springs a t Lilani, Natal and their geological setting. Trans. Geol. SOC.S. Afr., 6 7 : 35-41. Gold, D., 1957. T h e Geology of t h e Area around Pinetown and Hillcrest, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Grout, F.F., 1 9 3 5 . T h e compositions of some African granitoid rocks. J. Geol., 4 3 : 281296. Guy, J.E., 1974. The Geology of an Area to t h e northwest of Mandini, Zululand. M.Sc. Thesis, Univ. of Natal (unpubl.). Hargraves, R.B., 1 9 5 2 . T h e Geology of t h e Area north of Botha’s Hill, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Hill, J.H., 1 9 5 8 . The Geology of an Area around Mid-Illovo, Natal. M S c . Thesis, Univ. of Natal (unpubl.). Jennings, C.M.H., 1 9 6 3 . T h e Geology of the Umkomaas sheet, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Key, R.M., Litherland, M. and Hepworth, J.V., 1976. T h e evolution of the Archaean crust of northeast Botswana. Precambrian Res., 3 : 375-413. Kent, L.E., 1938. The geology of a portion of Victoria County, Natal. Trans. Geol. SOC. S. Afr., 4 1 : 1-36. Lambert, J.F., 1954. T h e Geology of a Portion of t h e Valley of a Thousand Hills. M.Sc. Thesis, Univ. of Natal (unpubl.). Lambert, J.F., 1962. T h e Petrology of t h e ultrabasic Rocks of t h e Tugela Valley, Natal. Ph.D. Thesis, Univ. of Natal (unpubl.).

795 Leith, M.J., 1966. The Geology of an Area to the south of Mapumulo, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Leith, M.J. and Rhodes, R.C., 1971. Granitic rocks in the basement complex around Mapumulo, Natal in the light of zircon studies. Trans. Geol. SOC.S. Afr., 74: 5 7 - 6 8 . Matthews, P.E., 1952. A Contribution to the Geology of the Valley of a Thousand Hills, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Matthews, P.E., 1959. The metamorphism and tectonics of the pre-Cape formations in the post-Ntingwe thrust-belt, southwestern Zululand, Natal. Trans. Geol. SOC.S. Afr., 62: 257-322. Matthews, P.E., 1972. Possible Precambrian obduction and plate tectonics in southeastern Africa. Nature, 240: 37-39. McCarthy, M.J., 1961. The Geology of the Empangeni Fault Area. M.Sc. Thesis, Univ. of Natal (unpubl.). McIver, J.R., 1966. Orthopyroxene-bearing granitic rocks from southern Natal. Trans. Geol. Soc. S. Afr., 69: 99-117. McIver, J.R. and Gevers, T.W., 1968. Charnockites and associated hypersthene-bearing rocks in southern Natal, South Africa. Int. Geol. Congr. 22nd India, 13: 151-168. Nicolaysen, L.O. and Burger, A.J., 1965. Note on an extensive zone of 1000 million-year old metamorphic and igneous rocks in southern Africa. Sci. Terre, 1 0 : 497-518. Otto, J.D.T., 1977. The geology and petrology of the Marble Delta. Ann. Univ. of Stellenbosch, 2: 249-365. Rhodes, R.C., 1966. The Geology of an Area to the north of Mapumulo, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Roberts, N.J., 1971. The Geology of an Area to the northeast of Inanda, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Robinson, D.N., 1969. The Geology of an Area northeast of Mapumulo, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Schulze-Hulbe, A., 1977. A Study of the Structure and Metamorphism of the Rocks of the Tugela Group exposed in the Mambula-Mbongolwane Area, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Scheepers, A.I.H., 1954. The Geology of the Ndwedwe-Inanda native Reserve, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Simpson, E.S.W. and Tregidga, J.A., 1956. The Archaean rocks of the Marble Delta district, Natal. Trans. Geol. SOC.S. Afr., 59: 237-258. Thompson, H.G., 1955. The Geology of the Area about Tafamasi, Inanda, Natal. M.Sc. Thesis, Univ. of Natal (unpubl.). Tuttle, O.F. and Bowen, N.L., 1958. Origin of granite in the light of experimental studies Geol. SOC.Am.. Mem., 75: 1-53. in the system NaAlSi308-KAlSi308-SiO~-H~O. White, A.J.R. and Chappell, B.W., 1977. Ultrametamorphism and granitoid gneesis. In: D.H. Green (Editor), Experimental Petrology Related to Extreme Metamorphism. Tectonophysics, 4 3 : 7-22. Williams-Jones, A.E., 1968. The Geology of an Area to the west of Glendale, Natal.M.Sc. Thesis, Univ. of Natal (unpubl.). Winkler, H.G.F., 1976. Petrogenesis of Metamorphic Rocks. Springer-Verlag, Berlin, 334 PP.

Part D Bahnemann, K.P. and Lockett, N.H., 1979. The geology and copper mineralization of the Gwai River Mine area, Wankie District, Rhodesia. Geol. SOC.S. Afr., Spec. Publ., 5: 209-217. Broderick, T.J., 1976. Explanation of the Geological Map of the Country East of Kariba. Rhod. Geol. Surv. Short Rep. No. 43.

Cahen, L. and Snelling, N.J., 1966. The Geochronology of Equatorial Africa. NorthHolland, Amsterdam, 1 6 1 pp. Ewart, A., 1960. The Geology and Mineralization of the Kamativi Tin Belt, S. Rhodesia. Ph.D. Thesis, Univ. London (unpubl.). Fick, L.J., 1960. The geology of the tin pegmatites a t Kamativi, Southern Rhodesia. Geol. Mijnb., 3 9 : 472-491. Lockett, N.H., in press. The geology of the country around Dett, Wankie District. Zimbabwe Geol. Surv. Bull., No. 85. Priem, H.N.A., Boelrijk, N.A.I.M., Hebeda, E.H., Verdurmen, E.A.Th. and Verschure, R.H., 1971. Isotopic dating in the Kamativi tin belt, Southern Rhodesia. Geol. Mijnb., 50: 6 1 9 - 6 2 4 . Rijks, H.R.P. and Van der Veen, A.H., 1972. The geology of the tin-bearing pegmatites in the eastern part of the Kamativi District, Rhodesia. Miner. Deposita, 7: 383-95. Vail, J.R. and Dodson, M.H., 1969. Geochronology of Rhodesia. Trans. Geol. Soc. S. Afr., 72: 79-114. Watson, R.L.A., 1962. The geology of the Kamativi and Lubimbi areas. Rhod. Geol. Surv. Bull., No. 57.

Part E

Ackermann, E. and Forster, A., 1960. Grundzuge der Stratigaphie und Struktur des Irumiden Orogens. 21st. Int. Geol. Congr. Norden, 18: 182-192. Amstutz, G.C. and Bernard, A.J., 1973. Ores in Sediments. Int. Union Geol. Sci., Series A, No. 3. Springer-Verlag, Heidelberg, 350 pp. Bartholom6, P. (Editor), 1974. Gisements stratiformes et provinces cuprif6res. Soci6t6 G6ologique de Belgique, 427 pp. Benham, D.G., Greig, D.D. and Vink, B.W., 1976. Copper occurrences of the Mombezhi Dome area, Northwestern Zambia. Econ. Geol., 71(2): 433-442. Bigotte, G. and Bonifas, G., 1968. Faits nouveaux sur la g6ologie d e la r6gion d e Bakouma. Chron. Mines Rech. Miniere, 370: 43-45. Binda, P.L., 1972. Zircons of the Nchanga granite and overlying metasediments, Zambia. 24th Int. Geol. Congr., Section 1,Precambrian, pp. 179-186. Binda, P.L. and Van Eden, J.G., 1972. Sedimentological evidence on the origin of the Precambrian Great Conglomerate (Kundelungu Tillite), Zambia. Palaeogeogr., Palaeoclimatol., Palaeoecol., 1 2 : 151-168. Bird, H.H., 1968. Falconbridge’s copper operations in Uganda. Can Min. Metal. Bull., Sept. 1968. Cahen, L. and Lepersonne, J., 1967. The Precambrian of the Congo, Rwanda, and Burundi. In: (K. Rankama (Editor), The Precambrian, Vol. 3. Interscience, New York, N.Y., pp. 143-290. Cahen, L. and Snelling, N.J., 1966. The Geochronology of Equatorial Africa. NorthHolland, Amsterdam, 1 6 1 pp. Clifford, T.N., 1966. Tectono-metallogenic units and metallogenic provinces of Africa. Earth Planet. sci. Lett., l ( 6 ) : 421-434. Clifford, T.N., 1967. The Damaran Episode in the Upper Proterozoic Lower Paleozoic structural history of Southern Africa. Geol. Soc. Am. Spec. Pap. 92. Clifford, T.N. and Gass, I.G. (Editors), 1970. African Magmatism and Tectonics. Oliver and Boyd, Edinburgh, 461 pp. De Swardt, A.M.J. and Drysdall, A.R., 1964. Precambrian geology and structure in central Northern Rhodesia. Geol. Surv. Mem., 2: 82 pp. De Wetter, D.R., Dijkstra, S. and Ter Bruggen, J.W., 1972. Zambia Issue. Geol. Mijnbouw, 51(3): 247-468. Des Ligneris, X. and Bernazeaud, J., 1959. Le Gisement de Mounana (Gabon). Science et Technique, pp. 1-13.

797 Dechow, E., Thomson, R. and Andreazza, P., 1976. An African model for base metal exploration in Brazil. Presented 29th Congress0 Sociedade Brasiliero d e Geologia Belo Horizonte, Nov. 1976. Dhvigne, J.P., 1959. Le Pr6cambrien du Gabon Occidental en Afrique Equatoriale FranGaise et regions limitrophes. Direction des Mines et d e la Geologie, Afrique Equat. Franqaise, Bull. 1 1 , 3 1 5 pp. Ellis, M.W. and McGregor, J.A., 1967. The Kalengwa copper deposits in Northwestern Zambia. Econ. Geol., 62(6): 781-797. Fleischer, V.D., 1967. Relation between folding, mineralization and Sub-Katanga topography a t Mufulira Mine, Zambia Trans. Geol. SOC.S. Afr., 70: 1-14. Fyfe, W.S. and Leonardos, O.H., Jr., 1973. Ancient metamorphic-migmatite belts of the Brazilian African Coasts. Nature, 244 : 501-502. Gerard, G., 1958. Carte Geologique de 1’Afrique Equatoriale Francaise, au 1/2,000,000. Notice explicative. Direction des Mines et de la Geologie, Afrique Equat. Francaise. Haughton, S.H., 1961. Review of a probable late Precambrian glacial period in Central and West Africa. Trans. Geol. SOC.S. Afr., 64: 73-84. Haughton, S.H., 1963. The Stratigraphic History of Africa South of the Sahara. Hafner, New York, N.Y., 3 6 5 pp. Hurley, P.M., 1968. The confirmation of continental drift. Sci. Am., 218(4): 5 3 4 3 . Hurley, P.M., 1974. Pangeaic orogenic system. Geology, 2(8): 373-376. Hurley, P.M. and Rand, J.R., 1969. Pre-drift continental nuclei. Science, 164(3885): 3 2 29-1 242. Isotta, C.A.L., Rocha-Campos, A.C. and Yoshida, R., 1969. Striated pavement of the upper Pre-Cambrian glaciation in Brazil. Nature, 222 : 466-468. Kennedy, W.Q., 1964. The structural differentiation of Africa in the Pan-African ( t 5 0 0 m.y.) tectonic episode. Univ. Leeds. Res. Inst. Afr. Geol., Annu. Rep., 8: 4849. Lombard, J. and Nicolini, P., 1962. Stratiform copper deposits in Africa. Symposium, Pt. 1: Lithology Sedimentology. Assoc. African Geol. Surveys (ASGA), Paris, 212 pp. Lombard, J. and Nicolini, P., 1963. Stratiform copper deposits in Africa. Symposium, Pt. 11: Tectonics. Assoc. African Geol. Surveys (ASGA), Paris, 265 pp. Lister, L.A. (Editor), 1973. Symposium on granites gneisses and related rocks. Geol. SOC. S. Afr., Spec. Publ. No. 3 , 509 pp. Mendelsohn, F. (Editor), 1961. The Geology of the Northern Rhodesian Copperbelt. MacDonald, London, 523 pp. Mendes, F. and Vialette, Y., 1972. Le Precambrian d e 1’Angola. 24th Int. Geol. Congr. section 1 , Precambrian, pp. 213-220. Mouta, F., 1954. Noticia explicativa d o Esboco Geolbgico de Angola, Ministerio d o Ultramar, Lisbon, 1 4 8 pp. Pelletier, R.A., 1964. Mineral Resources of South-Central Africa. Oxford Univ. Press, Capetown; 277 pp. Sawkins, F.J., 1976. Widespread continental rifting: Some considerations of timing and mechanism. Geology, 4(7): 427-430. Stanton, W.I., 1963. The geology of degree sheet Sul B-33/JP (Maquela d o ZombeDamba). Serv. Geol. Minas, Angola, Bol., 8: 79-136. Stanton, W.I., Schermerhorn, L.J.G. and Korpershoek, H.R., 1963. The West Congo System. Serv. Geol. Minas, Angola, Bol., 8: 69-78. Van Doorninck, N.H., 1928. De Lufilische Plooing. G. Naeff, The Hague, 202 pp. Vail, J.R., Snelling, N.J. and Rex, D.C., 1968. Pre-Katangan geochronology of Zambia and adjacent parts of Central Africa. Can. J. Earth Sci., 5(3): 621-4328. Walter, M.R. (Editor), 1976. Stromatolites. Elsevier, Amsterdam, 790 pp. Wolf, K.H. (Editor), 1976. Strata-bound and Stratiform Ore Deposits. Elsevier, Amsterdam, 6, 585 pp.

798

Part F Broderick, T.J., 1976. Explanation of the geological map of the country east of Kariba. Rhod. Geol. Surv. Short Rep. 43. Carta Geologica, Provincia de Mocambique, 1 : 250,000, Folha SUL-E-36/J. Lourenco Marques 1968. Clifford, T.N., Rex, D.C. and Snelling, N.J., 1967. Radiometric data for the Urungwe and Miami granites of Rhodesia. Earth Planet. Sci. Lett., 2: 5-12. Johnson, R.L. and Vail, J.R., 1965. The junction between the Moqambique and Zambezi orogenic belts; north-east Southern Rhodesia. Geol. Mag., 102: 489-495. Kirkpatrick, I.M., 1976. The geology of the country around Tengwe, Lomagundi District. Rhod. Geol. Surv. Bull. 75. Leitner, E.G. and Phaup, A.E., 1974. The geology of the country around Mount Darwin. Rhod. Geol. Surv. Bull. 73. Loney, P.E., 1969. The Geology of the Kariba District, Rhodesia, with special Reference to Geochronology and amphibolite Petrochemistry. Ph.D. Thesis, Leeds Univ. (unpubl.). Stagman, J.G., 1962. The geology of the southern Urungwe District. Rhod. Geol. Surv. Bull. 55. Talbot, C.J., 1973. The geology of the eastern end of the Zambezi Orogenic Belt, Rhodesia. Trans. Geol. SOC.S. Afr., 76: 113-133. Thole, R.H., 1974. The Geology of and Controls to the Distribution of Copper at the Shamrocke Mine, near Karoi, Rhodesia. Ph.D. Thesis, Univ. Rhod. (unpubl.). Wiles, J.W., 1968. Some aspects of the metamorphism of the Basement Complex in the Sipolilo District. Trans. Geol. SOC.S. Afr., Annex., 71: 79-88. Workman, D.R., 1961. The Geology of Mhavare, north-west Lomagundi District, Southern Rhodesia, with special Reference to the Stratigraphy and Metamorphism of the Lomagundi System. Ph.D. Thesis, Leeds Univ. (unpubl.).

Part G Bloomfield, K., 1968. The pre-Karoo geology of Malawi. Mem. Geol. Surv. Malawi, 5: 166 pp. Cannon, R.T., Hopkins, D.A., Thatcher, E.C., Peters, E.R., Kemp, J., Gaskell, J.L. and Ray, G.E., 1969. Polyphase deformation in the Mozambique Belt, northern Malawi. Geol. SOC.Am. Bull., 80: 2615-2622. Carter, G.S. and Bennett, J.D., 1973. The geology and mineral resources of Malawi. Bull. Geol. Surv. Malawi, 6 (2nd revised ed.): 6 2 pp. Clifford, T.N., 1970. The structural framework of Africa. In: T.N. Clifford and I.G. Gass (Editors), African Magmatism and Tectonics, Oliver and Boyd, Edingburgh, pp. 1-26. Clifford, T.N., 1974. Review of African granulites and related rocks. Geol. SOC.Am. Spec. Pap., 156: 49 pp. Dodson, M.H., Cavanagh, B.J., Thatcher, E.C. and Aftalion, M., 1975. Age limits for the 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., 1972. Outline of the geology of Zambia. Geol. Mijnbouw, 51 : 265-276. Fitches, W.R., 1971. Sedimentation and tectonics a t the northeast end of the Irumide Orogenic Belt, N. Malawi and Zambia. Geol. Rundsch., 6 0 : 5 8 9 - 6 1 9 . Kemp, J., 1973. The Precambrian Geology of two Areas in Malawi: its Significance for the Mozambiquian and Ubendian Belts. Ph.D. Thesis, Univ. Leeds (unpubl.). Kennedy, W.Q., 1964. The structural differentiation of Africa in the Pan-African (*500 million years) tectonic episode. Univ. Leeds Res. Inst. Afr. Geol. Annu. Rep. 8 (1962-1963): 48-49.

799 Klinck, B.A., 1980. The geology of the Chisomo area. Rep. Geol. Surv. Zambia (in press). Ramsay, C.R. and Ridgway, J., 1976. Metamorphic patterns in Zambia and their bearing on problems of Zambian tectonic history. Occ. Papers Geol. Surv. Zambia: 76. Ray, G.E., 1974. The structural and metamorphic history of northern Malawi. J. Geol. SOC.London, 130: 427-440. Snelling, N.J., Johnson, R.L. and Drysdall, A.R., 1972. The geochronology of Zambia. Rec. Geol. surv. Zambia, 1 2 : 19-30. Thatcher, E.C., 1973a. Structural and Stratigraphic Relationships in the Precambrian Rocks of the Nyika Area, northern Malawi. Ph.D. Thesis, Univ. Leeds (unpubl.). Thatcher, E.C., 1973b. Gneisses and granites in the Malawi Province of the Mozambique Belt. In: L.A. Lister (Editor), Symposium on granites, gneisses and related rocks. Geol. SOC.S. Africa, Spec. Pub]., 3: 149-155. Vail, J.R., 1965. An outline of the geochronology of the late Precambrian formations of eastern central Africa. Proc. R. SOC.London, Ser. A, 284: 354-369. Vavrda, I. and Vrana, S., 1972. Sillimanite, kyanite and andalusite in the granulite facies rocks of the Basement Complex, Chipata District. Rec. Geol. Surv. Zambia, 1 2 : 6980.

Part H Ahrendt, H., Hunziker, J.C. and Weber, K., 1978. Age and degree of metamorphism and time of nappe emplacement along the southern margin of the Damara Orogen/Namibia (S.W. Africa). Geol. Rundsch., 6 7 : 719-742. 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. Berning, J., Cooke, R., Hiemstra, S.A. and Hoffmann, U., 1976. The Rossing Uranium deposit, South West Africa. Econ. Geol., 71: 351-368. Blaine, J.L., 1978. Tectonic evolution of the Waldan Ridge Structure and the Okahandja Lineament in part of the central Damara orogen, west of Okahandja, South West Africa. Precambrian Res. Unit, Univ. Cape Town, Bull. 21. Bunting, I.J.L., 1977. Geology of Part of the Central Damara Belt around the Tumas River, South West Africa. M.Sc. Thesis, Rhodes Univ. Burger, A.J. and Coertze, F.J., 1973. Radiometric age measurements on rocks from Southern Africa t o end of 1971. Geol. Surv. S. Afr. Bull., 58. Burger, A.J. and Coertze, F.J., 1975. Age determinations - April 1972 to March 1974. Geol. Surv. S. Afr. Ann., 10: 135-141. Burger, A.J., Clifford, T.N. and Miller, R.McG., 1976. Zircon U-Pb ages of the Franzfontein granitic suite, northern South West Africa. Precambrian Res., 3, p. 415-431. Cahen, L. and Snelling, N.J., 1966. The Geochronology of Equatorial Africa. NorthHolland, Amsterdam, 1 9 5 pp. Clifford, T.N., 1962. Note on nappes in the Otavi Facies of northern South West Africa. Univ. Leeds Res. Inst. Afr. Geol. Annu. Rep., 6 : 44. Clifford, T.N., 1967. The Damaran episode in the Upper Proterozoic-Lower Paleozoic structural history of southern Africa. Geol. SOC.Am. Spec. Pap. 92. Clifford, T.N., Nicolaysen, L.O. and Burger, A.J., 1962. Petrology and age of the PreOtavi Basement Granite a t Franzfontein, northern South West Africa. J. Petrol., 3 : 244-278. Clifford, T.N., Rooke, J.M. and Allsopp, H.L., 1969. Petrochemistry and age of the Franzfontein granitic rocks of northern South West Africa. Geochim. Cosmochim. Acta, 33: 973-986. Cochran, P., Comline, S.R., Ingram, R., Hoffman, K. and Page-Green, B.D., 1975. Various internal reports and results of field mapping. Companhia Mineira d o Cunene S.A.R.L. Luanda, Angola.

800 Cole, M.M. and le Roux, H.D., 1978. The role of geobotany, biogeochemistry and geochemistry in mineral exploration in South West Africa and Botswana - a case history. Trans. Geol. SOC.S. Afr., 8 1 : 277-318, De Kock, W.P., 1934. The Geology of the Western Rehoboth, South West Africa. Geol. Sum. S. Afr. Mem. 1. De Villiers, J. and Simpson, E.S.W., 1974. Late Precambrian tectonic patterns in SouthWestern Africa. Precambrian Res. Unit, Univ. Cape Town, Bull., 15: 141-152. De Waal, S.A., 1966. The Alberta Complex, a Metamorphosed Layered Intrusion, North of Nauchas, South West Africa, the surrounding Granites and Repeated Folding in the Damara System. D.Sc. Thesis, Univ. Pretoria. Eskola, P., 1949. The problem of mantled gneiss domes. Q.J. Geol. SOC.London, 104: 46 1-4 76. Finnemore, S.H., 1978. The geochemistry and origin of the Matchless Amphibolite Belt, Windhoek District, South West Africa. In: Mineralization in Metamorphic Terranes. Geol. SOC.S. Afr., Spec. Publ., 4 : 433-447. Frets, D.C., 1969. Geology and structure of the Huab-Welwitschia area, South West Africa. Precambrian Res. Unit, Univ. Cape Town, Bull. 5. Germs, G.J.B., 1972. The stratigraphy and palaeontology of the lower Nama Group, South West Africa. Precambrian Res. Unit, Univ. Cape Town, Bull. 12. Germs, G.J.B., 1974. The Nama Group in South West Africa and its relationship to the Pan-African geosyncline. J. Geol., 82: 301-317. Gevers, T.W., 1931. The Fundamental Complex of Western Damaraland, South West Africa. D.Sc. Thesis, Univ. Cape Town. Gevers, T.W., 1934. The geology of the Windhoek District in South West Africa. Trans. Geol. SOC.S. Afr., 39: 221-251. Gevers, T.W., 1963. Geology along the northwestern margin of the Khomas Highlands between Otjimbingwe-Karibib and Okahandja, South West Africa. Trans. Geol. SOC. S. Afr., 66: 199-258. Goldberg, I., 1976. A preliminary account of the Otjihase copper deposit, South West Africa. Econ. Geol., 71: 384-390. Guj, P., 1970a. The Damara mobile belt in the south-western Kaokoveld, South West Africa. Precambrian Res. Unit, Univ. Cape Town, Bull. 8. Guj, P., 1970b. The relationship between the “Franzfontein Granite” and the Huab and Khoabendus Formations northwest of Franzfontein, South West Africa. Geol. Surv. S. Afr. Ann., 8: 45-51. Halbich, I.W., 1970. The Geology of the Western Windhoek and Rehoboth Districts: a Stratigraphic-Structural Analysis of the Damara System. D.Sc. Thesis, Univ. Stellenbosch. Hartnady, C.J., 1978a. The stratigraphy and structure of the Naukluft nappe complex. Precambrian Res. Unit, Annu. Rep., 14/15: 163-170. Hartnady, C.J., 1978b. Tectonic evolution of the southwestern Hakos-Auas Mountains zone in the Damara Orogenic Belt. Precambrian Res. Unit Annu. Rep. 14/15: 171182. Hedberg, R.M., 1979. Stratigraphy of the Ovamboland Basin, South West Africa. Precambrian Res. Unit, Univ. Cape Town, Bull. 24. Hugo, P.J. and Schalk, K.E.L., 1975. The isotopic ages of certain granites and acid lavas in the Rehoboth and Maltahohe districts, South West Africa. Geol. Surv. S. Afr. Ann., 9: 103-105. Jackson, M.P.A., 1976. Highgrade metamorphism and migmatisation of the Namaqua Metamorphic Complex around Aus in the southern Namibia Desert, South West Africa. Precambr. Res. Unit, Univ. Cape Town, Bull. 18.

801 Jacob, R.E., 1974. Geology and metamorphic petrology of part of the Damara orogen along the lower Swakop River, South West Africa. Precambrian Unit, Univ. Cape Town, Bull. 17. Jacob, R.E., 1978. Granite genesis and associated mineralisation in part of the Central Damara Belt. In: Mineralization in Metamorphic Terranes. Geol. Soc. S. Afr. Spec. Publ., 4 : 417-432. Jacob, R.E., Kroner, A. and Burger, A.J., 1978. Areal extend and first U-Pb age of the Pre-Damaran Abbabis Complex in the Central Damara Belt South West Africa (Namibia). Geol. Rundsch., 67 : 706-718. Jeppe, J.F.B., 1952. The Geology of the Area along the Ugab River west of the Brandberg. Ph.D. Thesis, Univ. Witwatersrand, Johannesburg. Kasch, K.W., 1979. A continental collision model for the tectono-thermal evolution of the (southern) Damara Belt. Precambrian Res. Unit, Univ. Cape Town, Annu. Rep., 1 6 : 101-107. Kennedy, W.Q., 1964. The structural differentiation of Africa on the Pan-African (k500 m.y.) tectonic episode. Univ. Leeds Res. Inst. Afr. Geol. Annu. Rep. 1962/63, 8: 48-49. Korn, H. and Martin, H., 1959. Gravity Tectonics in the Naukluft Mountains, S.W.A. Bull. Geol. SOC.Am., 70: 1047-1078. Kroner, A., 1971. Late Precambrian correlation and the relationship between the Damara and the Nama Systems of South West Africa. Geol. Rundsch., 60: 1513-1523. Kroner, A., 1975. Geochronology. Precambrian Res. Unit, Univ. Cape Town, Annu. Rep., 1 3 : 139-143. Kroner, A., 1977. Precambrian mobile belts of southern and eastern Africa - ancient sutures or sites of ensialic mobility? A case for crustal evolution towards plate tectonics. Tectonophysics, 40: 101-135. Kroner, A. and Hawkesworth, C.J., 1977. Late Pan-African emplacement ages for Rossing alaskitic granite (Damara Belt) and Rooi Lepel bostonite (Gariep Belt) in Namibia, and their significance for the timing of metamorphic events. Univ. Leeds Res. Inst. Afr. Geol. Annu. Rep. 1976, 20: 16-20. Kroner, A.,Halpern, M. and Jacob, R.E., 1978. Rb-Sr geochronology in favour of polymetamorphism in the Pan African Damara Belt of Namibia (South West Africa). Geol. Rundsch., 67: 688-705. Lee, J.E. and Glenister, D.A., 1976. Stratiform sulphide mineralisation at Oamites Copper Mine, South West Africa. Econ. Geol., 71: 369-383. Malling, S.,1975. Preliminary report on the investigation of the pre-Damaran rocks of the Nauchas area, South West Africa. Precambrian Res. Unit, Univ. Cape Town, Annu. Rep., 13: 97-101. Martin, H., 1965. The Precambrian Geology of South West Africa and Namaqualand. Precambrian Res. Unit, Univ. Cape Town, Rustica Press, 159 pp. Martin, H., 1978. The mineralisation of the ensialic Damara Orogenic Belt. In: Mineralisation in Metamorphic Terranes. Geol. SOC.S. Afr. Spec. Publ., 4 : 405-415. Martin, H. and Porada, H., 1977a. The intracratonic branch of the Damara Orogen in South West Africa - discussion of geodynamic models. Precambrian Res., 5: 311338. Martin, H. and Porada, H., 1977b. The intracratonic branch of the Damara Orogen in South West Africa - discussion of relationships with the Pan-African mobile belt system. Precambrian Res., 5: 339-358. Miller, R.McG., 1972. The Geology of a Portion of southern Damaraland, South West Africa, with Particular Reference to the Petrogenesis of the Salem Granite. Ph.D. Thesis, Univ. Cape Town.

Miller, R.McG., 1974. The stratigraphic significance of the Naauwpoort Formation of east central Damaraland. Trans. Geol. SOC.S. Afr., 77: 363-367. Nash, C.R., 1971. Metamorphic petrology of the S.J. area, Swakopmund district, South West Africa. Precambrian Res. Unit, Univ. Cape Town, Bull. 9. Porada, H., 1974. The Khoabendus Formation in the area northwest of Kamanjab and in the southeastern Kaokoveld, South West Africa. Geol. Surv. S. Afr. Bull. 4, S.W.A. series. Roper, H., 1959. The Geology of the Otjosondu Manganese Area, South West Africa. Ph.D. Thesis, Univ. Witwatersrand, Johannesburg. Sawyer, E.W., 1978. Structure and Metamorphic Geology of the Damara in an Area Southeast of Walvis Bay. M.Sc. Thesis, Univ. Cape Town. Schalk, K.E.L., 1973. Some Late Precambrian formations in Central South West Africa. Geol. Suw. S. Afr. Ann., 8: 29-47. Smith, D.A.M., 1961. The Geology of the Area around the Khan and Swakop Rivers in South West Africa. Ph.D. Thesis, Univ. Witwatersrand, Johannesburg. Smith, D.A.M., 1965. The geology of the area around the Khan and Swakop Rivers in South West Africa. Geol. Surv. S. Afr., Mem. 3. Sohnge, P.G., 1964. Geology of the Tsumeb Mine in The Geology of some Ore Deposits in Southern Africa. Geol. SOC.S. Afr. Spec. Publ., 1: 367-382. Sohnge, P.G., 1974. Sedimentary Ore Deposits: A review of Research Trends. Trans. Geol. SOC.S. Afr., 77 : 159-168. Sutton, J., 1967. The extension of the geological record into the Precambrian. Proc. Geol. ASSOC.,78: 493-534. Thomas, C.N., 1969a. A brief description of the geology of south Ngamiland. Rep. Geol. Suw. Botswana (unpubl.). Thomas, C.N., 1969b. Geological notes on the Chobe District. Rep. Geol. Surv. Botswana (unpu bl. ). Toens, P.D., 1975. The geology of part of the southern foreland of the Damara orogenic belt in South West Africa and Botswana. Geol. Rundsch., 6 4 : 175-192. Von Backstrom, J.W. and Jacob, R.E., 1979. Uranium in South Africa and South West Africa (Namibia). Philos. Trans. R. SOC. London, Ser. A, 291, 307-319. Von Brunn, V., 1967. Acid and Basic Igneous Rock Associations West of Helmeringhausen, South West Africa. Ph.D. Thesis. Univ. Cape Town. Watters, B.R., 1974. Stratigraphy, igneous petrology and evolution of the Sinclair Group in southern South West Africa. Precambrian Res. Unit, Univ. Cape Town, Bull. 16. Watters, B.R., 1977. The Sinclair Group: Definition and Regional Correlation. Trans. Geol. SOC.S. Afr., 80: 9-16. Watters, B.R., 1978. Petrogenesis of the felsic rock units of the late-Precambrian Sinclair Group, South West Africa. Geol. Rundsch., 6 7 : 743-773. Windley, B.F., 1977. The Evolving Continents. Wiley, New York, N.Y., 385 pp.

80 3

Chapter 11

CARBONATITES, ALKALINE AND RELATED INTRUSIONS D.R. HUNTER

INTRODUCTION

Carbonatites and related alkaline intrusions of Precambrian age are most extensively developed within the Kaapvaal craton or structural province of the Republic of South Africa. At least two ages of emplacement have been recognized. The older, represented by the Palabora Complex and a number of small peralkaline intrusions in the collar of the Vredefort dome, were emplaced at -2000 Ma. The Pilanesberg alkaline province, dated at 1300 Ma, extends for a distance of at least 300 km in a north-northwesterly direction in the western Transvaal, and comprises dyke swarms, central intrusions, pipes and irregular plugs.

-

THE

- 2000 Ma INTRUSIONS

The Palabora Complex in the northeastern Transvaal is unique among other African alkaline complexes in that its carbonatite member contains economically viable concentrations of copper. It consists of a main composite intrusion with numerous dykes and separate plugs of syenite and carbonate-bearing injection breccia extending along an approximately east-west zone, 80 km long and 20 km wide. The main composite plug is elongated in a n o r t h s o u t h direction 6.5 km long and 2.5 km wide, underlying an area of about 1 7 km2. (Fig. 11.1).The first intrusive stage is represented by a vertical plug of pyroxenite, consisting mainly of diopside, phlogopite and apatite, which is surrounded by a nearly continuous outer rim about 100 m across composed of feldspathic pyroxenite. Interaction between the pyroxenite and the various Archaean granitic gneisses forming the wall of the plug is believed to be responsible for the development of a feldspathic outer rim. Concentrations of apatite are sufficiently high in parts of the pyroxenite to warrant economic exploitation which is currently being undertaken by the Phosphate Development Corporation. Beyond the limits of the pyroxenite intrusion the second intrusive stage is marked by the emplacement of numerous syenite plugs. Where syenite occurs close to the pyroxenite contact, the feldspathic nature of the

804

Fig. 11.1. Simplified geological map of the Palabora Complex.

marginal pyroxenite is prominent. Strong brecciation around the outer contacts of, and occasionally throughout the syenite testifies to explosive emplacement. The granitic gneisses in contact with syenite and pyroxenite are locally fenitized, mainly as a result of potassium-sodium-iron metasomatism during the pneumatolytic phases of the intrusive stages. The syenites vary in grain size and texture from dense, fine-grained to porphyritic and pegmatitic types. Three pegmatoid bodies are located within the pyroxenite. The northern pegmatoid has an oval shape in plan and consists of a core of phlogopite-serpentine rock with minor diopside surrounded by an outer zone composed of

805 medium- to fine-grained phlogopite--diopside rock which contains economic concentrations of apatite. Alteration of the phlogopite to vermiculite and hydrophlogopite is common in the weathered zone. The host t o the copper mineralization is a composite pipe of carbonatite enclosed by an outer rim of magnetite-olivine-apatite rock located slightly to the west of the centre of the pyroxenite plug within the third pegmatoid body. The vertical pipe has an elliptical shape in plan, elongated in an eastrock (locally termed foswest direction. The magnetite-olivine-apatite korite) in the outer part of the pipe consists of a series of concentric layers interbanded with the pyroxenite host. The boundaries between foskorite and both the pyroxenite and the banded carbonatite is either sharp or gradational. Towards the centre of the pipe the foskorite and carbonatite are interbanded in a similar fashion, both rock types s h o ~ i n ga near vertical mineral banding due t o the alignment of magnetite concentrations. Intense fracturing trending WNW and ENE preceded the emplacement of a second period of carbonatitic magmatism, which is concentrated at the intersection of these two fracture directions. This late-stage carbonatite forms a dykelike body that transects the concentric layering of the foskorite and older carbonatite, which are also cut by a network of closely spaced, impersistent veinlets of the transgressive carbonatite. Magnetite constitutes between 25 and 50 weight percent and apatite 1 2 t o 25 percent of the foskorite, the olivine in which is serpentinized to depths of 1000 m below surface. The banded carbonatite is composed of magnesian calcite, magnetite, and subordinate apatite with chondrodite, olivine, phlogopite and biotite as accessory minerals. The late-stage carbonatite has a similar mineralogy with the calcite being more magnesian but lacks the mineral banding. The main body of this late carbonatite has a very variable grain size ranging from fine- to very coarse-grained in contrast to the more mediumto coarse-grained size of the banded carbonatite. The carbonatite is essentially magnetite-rich sovite with 3 to 8%MgO. Its most unusual feature is the abundance of sulphide minerals which are found in fine discontinuous fractures in zones up to 10 m wide. Mineralization also accompanies the carbonatite veinlets in the older rock types. Sulphide blebs are found in the banded carbonatite and foskorite unrelated to fracturing and it has been concluded that these older rocks suffered an earlier phase of mineralization. The principal copper sulphides are chalcopyrite and bornite. Valeriite, a very latestage copper sulphide, is mainly concentrated in broad shear zones and replaces all other sulphides as well as magnetite, carbonate and silicate gangue. Pyrrhotite, pentlandite, millerite bravoite, tetrahedrite, sphalerite, galena, and pyrite are also present. In 1973, tons of ore mined amounted to 19184,961 at an average grade of 0.54% copper. The only age determination made on the Palabora Complex dates back to 1957 when a date of 2060 f 100 Ma was given for thorianite from the complex (Holmes and Cahen, 1957). Palaeomagnetic study of some of the post-

806

27'30'

0

Mesozoic cover rocks Alkaline intrusions

Basic intrusions

H :::: :

TR A NSVA A L Ventersdorp Witwatersrand Granite

Fig. 11.2. Locations of alkaline intrusions around the Vredefort dome, south of Johannesburg. Note the circle encloses the area in which alkali granite/aplite dykes are known.

complex dolerite dykes has indicated that they may be about 1880 Ma-old (J.C. Briden, 1974, pers. commun. to Palabora Mining Company). If the age of 2060 Ma is confirmed by more detailed studies it would suggest that the emplacement of the Palabora Complex is nearly contemporaneous with the Bushveld Complex. Minor peralkaline intrusions are located around the periphery of the Vredefort Dome (Bisschoff, 1973). Geochronological, palaeomagnetic and field data point to: (1)a possible genetic relationship between these intrusions and the small Roodekraal and Lindesquesdrift mafic complexes; (2) emplacement at a late stage in the evolution of the Vredefort Dome; and

807 (3) contemporaneity with the late stages of magma activity in the Bushveld Complex. The three alkali granite plutons in the collar of the Vredefort Dome were emplaced along strike faults at Rietfontein, Schunvedraai and Baviaan Kranz (Fig. 11.2). Dykes of aplitic alkali granite and alkali syenite are found mainly in the eastern part of the Vredefort dome, and have been previously described under a variety of names (Rogers, 1922; Nel and Jansen, 1957). The alkali granite is a sodic variety, composed on average of 80% feldspar, 10% quartz and 8% alkali amphibole (riebeckite or arfvedsonite) and/or alkali pyroxene (aegirine). Quartz-poor syenitic phases have a patchy distribution in the finer-grained alkali granite into which they are gradational. Bisschoff (1973) considers that the alkali granite originated from an alkali syenitic magma by assimilation of quartz-rich rocks. The nepheline syenite dykes, previously termed canadite and litchfieldite (Hall and Molengraaff, 1925), are classified as mariupolites by Bisschoff (1973) because of their paralkaline characteristics. The dykes are intrusive mainly into the roof and hood zones of the southeastern side of the Schunvedraai alkali granite pluton. Bisschoff (1973) notes that mariupolite dykes are absent from those parts of the pluton that have been deeply eroded and suggests that this mode of occurrence points to the dykes being rooted in the alkali granite. The dykes consist of fine- to medium-grained and coarse-grained, pegmatitic portions, the latter forming narrow bands or irregular elongated patches parallel to the strike of the dykes less than a centimetre to 40 cm in width. The mariupolite dykes are composed of albite, nepheline, aegirine and microcline, lepidomelane being present occasionally. Bisschoff (1973) proposed that the mariupolite dykes originated by desilication of alkali syenite liquids under conditions of high PHZO. The close spatial relationship between the dykes and quartz syenite is considered t o be indicative of a possible genetic connection between the rock types. Bisschoff (1973) also regarded the alkali granite plutons as crystallization products of a magma evolved from the original quartz syenitic liquid by assimilation. Minor syenite intrusions are common in southeastern Botswana. Samples from the vicinity of Letilapana Pan have yielded whole-rock Rb/Sr dates of around 1900 Ma and K/Ar hornblende ages of 1805 t o 1908 Ma, which are regarded as minima (Key, 1977). If these dates are confirmed it suggests the presence of syenitic, dioritic, gabbroic and noritic rocks comparable in age t o the Bushveld Complex. THE

- 1300 Ma INTRUSIONS

The Pilanesberg alkaline province in the western area of the Transvaal represents a major igneous event. The province consists of swarms of narrow dykes of subalkaline and femic character which pre-date passively emplaced intrusions that are exclusively salic and alkaline whereas abundant kimber-

lite and carbonatite occurrences demonstrate forceful intrusions (Ferguson, 1973). The Leeuwfontein alkaline complex has been dated at 1420 70 Ma (Oosthuysen and Burger, 1964). A K/Ar biotite age of 1250 Ma has been obtained from the Pilanesberg complex (Snelling, 1963). Pilanesberg dykes have been dated at 1290 ? 180 Ma and 1330 2 80 Ma (Schreiner and Van Niekerk, 1958; Van Niekerk, 1962). The post-Waterberg kimberlite pipe at the Premier mine is cut by a mafic dyke dated at 1115 k 15 Ma (Allsopp et al., 1967). The main sites of Pilanesberg igneous activity are located within a 200 kmwide belt extending from beneath the Mesozoic cover south of Johannesburg for at least 300 km in a northwesterly direction into eastern Botswana. Ferguson (1973) interpreted this distribution as a reflection of the structural control exerted by a northwesterly trending anticlinal flexure the trace of which passes through the three domal structures, namely the Johannesburg, Crocodile and Makoppa domes. To the east of this structure lies a basin filled by thick accumulation of Waterberg sediments. There is no evidence for the existence of a well-defined basin to the west during Waterberg times, although other second order flexures trending both northwest and northeast were probably present. The dominant site of emplacement of the Pilanesberg dykes and central intrusions is considered to be a monoclinal structure. Ferguson (1973) noted that support for this contention is provided by aeromagnetic data which indicate that the main concentration of dyking between the Pilanesberg central intrusions and the Botswana border is limited t o a 40 km-wide zone trending N40”W. Verwoerd (1967) pointed out that, with the exception of the most northerly intrusion (Glenover) in the Transvaal, all the remaining carbonatite intrusions lie along an east-west belt and that the main concentration of these intrusions near Pretoria is coincident with the intersection of this belt with a southward projection of the Great “Dyke” lineament. Although the Great “Dyke” is of greater antiquity than the Pilanesberg intrusions, the possibility of re-activation of this lineament is suggested by the age of the Trompsburg Complex at the southern end of the lineament which is dated at -1372 Ma (Davies et al., 1969). There is a general absence of rhythmic layering in the Pilanesberg province which Ferguson (1973) considers to be probably a reflection of the volcanic to sub-volcanic nature of the central intrusions. The non-composite, northwesterly trending dykes generally have tholeiitic basalt or tholeiitic andesite composition (Ferguson, 1973). The larger dykes exceeding 40 m in width are frequently composite having margins that are more femic than the centres. Ferguson (1973) estimated that the proportion of femic to salic units in these dykes is of the order of 20 : 80 but femic borders can make up to 40% by volume. The composition of the femic margins is similar to that of the non-composite dykes, grading into potassic tristanite or hawaiite. In two dykes hawaiite or nepheline-bearing trachyte constitute the margins grading in both cases t o centres composed of nepheline syenite.

*

809 Dykes less than a metre wide, frequently randomly oriented, are found on the west side of the Pilanesberg central intrusion. The dykes are lamprophyric, carbonatitic or, less frequently, kimberlitic in composition. The lamprophyres vary considerably in composition and have suffered late-stage carbonatization. Camptonite containing kaersutite, titanaugite, olivine and ocelli showing carbonate and silicate phases (Ferguson, 1973) is the more general variety. Verwoerd (1967) noted the similarity in the chemistry of some lamprophyres t o alnoite. Minor occurrences of spessatite, melteigite, jacupirangite and urtite have also been reported (Toens, 1952; Verwoerd, 1967). Narrow carbonatite intrusions, either sovite or beforsite, are found throughout the Pilanesberg province, particularly adjacent to carbonatite plugs where they occur as radial and ring-dykes or infrequently as conesheets (Ferguson, 1973). The central intrusions have compositions equivalent to those displayed by the hypabyssal and volcanic rocks with the exception of tholeiitic basalt and andesite, hawaiite and lamprophyres (Ferguson, 1973). These central complexes usually post-date the northwesterly trending dyke swarms and are exclusively salic. Within individual complexes the later rock types are generally more alkaline than the earlier ones. The Pilanesberg central complex differs in that later alkaline rocks include both over- and undersaturated varieties (Ferguson, 1973). Ferguson (1973) has estimated the areal extent of rock types in the Pilanesberg province (Table ll.I), and shows that nepheline syenites and

TABLE 11.1 Areal extent of rock types in the Pilanesberg alkaline province Rock type ~~

Basaltic Tholeiitic andesite Hawaiite Nepheline-bearing trachyte Miaskitic nepheline syenite Tristanite (potassic) Tristanite Sodic rhyolite Agpaitic nepheline syenite Lamprophyre Kimberlite Carbonatite (sovite) Carbonatite (beforsite) Total

Area ( km2 )

%’

~~

21.80 23.41 3.62 12.04 91.44 76.96 28.60 2.81 42.90 1.45 0.40 2.24 16.85 324.52

6.72 7.21 1.12 3.71 28.18 23.72 8.81 0.87 13.22 0.45 0.12 0.69 5.19

810 potassic tristanites have by far the greatest areal extent being concentrated in the central intrusions at Pilanesberg, Roodeplaat and Leeuwfontein. Agpaitic nepheline syenites, that occur exclusively in the Pilanesberg intrusion are the next most abundant. Beforsite is the dominant type among the ultramafic and COz-rich rocks. Ferguson (1973) notes a similarity between the Pilanesberg province and those of Gardar and Oslo in the dominance of nepheline syenite and the small amount of femic material, almost exclusively confined t o narrow dykes and effusives, which is a marked feature of continental alkaline provinces. Ferguson (1973) concludes that the parallelism of the faulting, shear and dyke directions suggests emplacement of the femic magmas during tensional conditions, but periods of compression must have alternated with tension as indicated by shearing along some dyke margins. The frequent presence of partly digested mafic inclusions in the more salic centres of composite dykes implies that the salic magma was emplaced passively by stoping. Lack of chill-zones suggests contemporaneity of emplacement (Ferguson, 1973). The central intrusions display both passive and forceful intrusion, the latter being demonstrated by the kimberlite and carbonatite diatremes. The Premier mine kimberlite pipe has a typical funnel-shaped vent in keeping with the concept that forcible emplacement by a gas fluidization process is a mechanism by which these bodies were emplaced (Ferguson, 1973). The C0,-poor rocks of the central intrusions appear to have been intruded passively, mainly by ring-faulting and cauldron subsidence. All the central complexes are volcanic or sub-volcanic so that ejectamenta accompany the intrusions (Ferguson, 1973). There is some evidence that the central intrusions are located at the intersection of northwest- and east-trending faults with the former being the dominant surface expression throughout the province. Ferguson (1973) notes that aeromagnetic data define long, narrow, linear easterly-trending anomalies, and suggests that the intersection of these two directions may be important in locating the sites of central intrusions. Only limited trace element studies have been undertaken on the Pilanesberg rocks. Ferguson (1973) finds that the data show covariance between Ba and K, Ba and Sr, Sr and K, A1 and Gay Zr and Hf and Nb and Ta, and that most of the rocks show Rb enrichment relative to K. With the exception of diamonds produced from the Premier kimberlite pipe no economically viable mineralization is yet known. REFERENCES Allsopp, H.L., Burger, A.J. and van Zyl, C., 1967. A minimum age for the Premier Kimberlite pipe yielded by Rb/Sr measurements, with related galena isotopic data. Earth Planet. Sci. Lett., 3: 161-166. Bisschoff, A.A., 1973. The petrology of some mafic and peralkaline intrusions in the Vredefort dome, South Africa. Trans. Geol. SOC.S. Afr., 7 6 : 27-52.

811 Davies, R.D., Allsopp, H.L., Erlank, A.J. and Manton, W.I., 1970. Sr-isotopic studies on various layered mafic intrusions in southern Africa. Geol. Soc. S. Afr., Spec. Publ., 1: 5 76-5 93. Ferguson, J., 1973. The Pilanesberg alkaline province, southern Africa. Trans. Geol. Soc. S. Afr., 76: 249-270. Hall, A.L. and Molengraaff, G.A.F., 1925. The Vredefort Mountain Land in the southern Transvaal and northern Orange Free State. Verh. K. Akad. Wet. Amsterdam, 2, pt. 24, no. 3: 1-183. Holmes, A. and Cahen, L., 1957. Geochonologie africaine 1956. Acad. R. Sci. Colon. (Brussels) C1. Sci. Nat Med., Mem. Collect. 8', N.A. Tome V, Fasc. 1: 1-169. Key, R.M., 1977. The geochronology of Botswana. Trans. Geol. Soc. S. Afr., 80: 31-42. Nel, L.T. and Jansen, H., 1957. The geology of the country around Vereeniging. Explan. Sheet 6 2 (Vereeniging). Geol. Surv., Dept. Min. S. Afr., 1-90. Oosthuysen, E.J. and Burger, A.J., 1964. Radiometric dating of some igneous intrusives associated with the Waterberg System. Ann. Geol. Surv., Dept. Min. S. Afr., 3: 87 106. Palabora Mining Company Limited Mine Geological and Mineralogical Staff, 1976. The geology and the economic deposits of copper, iron, and vermiculite in the Palabora Igneous Complex: a brief review. Econ. Geol., 71: 177-192. Rogers, A.W., 1922. The geology of the country around Heidelberg. Explan. Geol. Map. Geol. Surv., Min Dept. S. Afr., 1-84. Schreiner, G.D.L. and van Niekerk, C.A., 1958. The age of a Pilanesberg dyke from the central Witwatersrand. Trans. Geol. Soc. S. Afr., 61: 197-207. Snelling, N.J., 1963. Age determination unit. Annu. Rep. Overseas Geol. Surv. p. 30. Toens, P.D., 1952. The Geology around Leeuwfontein Northeast of Pretoria. M. Sc. Thesis, Univ. Pretoria, (unpubl.). Van Niekerk, C.B., 1962. The age of the Gemspost dyke from the Ventersdorp Gold Mine. Trans. Geol. SOC.S. Afr., 65: 105-111. Verwoerd, W.J., 1967. The carbonatites of South Africa and South West Africa. Geol. Surv., Dept. Min. S. Afr., Handb. 6 : 1-452.

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813

Chapter 12

PRECAMBRIAN PALAEONTOLOGY OF SOUTHERN AFRICA K.A. ERIKSSON and I.M. MACGREGOR

INTRODUCTION

Widespread evidence of Early and Late Precambrian life is to be found in the stratigraphic record of Southern Africa. Well-preserved Archaean and unique Early Proterozoic sedimentary sequences contain diverse assemblages of fossilized life-forms. Middle Precambrian stratigraphic units are, in contrast, poorly represented in the subcontinent. In this review evidence of Precambrian life, in the form of stromatolites, microfossils, extractable and non-extractable organic matter, carbon-isotopes, and body and trace fossils, are first documented (Table 12.1). A number of broader implications of the Precambrian fossil record in Southern Africa are then evaluated, including the age of biogenesis, the Archaean-Proterozoic boundary and the Upper Proterozoic Phanerozoic boundary. STROMATOLITES

I t is now widely acknowledged that stromatolites provide the most conclusive evidence for the existence of Early and Middle Precambrian life (see for example Schopf, 1976). Whether all such structures are necessarily algal, however, remains uncertain following the discovery of bacterial stromatolites in hot springs (Walter, 1972a). Columnar stromatolites furthermore may have biostratigraphic significance in that there is considered t o be a relationship between the species composition of the stromatolite-building algal mats and the gross morphology of the stromatolites (Kryolov and Semikhatov, 1976). Numerous palaeoenvironmental analyses have also made use of stromatolites, but the necessity for developing non-uniformitarian depositional models has been encountered by most workers (see for example Garret, 1970). The oldest recognizable stromatolites occur in the Pongola Supergroup of northern Natal and consist of weakly developed domical and columnar forms (Mason and Von Brunn, 1977). They are laterally impersistent and are associated with clastic tidal flat sediments. The preponderance of convex-up structures, coupled with the fine mm-scale laminations have been taken as

814 TABLE 12.1 Evidence of life in Southern African stratigraphic units Local stratigraphic unit

Age (X106 yr)

Evidence of life

References

Nama SWA, S.A.

> 550 probably,

stromatolites microfossils shelly fossils ediacaran fauii a trace fossils

Germs (1972a) Germs (1972a) Germs (1972b) Germs (1973,1974) Germs ( 1 9 7 2 ~ )

stromatolites

Kruger (1969), Cloud and Semikhatov

Damara SWA.

<700 Welke e t al. (1973), Kroner (1971)

*

>500 100 <1200 f 100

Clifford et al. (1969) Katanga

Umkondo R

(1969)

> 6 2 0 < 1050

stromatolites

Cahen (1970) Cahen (in Clemmey, 1976)

micro fossils trace fossils

>1100 probably < 1785

Malan (1964), Gunatilaka (1977) Binda (1977) Clemmey (1976)

stromatolites

Button (1977), Tyrrwhitt (1966)

stromatolites

Winnall (1971)

stromatolites

Button (1971), Button and Vos (1977)

stromatolites

Young (1934), Truswell and Eriksson (1972,

Vail and Dodson (1969)

Allsopp et al. (1973) Lomagundi R

>1940

* 70

Vail and Dodson (1969)

Transvaal/Griqualand West (S.A.) (1) Pretoria 2224 k 2 1 D. Crampton (pers. commun., 1972) > 2224 ?I 2 1 (2) Chuniespoort/ Campbell/ Griquatown

1973),

Eriksson and Truswell (1974), Eriksson (1977a),

Bertrand and Eriksson microorganisms

(1977), Cloud and Semikhatov (1969) Nagy (1974), Macgregor e t al. (1974),

Cloud and Licari (1968), La Berge (1973)

organic matter

Harrington and Cilliers (1963), Nagy e t al. (1976)

815 TABLE 12.1 (continued) Local stratigraphic unit

Age (X106 yr)

(3) Wolkberg

>2224

Ventersdorp S.A.

Witwatersrand S.A.

Evidence of life

References

biog. carbon stromatolites

Hoering (1962) Button (1973)

2300 f 100 Van Niekerk and Burger (1964) 2620 2 50 Van Niekerk and Burger (1977)

stromatolites

Winter (1963), Grobler and Emslie (1976), Walter (1972a,b)

<2720 f 100 Allsopp (1964)

“Lichen”

Halbauer and Van Warmelo (1974), Halbauer (1975) Schidlowski (1965), Nagy e t al. (1976) Prashnowsky and Schidlowski (1967), Oberlies and Prashnowsky (1968) Hoefs and Schidlowski (1967)

* 21

microfossils organic matter

biog. carbon

Bulawayan (R) (1)Belingwe

(2) Huntsman

* 80

stromatoli tes

Bickle et al. (1975)

Hawkesworth and Bickle (1976) 2800-2 500 Hawkesworth et al. (1975)

stromatolites

Macgregor (1940), Schopf et al. (1971), Winnall (1971) Nagy e t al. (1976) Nagy et al. (1976) Schopf et al. (1971)

< 2970 > 2700

micro fossils organic matter biog. carbon Pongola S.A.

>2870 f 30 Davies et al. (1962) <3070 f 60 Allsopp e t al. (1962)

stromatolites

Mason and Von Brunn (1977), Von Brunn and Mason (197 7)

Swaziland S.A.

=3 3 00-3 5 0 0

microfossils

Jahn and Shih (1974), Van Niekerk and Burger (1969)

organic matter

Brooks et al. (1973), Muir and Hall (1974), Muir and Grant (1976), Schopf (1970) Brooks e t al. (1973), Nagy e t al. (1976), McKirdy (1974) Oehler e t al. (1972)

biog. carbon

-

S.A. = South Africa; SWA = South West Africa; R = Zimbabwe; Z = Zambia.

816

Fig. 12.1. Contiguous domical stromatolites from the Belingwe greenstone belt, Zimbabwe (pen for scale).

evidence, albeit indirect, for an algal origin for these structures. Bulawayan stromatolites are now known from the two widely spaced localities in Zimbabwe (Fig. 12.1; MacGregor, 1940; Schopf et al., 1971; Bickle et al., 1975), both of which have now been shown t o be less than 3000 Ma old (Table 12.1). Well-developed columnar and domical forms occur within thin calcareous units. Stromatolites in the Ventersdorp Supergroup (Winter, 1963; Grobler and Emslie, 1976) and Wolkberg Group (Button, 1973) are similarly localized in their distribution and probably developed in lakes fronting alluvial fans or during periods of volcanic quiescence. The stromatolites in the Rietgat Formation of the Ventersdorp Supergroup contain the oldest named form; Gruneria f. nov (Walter, 197213) which may be represented in the Wolkberg Group as well. The Chuniespoort, Campbell and Griquatown Groups constitute the first widespread and thick development of stromatolitic shelf carbonates (Button, 1973). Evidence has been led for the accumulation of these sediments in supratidal as well as subtidal depositional environments (see for example Truswell and Eriksson, 1973, 1975; Eriksson et al., 1976). The discovery of algal filaments and unicells in these stromatolitic carbonates (Macgregor et al., 1974; Nagy, 1974) has conclusively shown them to be of algal origin. Furthermore, following the work of Walter (1975, 1976), the morphologies of finely laminated supratidal and upper-intertidal stromatolites are con-

817

Fig. 12.2. Pilbaria boetsapia columnar stromatolite from the Campbell Group; Groot Boetsap river, northwest Cape Province, South Africa.

sidered to be a response to the growth habits of different algal species (Eriksson and Truswell, 1974; Eriksson, 1977a). Various columnar stromatolites also occur. These include two forms from each of Group Pilbaria (Fig. 12.2; Walter, 197213) and Group Katerniu (Cloud and Semikhatov, 1969) which are respectively similar t o columnar stromatolite groups from the Duck Creek Dolomite of the Wyloo Group of Western Australia and possibly from the Gunflint Formation of Canada (Walter, 1972b). Other named stromatolites from the Campbell Group are, at this stage, unique t o Southern Africa and are considered to differ at the form level-from later Precambrian columns (Bertrand and Eriksson, 1977). Flat-laminated, domical and unnamed columnar stromatolites also occur in the Pretoria Group (Button, 1972; Button and Vos, 1977), the Lomagundi (Winnall, 1971; and Umkondo (Tyrrwhitt, 1966; Button, 1977) of Rhodesia, and the Kantangan of Zambia (Malan, 1964; Gunatilaka, 1977). These stromatolites are considered to have grown in a diversity of environments including supratidal, intertidal and subtidal. Widespread stromatolitic carbonates occur in the Late Precambrian and possibly partly time equivalent Nama and Damara Groups of South West Africa (Figs. 12.3 and 12.4). Cloud and Semikhatov (1969) have described Buiculia uff. B. ruru Semikhatov from the lower part of the Abenab Subgroup in the Tsumeb district and consider this a Middle Riphean form (1350 f 50 to 950 f 50 Ma); and also Conophyton ressoti Menchikoff from the upper

Fig. 12.3. Columnar strornatolite (Baicalia aff. B. raru?) from the Abenab Subgroup, South West Africa. SOUTH ERN SOUTH WEST AFRICA AGElmyl

I

5 5 3 1 30

F i s h Rjver Formotion

SO0 - 700

, 5 5 0 1 100

5%

$; ;P O ,"

TSUMEB SUBGROUP

0 k

4:

a . 0 CHUOS T I L LI TEAAA

2 $' fOQ0

q

0

P. T

a ?a.

d

Qk ZO 1100

I

Numees Formation AAA

7-7-7-

ABENAB SUBGROUP

Hilda Form at,on

Sttnkfonte,n Formation

I

-

UNCONFORMITIES

Fig. 12.4. Tentative correlations of stratigraphic units in northern and southern South West Africa (modified after Germs, 1974).

819 Abenab at the same locality, a suggested Upper Riphean form (950 f 50 to 680 f 20 Ma). Several columnar stromatolites within carbonate members of the Kuibis and Schwarzrand Formations in the Nama Group have been most tentatively identified and taken to suggest an Upper Riphean or Vendian age (950 f 50 to 570 f 20 Ma), and perhaps, but less possibly, a Lower Cambrian age (Germs, 1972a). A Middle to Upper Riphean to Vendian age for the Nama Group is compatible with available chronostratigraphic data (see Table 12.1). This evidence has led Germs (1972a) to argue that the Abenab is older than the Schwarzrand Formation and that the Chuos Tillite and Numees Formation (Fig. 12.4) may be coeval. Detailed comparative studies on the Damara and Nama stromatolites and additional age dating are required to test this hypothesis. MICROFOSSILS, ORGANIC MATTER AND CARBON ISOTOPES

The recognition of microfossils and organic matter in Precambrian sediments has far reaching implications with respect to the origin and evolution of life prior to 600 Ma ago. Considerable debate continues, however, as to the syngenetic nature of certain microfossil and extractable organic matter discoveries with the possibility of contamination existing (see for example McKirdy, 1974; Schopf, 1976). Biogenic carbon-isotope ratios from insoluble organic matter are generally taken as reliable evidence for Precambrian organic activity. Scant, and often controversial evidence of life has been discovered in chemical sediments from the Swaziland Supergroup in the Barberton Mountain Land. Cherts in the Theespruit, Kromberg (Brooks et al., 1973; Muir and Hall, 1974) and Swartkoppie Formations (Engel et al., 1968; Nagy and Nagy, 1969; Nagy, 1971) of the volcanic Onverwacht Group; and the sedimentary Fig Tree Group (Barghoorn and Schopf, 1966; Schopf, 1970), contain microstructures. The Fig Tree structures have generally been accepted as biogenic (Barghoorn and Schopf, 1966) although Schopf (1975) has recently questioned the authenticity of these. Scepticism has also been expressed concerning the microstructures in the Onverwacht (see for example Engel et al., 1968; Schopf, 1975). The discovery (Brooks et al., 1973; Muir and Hall, 1974) and recent reappraisal (Muir and Grant, 1976) of unicells, paired unicells, agglomerations of cells and filaments in the Onverwacht Group has been taken to suggest more strongly that these are biogenic. Especially compelling evidence, such as chains of small cells, terminations of filaments resembling pointed terminal cells of blue-green algae, and probable abandoned sheaths, led Muir and Grant (1976) to conclude that the microstructures in the Onverwacht sediments are blue-green algal or flexibacterial remains. Insoluble organic matter, suitable for isotopic studies and resembling naturally occurring substances of known biogenic origin such as aromatic hy-

820

drocarbons, have been recovered from Onverwacht cherts (Brooks et al., 1973; Nagy et al., 1976). Oehler et al. (1972) noted that carbon-isotopic values from the Lower Onverwacht were anomalously heavy in comparison with those from the Upper Onverwacht. They suggested that these differences may reflect a major event in biological evolution from non-photosynthetic to photosynthetic organisms. Muir and Grant (1976), however, attribute this change in carbon-isotopic values to variable degrees of metamorphism. Organic components, including aliphatic aromatic hydrocarbons, have been extracted from sediments of the Onverwacht and Fig Tree Groups (McKirdy, 1974). If it can be shown unequivocally that these date from the time of sedimentation, rather than having been introduced later, the presence of such compounds as porphyrins will indicate that photosynthesis occurred at least 3300 Ma ago (McKirdy, 1974). The Bulawayan limestone in Zimbabwe contains morphologically complex unicellular microfossils as well as different types of hydrocarbons (Nagy et al., 1976) while carbon-isotope ratios provide evidence of photosynthetic fractionation (Schopf et al., 1971). Although sceptical about microfossils in sediments of this age and older, Schopf (1976) acknowledges that the most plausible interpretation for insoluble carbonaceous material is some form of primitive biological activity. Halbauer and Van Warmelo (1974) and Halbauer (1975) have recognized a well-differentiated plant-life in the Witwatersrand Supergroup, including bacteria and algae as well as the evolutionary highly advanced fungi. They proposed that lichen-like plants existed, comprising a symbiosis of algae and fungi. Cloud (1976), however, has produced non-biogenic replicates of these structures which he considers to be artifacts of preparation. Mineralized bacteria (Schidlowski, 1965) and unicellular and filamentous cyanophytes or blue-green algae (Nagy et al., 1976) appear to represent more authentic and certainly less controversial Witwatersrand microfossils. Extractable organic compounds such as amino acids and carbohydrates (Prashnowsky and Schidlowski, 1967; Oberlies and Prashnowsky, 1968) encounter the same syngenetic-epigenetic controversy as do those in older stratigraphic units. McKirdy (1974) argues that carbohydrates and amino acids lack geochemical stability and are almost certainly not syngenetic. Insoluble aromatic hydrocarbons (Nagy et al., 1976), from which biogenic carbon-isotope ratios have been obtained (Hoefs and Schidlowski, 1967), indicate biological activity at the time of deposition of the auriferous Witwatersrand sediments. Filamentous microfossils from the Chuniespoort, Campbell and Griquatown Groups (Macgregor et al., 1974; Nagy, 1974) constitute the oldest known evidence of cell diversification. A number of representatives of unbranched filamentous cyanophytes have been recognized, including Nostocaceae, Oscillatoriaceae and Rivulariaceae (Fig. 12.5). These are frequently multicellular, with distinct crosswalls, and often contain structures analogous t o akinetes or heterocysts in modem blue-green algae. Simple

821

Fig. 1 2 . 5 . Algal filament resembling the modern Rivulariaceae from the Campbell Group, northwest Cape Province, South Africa (X550).

spherical structures resembling unicellular or coccoid cyanophytes have also been reported from both the carbonates (Nagy, 1974) and banded ironformations (Cloud and Licari, 1968; La Berge, 1973). Reports of amino acids associated with asbestos from this same stratigraphic unit (Harrington and Cilliers, 1963) should be treated with reservation as a possible contaminant. Nagy et al. (1976) have isolated a number of advanced organic compounds, of probable carbohydrate derivation, from insoluble carbonaceous residues. Isotope ratios point t o a biogenic origin for the carbon (Hoering, 1962). The Lower Kundelungu sediments of the Zambian Copperbelt contain filamentous microfossils of probable algal affinity (Binda, 1976). Filaments are frequently branched and consist of several connected cells which occasionally separate from the thallus by constriction. Spherical structures, displaying a considerable variation and of unknown biological affinity, have also been reported (Binda, .1976). These occur singly or in agglomerations and have walls which may be smooth or sculptured. It is probably that a number of these spherical structures are deteched portions of the filamentous microfossils. No specific microfossils have been found in the Nama Group but several

822 possible biogenic structures have been recognized. Vaguely defined algae, together with possible reproductive structures of algal affinity, including acritarchs, palynomorphs and catagraphs, have been recognized (Germs, 1972a). BODY AND TRACE FOSSILS

Body or trace fossils, which commonly occur in Upper Precambrian stratigraphic units, provide evidence for advanced or metazoan life. Metazoa are multicellular animals, including such phylla as mollusca, coelenterata, arthropoda, annelida and echinodermata, whose cells are differentiated into tissues or organs that are specialized for different functions. Trace fossils have recently been discovered in the circa 1000 Ma Katangan of Zambia (Clemmey, 1976), thus indicating the existence of advanced or metazoan life some 300 Ma earlier than was previously known. Two types of burrows occur in the Mixed Formation of the Roan Group. Both varieties are dominantly horizontal with respect t o the bedding although occasional oblique-types do occur. They range from 2 to 1 0 mm in diameter and while one variety has a maximum length of 20 cm the second may be up to a metre in length. The shorter of the burrows contains walls which are laminated parallel to their long axes and also well-developed “spreite” which probably indicate active backfilling by the burrowing organism. The longer burrow-type displays less evidence of active infilling. Branching is common in both varieties and is dominantly y-shaped. The first major development of metazoans in Southern Africa occurs in the Nama Group (Germs, 1972a). The Nama Group contains both body, including representatives of the soft-bodied Ediaran fauna, and shelly fossils, and trace fossils. Ediacaran body fossils have been recovered from the Kuibis and Schwartzrand Formations of the Nama Group (Germs, 1972a, 1973 and 1974). These comprise Sprigginid worms as well as the dominant frond or “leaf-like” organisms which have been assigned to the group Petalonamae, in which five forms have been recognized: (1)Rangea: This form has a leaf-shaped main body with a medial axis, lateral branches separated by transverse lateral furrows (Fig. 12.6). (2) Pteridium : The body of Pteridium has a complex horseshoe-form built up by leaf-like bodies called petaloids which were possibly specialized in respertational functions. This form displays some features common to Echinoderms and Coelenterates. The form Cyclomedusa has been recognized by Germs (1974) within the phylum Coelenterata as a circular organism with a diameter range from 15 to 46 mm. Prominent concentric grooves with numerous radial striae, a raised central knob and a ring canal characterize this body fossil. (3) Erniettu: This is a sessile fossil with a thick dorsal exoskeleton which is

823

Fig. 12.6. Rangea schneiderhohni; an Ediacaran body fossil from the Schwarzrand Formation, Nama Group, South West Africa (photograph taken by G.J.B. Germs).

bent ventrally into a horseshoe-shape. The exoskeleton is built up of segments which are orientated left and right from a medial zig-zag line. (4) Nasepia: This consists of leaf-like bodies with thin ribs which are mostly oriented in one direction and subparallel to the long axis. The leaflike bodies (petaloids) are nearly flat or strongly convex and have a skeletal structure. The petaloids are generally bundled together t o form a petalodium. (5) Namalia: A conical morphology with longitudinal ridges characterizes this form. I t is oval to round in cross-section and there is evidence for possible articulation. Namalia strongly resembles an individual petaloid of Nasep ia . The soft-bodied Ediacaran fauna provide the most specific evidence for the development of metazoans in the Late Precambrian (Glaessner, 1971). Coelenterata appear to be the most abundant of the Ediacaran and, in addition to annelids (worms), representatives of the Echinodermata may also occur in the Nama Group. In addition t o the Ediacaran fauna, shelly fossils, resembling tubiculous annelids (Fig. 12.7) and probably related t o the Cribricyathids, also occur in limestones of the Schwarzrand and Kuibis Formations (Germs, 197213). These fossils consist of sinuous calcareous tubes composed of stacked cones so arranged as to give the appearance of a composite structure having two walls.

824

Fig. 1 2 . 7 . Shelly fossil; Cloudina hartmannae from the Kuibis Formation, Nama Group, South West Africa (photograph taken by G.J.B. Germs).

Fig. 12.8. Phycodes p e d u m trace fossil from the Fish River Formation, Nama Group, South West Africa (photograph taken by G.J.B. Germs).

825 Trace fossils are common in sediments in all three formations of the Nama Group but are most prolific in sandstones, siltstones and shales of the Schwarzrand Formation. Numerous types of traces have been recognized including Phycodes pedum (Fig. 12.8), Planolites, Helminthoichnites and Skolithos (Germs, 1972c) and Buchholzbrunnichnus kroneri (Germs, 1973). These are dominantly tubular but also include double-spiral burrows and trails. It is difficult to state which organisms were responsible for the Nama traces but Phycodes pedum and Skolithos, amongs others, were probably formed by worm-like animals. Trails with two or three longitudinal ridges and transverse ribs, specifically Buchholzbrunnichnus kroneri, may have been formed by molluscs (Germs, 1972c, 1973). With exception of Phycodes pedum, which suggests a late-Precambrian to upper Cambrian age, the trace fossils in the Nama Group are of little chronostratigraphic importance. BROADER IMPLICATIONS OF T H E PRECAMBRIAN FOSSIL RECORD

The age of biogenesis The recognition of microfossils in cherts of the Onverwacht and Fig Tree Groups has shown that life in the Archaean extended back to at least 3400 Ma. Older banded iron-formations may provide indirect evidence for the existence of earlier life. Cloud (1974, 1976) has suggested that there may be a link between the production of oxygen by primitive photosynthetic microorganisms and its consumption by reaction with ferrous iron to form hematite and similarly oxidized mineral species. Extrapolation back in time in search of an age of biogenesis encounters a murky rock record and the enigmatic problems of both the nature and age of the original crust. In Southern Africa the Ancient Gneiss Complex of Swaziland, which contains banded iron-formations, may (Hunter, 1970) or may not (Viljoen and Viljoen, 1969; Anhaeusser, 1973) represent an earlier preOnverwacht floor. Geochronological studies have not, as yet, resolved this controversy as Davies and Allsopp (1976) consider that an extensive siallic crust did not begin forming much before 3500 Ma ago. Similar ages have been obtained for early gneisses and granites in Zimbabwe (Hickmann, 1974, 1976; Hawkesworth et al., 1975). Within the Limpopo mobile belt a major period of metamorphism, probably involving sedimentary rocks, has been dated at 3 8 6 0 t 100 Ma (Barton and Ryan, 1977). Younger cover rocks, which include banded iron-formations, are younger than 3640 t 100 Ma and older than 3250 Ma (Barton and Ryan, 1977). Banded iron-formations in Southern Africa can thus not be shown conclusively to pre-date the Onverwacht Group while relatively unmetamorphosed sediments older than this group, and which may contain microfossils, are not known from the subcontinent. At this stage the Onverwacht Group contains the oldest evidence of life, not only in Southern Africa but in the world, and

826 it is plausible that 3500 Ma represents the age of biogenesis. The c. 3800 Ma. Isua iron ores of Greenland (Moorbath, 1975) contain neither organic compounds nor microfossils (Nagy et al., 1976) suggesting that oxygen in hematite could have a non-biogenic source. Photodissociation of water induced by ultra-violet light is one possible source (Schopf, 1976).

The Archaean-Proterozoic

boundary

Numerous general terms have been suggested as subdivisions of the Precambrian (see for example Hedberg, 1974; Cloud, 1976). Most of them have been discarded, and the only two that have remained in common usage are Proterozoic and Archaean. Rankama (1970), in a survey of the literature, found that the age of the Proterozoic-Archaean boundary ranged from 1700-3000 Ma, while Salop (1972) suggested a figure of 3500 Ma for this boundary. The Archaean-Proterozoic boundary in Southern Africa is currently considered t o be represented by the unconformable contact between the granite-greenstone terrain, termed the basement, and the younger cratonic basins (De Villiers, in Rankama, 1970). Material lying above this contact in South Africa varies in age from 3100 Ma t o 2300 Ma where the Pongola and Transvaal Supergroups rest on the basement (see Table 12.1). Furthermore, in Zimbabwe Archaean-style volcanism and tectonism continued until at least 2600 Ma. These figures highlight the need for the recognition of an event of chronostratigraphic significance by which the Archaean-Proterozoic boundary can be more accurately defined. It is believed that the unconformity separating the Chuniespoort, Campbell and Griquatown Groups from the overlying Pretoria and Postmasburg Groups represents an Archaean-Proterozoic boundary which has biogenic as well as atmospheric significance. Photosynthetic conversion of carbon dioxide t o oxygen occurred through a considerable time period during the accumulation of the underlying, earliest widespread stromatolitic carbonates, and, as manifested in lithologies above and below the unconformity, had a profound effect on the composition of the atmosphere. The stratigraphic record below this unconformity contains abundant evidence for a reducing atmosphere. Firstly, alluvial sediments, such as occur in the Moodies Group (Eriksson, 1977b), the Witwatersrand Supergroup (Vos, 1975) and the Black Reef Formation, are ubiquitously pale-coloured with no indication of iron oxidation during weathering. Secondly, the presence of detrital pyrite and uraninite in the Witwatersrand Supergroup has often been taken as indicative of a reducing atmosphere (see for example Cloud, 1972). Thirdly, dolomites in the Chuniespoort, Campbell and Griquatown Groups are commonly ferruginous and manganiferous, indicating that these metals were taken into solution in their reduced states before being precipitated within the dolomite lattice in the same divalent forms (Eriksson e t al., 1975). Immediately above the erosional unconformity both iron and manganese

827 exist exclusively in their higher valency states. The palaeokarst breccia, which commonly occupies the erosional surface, is characteristically rich in tri- and quadri-valent manganese, either in the form of amorphous wad or less commonly pyrolucite (Button, 1974). Rich hematitic and manganiferous ore deposits rest on this unconformity in the northern Cape Province and are probably, in part, syngenetic (Eriksson and Truswell, 1978) while the same stratigraphic unit in the Transvaal, namely the Timeball Hill Formation, contains oolitic ironstones which are considered t o provide evidence of colloidal transport of trivalent iron (Garrels et al., 1973) and deposition mder oxidizing conditions (Button, 1974). Unquestionably syngenetic manganese deposits occur higher up in the Postmasburg Group (Boardman, 1964) and developed as a result of the geochemical separation of manganese in an oxidizing environment. The Waterberg Supergroup with an age of c. 1800 Ma is comprised largely of red beds which are taken as conclusive evidence of an oxidizing atmosphere. Even older red beds may occur above the erosional unconformity in the Transvaal Supergroup. In addition to the foregoing lithologic changes which can be used to define the Archaean-Proterozoic boundary, a number of palaeontological criteria are also of chronostratigraphic significance in this regard. The columnar stromatolite Gruneria f. nov. appears to be confined to Archaean strata. In addition to the Ventersdorp occurrence this form is also known from the Fortescue Group in the Hamersley basin of Western Australia (Walter, 1972b) which has a similar stratigraphic setting t o the Wolkberg Group. Within Southern Africa the Groups Pilbaria and Katernia are of Archaean age but in Australia the former occurs above an equivalent erosional unconformity in the Hamersley basin while a Katernia-like form occurs in the younger Gunflint strata of Canada (Walter, 1972b). Microorganisms also display characteristic changes around 2300 Ma. The first known evidence of cell diversification is found in the Chuniespoort and Campbell Groups (Nagy, 1974) and an evolutionary development of unicellular and filamentous algae occurred from 2300 Ma onwards (Schopf, 1976).

The Proterozoic and metazoans The base of the Cambrian is commonly defined by the first appearance of shelly fauna (Rosanov, 1967; Zhuravleva, 1970; and with certain qualifications, Glaessner, 1972) and arbitrarily defines the end of the Precambrian. According to the most widely held view (Stasek, 1972) Cambrian molluscs were derived from free-living flatworms. These developed a dorsal protective mucoid coat which gradually turned into a tough cuticle and formed a substrate for a subsequently developed calcareous exoskeleton. The earlier part of this evolutionary process must have taken place in the Late Precambrian (Runnegar and Pojeta, 1974). On a broader scale, if it is accepted that such varied Cambrian phyta as annelids, molluscs.and arthropods belong t o a com-

828 mon lineage (Runnegar et al., 1975), the conclusion can be drawn not that oxygen levels had of necessity to reach a specific level before calcareous exoskeletons could develop (Berkener and Marshall, 1966; Rhoads and Morse, 1971), but that this development took place within the latest Precambrian to Palaeozoic evolutionary development. Geological and biological evidence suggests a transition from a reducing t o an oxidizing atmosphere long before the end of the Precambrian. This change, of necessity, pre-dated the origin of aerobic eucaryotes which, in turn, must have preceded the evolution of metazoa (Margulis et al., 1976). A consideration of the soft-bodied Ediacaran fauna provides further support for the development of metazoans through an evolutionary process in the Late Precambrian. Absolute age determinations indicate that the fauna existed between 600 and 700 Ma ago (Glaessner, 1971). Thus, if the suggestion by Cloud (1976) is followed and the base of the Phanerozoic placed at the first appearance of metazoans, the Precambrian-Phanerozoic boundary would be dated at 680-700 Ma. In most parts of the world a pre-Cambrian or Vendian time unit, corresponding t o the occurrence of the Ediacaran fauna, would appear at the base of the Phanerozoic and pre-date the appearance of shelly fauna. In the Nama Group of South West Africa the presence of calcareous annelid tubes or Cribricyathids in association with the Ediacaran (Germs, 1974), introduces a complication and led Glaessner (1971) to suggest a Cambrian age for the Nama Group with the base of the Phanerozoic then coinciding with the base of the Cambrian at c. 700 Ma. The high degree of morphological organization exhibited by the Ediacaran fauna (Hoffman, 1972) implies a pre-700 Ma evolution of, as yet, unrecognized metazoa which may be manifested in the c. 1000 Ma trace fossils of the Kundelungu (Clemmey, 1976). REFERENCES Allsopp, H.L., 1964. Rubidium/strontium ages from the western Transvaal. Nature, 204: 361-363. Allsopp, H.L., Erlank, A.J. and Hornung, G., 1973. Rb-Sr age measurements o n the Umkondo Dolerite of Rhodesia. Geol. SOC.S. Afr., Spec. Publ., 3: 217. Allsopp, H.L., Roberts, H.R., Schreiner, G.D.L. and Hunter, D.R., 1962. Rb-Sr age measurements on various Swaziland granites. J. Geophys. Res., 67 : 5307-5313. Anhaeusser, C.R., 1973. The evolution of the early Precambrian crust of southern Africa. Philos. Trans. R. SOC.London, Ser. A, 273: 359-388. Barghoorn, E.S. and Schopf, J.W., 1966. Microorganisms three billion years old from the Precambrian of South Africa. Science, 152: 758-763. Barton, J.M. and Ryan, B., 1977. A review of the geochronologic framework of the Limpopo Mobile Belt. Botswana Geol. Sum. Bull., 1 2 : 183-200. Berkener, L.V. and Marshall, L.C., 1966. On the origin and rise of oxygen concentration in the Earth’s atmosphere. J. Atmos. Sci., 22: 225-261. Bertrand, J. and Eriksson, K.A., 1977. Columnar stromatolites from the early Proterozoic Schmidtsdrift Formation, northern Cape Province, South Africa - Part 1: systematic and diagnostic features. Palaeontol. Afr., 20: 1-26.

829 Bickle, M.J., Martin, A. and Nisbet, E.G., 1975. Basaltic and peridotitic komatiites and stromatolites above a basal unconformity in the Belingwe greenstone belt, Rhodesia. Earth Planet. Sci. Lett., 27: 155-162. Binda, P.L., 1977. Microfossils from the Lower Kundelungu (late Precambrian) of Zambia. Precambrian Res., 4: 235-306. Boardman, L.G., 1964. Further geological data on the Postmasburg and Kuruman manganese ore deposits, northern Cape Province. In: S.H. Haughton (Editor), The Geology of Some Ore Deposits in Southern Africa, 11. Geol. SOC.S. Afr., pp. 415-441. Brooks, J., Muir, M.D. and Shaw, G., 1973. Chemistry and morphology of Precambrian microorganisms. Nature, 224: 215-217. Button, A., 1971. Early Proterozoic algal stromatolites of the Pretoria Group, Transvaal Sequence. Trans. Geol. Soc. S. Afr., 74: 201-210. Button, A., 1973. Algal stromatolites from the early Proterozoic Wolkberg Group, Transvaal Sequence. J . Sediment. Petrol., 43: 160-167. Button, A., 1974. A Regional Study of the Stratigraphy and Development of the Transvaal Basin in the Eastern and Northeastern Transvaal. Ph.D. Thesis, Univ. Witwatersrand, Johannesburg, (unpubl.). Button, A., 1977. Stratigraphic history of the middle Proterozoic Umkondo basin in the Chipinga area, southeastern Rhodesia. Econ. Geol. Res. Unit, Inf. Circ. 108, Univ. Witwatersrand, 32 pp. Button, A. and Vos, R.G., 1977. Subtidal and intertidal clastic and carbonate sedimentation in a macrotidal environment: an example from the lower Proterozoic of South Africa. Sediment. Geol., 18: 175-200. Cahen, L., 1970. Igneous activity and mineralisation episodes in the evolution of the Kibaride and Katangide orogenic belts of central Africa. In: T.N. Clifford and I.C. Gass (Editors), African Magmatism and Tectonics. Oliver and Boyd, Edinburgh, pp. 99117. Clemmey, H., 1976. World’s oldest animal traces. Nature, 261: 576-578. Clifford, T.N., Rooke, J.M. and Allsopp, H.L., 1969. Petrochemistry and age of the Franzfontein granitic rocks of northern South West Africa. Geochim. Cosmochim. Acta, 33: 973-986. Cloud, P.E., 1972. A working model for the primitive earth. Am. J. Sci., 272: 537-548. Cloud, P.E., 1974. Evolution of ecosystems. Am. Sci., 62: 54-66. Cloud, P.E., 1976. Major features in crustal evolution. Trans. Geol. SOC.S. Afr., Annex. vol. 79, 32 pp. Cloud, P.E. and Licari, G.R., 1968, Microbiotas of the banded iron formations. Proc. Natl. Acad. Sci. U.S.A., 61: 779-746. Cloud, P.E. and Semikhatov, M.A., 1969. Proterozoic stromatolite zonation. Am. J. Sci., 267,1017-1061. Davies, R.D. and Allsopp, H.L., 1976. Strontium isotopic evidence relating to the evolution of the lower Precambrian granitic crust in Swaziland. Geology, 4: 553-556. Davies, R.D., Allsopp, H.L., Erlank, A.J. and Manton, W.I., 1962. Sr-isotopic studies on various layered mafic intrusions in southern Africa. Geol. SOC.S. Afr., Spec. Publ., 1: 5 7 6-593. Engel, A.E.J., Nagy, B., Nagy, L.A., Engel, C.G., Kremp, G.O.W. and Drew, C.M., 1968. Alga-like forms in Onverwacht Series, South Africa: oldest recognized life-like forms on earth. Science, 161: 1005-1008. Eriksson, K.A., 1977a. Tidal flat and subtidal sedimentation in the 2250 m.y. Malmani Dolomite, Transvaal, South Africa. Sediment. Geol., 18: 223-244. Eriksson, K.A., 1977b. A Palaeoenvironmental Analysis of the Archaean Moodies, Group, Barberton Mountain Land, South Africa. Ph.D. Thesis, Univ. Witwatersrand, Johannesburg (unpubl.).

Eriksson, K.A. and Truswell, J.F., 1974. Tidal flat associations from a Lower Proterozoic carbonate sequence in South Africa. Sedimentology, 21 : 293-309. Eriksson, K.A. and Truswell, J.F., 1978. Geological processes and atmospheric evolution in the Precambrian. In: D.H. Tarling (Editor), Evolution of the Earth’s Crust. Academic Press, London, pp. 219-238. Eriksson, K.A., McCarthy, T.S. and Truswell, J.F., 1975. Limestone formation and dolomitization in a lower Proterozoic succession from South Africa. J. Sediment. Petrol., 45: 604-614. Eriksson, K.A., Truswell, J.F. and Button, A., 1976. Palaeoenvironmental and geochemical models from a lower Proterozoic carbonate succession in South Africa. In: M.R. Walter (Editor), Stromatolites, Developments in Sedimentology. Elsevier, Amsterdam, pp. 635-643. Garrels, R.M., Perry, E.A. and MacKenzie, F.T., 1973. Genesis of Precambrian iron formations and the development of atmospheric oxygen. Econ. Geol., 68: 1173-1179. Garrett, P., 1970. Phanerozoic stromatolites: non-competitive ecologic restriction by grazing and burrowing animals. Science, 169: 171-173. Germs, G.J.B., 1972a. The stratigraphy and palaeontology of the lower Nama Group, South West Africa. Precambrian. Res. Unit. Bull. 1 2 , Univ. Cape Town. Germs, G.J.B., 1972b. New shelly fossils from the Nama Group, South West Africa. Am. J. Sci., 272: 752-761. Germs, G.J.B., 1972c. Trace fossils from the Nama Group, South West Africa. J. Paleontol., 46: 864-870. Germs, G.J.B., 1973. Possible sprigginid worm and a new trace fossil from the Nama Group, South West Africa. Geology, 1: 69-70. Germs, G.J.B., 1974. The Nama Group in South West Africa and its relationship to the Pan-African geosyncline. J. Geol., 82: 301-317. Glaessner, M.F., 1971. Geographic distribution and time range of the Ediacara Precambrian fauna. Geol. SOC.Am. Bull., 82, 509-514. Glaessner, M.F., 1972. Precambrian palaeozoology. Univ. Adelaide, Centre for Precambrian Research, Spec. Pap., 1: 43-52. Grobler, N.J. and Emslie, D.P., 1976. Stromatolitic limestone and chert in the Ventersdorp Supergroup at the T’kuip Hills area and surroundings, Brits District, South Africa. Trans. Geol. SOC.S. Afr., 79: 49-52. Gunatilaka, A., 1977. Environmental significance of Upper Proterozoic algal stromatolites from Zambia. In: E. Fliigel (Editor), Fossil Algae. Springer-Verlag, Berlin, pp. 74-79. Halbauer, D.K., 1975. The plant origin of the Witwatersrand “carbon”. Miner. Sci. Eng., 7 : 111-131. Halbauer, D.K. and Warmelo, K.B., 1974. Fossilized plants in thucholite from Precambrian rocks of the Witwatersrand, South Africa. Precambrian Res., 1: 199-212. Harrington, J.S. and Cilliers, J.J. le R., 1963. A possible origin of the oils and amino acids isolated from amphibole asbestos and banded ironstones. Geochim. Cosmochim. Acta, 27: 411-418. Hawkesworth, C.J. and Bickle, M.J., 1976. Rhodesian Rb-Sr Geochronology from 3.6 t o 2.6 b.y.: a brief review. Univ. Leeds Res. Inst. Afr. Geol. Annu. Rep. Hawkesworth, C.J., Moorbath, S., O’Nions, R.K. and Wilson, J.F., 1975. Age relations between greenstone belts and granites in the Rhodesian Archaean craton. Earth Planet. Sci. Lett., 25: 251-262. Hedberg, H.D., 1974. Basis for chronostratigraphic classification of the Precambrian. Precambrian Res., 1 : 165-177. Hickman, M.H., 1974. 3500-m.y. old granite in southern Africa. Nature, 251: 295. Hickman, M.H., 1976. Geochronological Investigations in the Limpopo Belt and Part of the Adjacent Rhodesian Craton. Ph.D. Thesis, Univ. Leeds, 188 pp. (unpubl.).

Hoefs, J. and Schidlowski, M., 1967. Carbon isotope composition of carbonaceous matter from the Precambrian of the Witwatersrand System. Science, 155: 1096-1097. Hoering, T.C., 1962. The stable isotopes of carbon in the carbonate and reduced carbon of Precambrian rocks. Carnegie Int. Washington, Yearb., 61 : 190-191. Hofmann, H.J., 1972. Precambrian remains in Canada: fossils, dubiofossils and pseudofossils. Int. Geol. Congr. 24th Sess., 1 , Precambrian Geol., pp. 20-31. Hunter, D.R., 1970. The Ancient Gneiss Complex in Swaziland. Trans. Geol. SOC.S. Afr., 73: 107-150. Jahn, B.M. and Shih, C.Y., 1974. On the age of the Onverwacht Group, Swaziland Sequence, South Africa. Geochim. Cosmochim. Acta, 38: 873-885. Kroner, A., 1971. Late-Precambrian correlations and the relationship between the Damara and Nama Systems of South West Africa. Geol. Rundsch., 60: 1513-1523. Kruger, L., 1969. Stromatolites and oolites in the Otavi Series, S.W.A. J. Sediment. Petrol., 36: 1046-1056. Krylov, I.N. and Semikhatov, M.A., 1976. Table of time-ranges of the principal groups of Precambrian stromatolites. In: M.R. Walter (Editor), Stromatolites, Developments in Sedimentology, 20. Elsevier, Amsterdam, pp. 693-694. La Berge, G.L., 1973. Possible biological origin of Precambrian iron-formations. Econ. Geol., 68: 1099-1109. Malan, S.P., 1964. Stromatolites and other algal structures a t Mufulira, Northern Rhodesia. Econ. Geol., 59: 397-415. Margulis, L., Walker, J.C.G. and Rambler, M., 1976. Reassessment of roles of oxygen and ultraviolet light in Precambrian evolution. Nature, 264: 620-624. Mason, T.R. and Von Brunn, V., 1977. 3-Gyr-old stromatolites from South Africa. Nature, 266: 47-49. MacGregor, A.M., 1940. A pre-Cambrian limestone in southern Rhodesia. Trans. Geol. Soc. S. Afr., 43: 9-15. Macgregor, I.M., Truswell, F.J. and Eriksson, K.A., 1974. Filamentous algae from the 2300 m.y. old Transvaal Dolomite. Nature, 247: 538-540. McKirdy, D.M., 1974. Organic geochemistry in Precambrian research. Precambrian Res., 1: 75-137. Moorbath, S., 1975. Constraints for the evolution of Precambrian crust from strontium isotopic evidence. Nature, 254: 395. Muir, M.D. and Grant, P.R., 1976. Micropalaeontological evidence from the Onverwacht Group, South Africa. In: B.F. Windley (Editor), The Early History of the Earth. Wiley Interscience, New York, N.Y., pp. 595-604. Muir, M.D. and Hall, D.O., 1974. Diverse microfossils in Precambrian Onverwacht group rocks of South Africa. Nature, 252: 376-378. Nagy, B. and Nagy, L.A., 1969. Early Precambrian Onverwacht microstructures: possibly the oldest fossils on earth. Nature, 223: 1226-1229. Nagy, B., Nagy, L.A., Zumberge, J.E., Sklarov, D.S. and Anderson, P., 1976. Biological evolutionary trends and related aspects of carbon chemistry during the Precambrian, between 3800 m.y. and 2300 m.y. Abstr. Int. Geol. Congr., 25th Sess., 1, 33-34. Nagy, L.A., 1971. Ellipsoidal microstructures of narrow size range in the oldest known sediments on earth. Grana, 1 1 : 91-94. Nagy, L.A., 1974. Transvaal stomatolite: first evidence for the diversification of cells about 2.2 X l o 9 years ago. Science, 183: 514-516. Oberlies, F. and Prashnowsky, A.A., 1968. Biogeochemische und elektromikroskopische Untersuchung prakambrischer Gesteine. Naturwissenschaften, 55 : 25-28. Oehler, D.Z., Schopf, J.W. and Kvenvolden, K.A., 1972. Carbon isotopic studies of organic matter in Precambrian rocks. Science, 175: 1246-1248. Prashowsky, A.A. and Schidlowski, M., 1967. Investigation of Precambrian thucholite. Nature, 216: 560-563.

832 Rankama, K., 1970. Global Precambrian stratigraphy: background and principles. Scientia, 7th Ser., 105: 382-421. Rhoads, D.C. and Morse, J.W., 1971. Evolutionary and ecologic significance of oxygendeficient marine basins. Lethaia, 4: 413. Rosanov, A.Yu., 1967. The Cambrian boundary problem. Geol. Mag., 104: 415-434. Runnegar, B. and Pojeta, J., 1974. Molluscan phylogeny: the paleontological viewpoint. Science, 186: 311-317. Runnegar, B., Pojeta, J., Morris, N.J., Taylor, J.D., Taylor, M.E. and McClung, G., 1975. Biology of the Hyolitha. Lethaia, 8: 181-191. Salop, L.I., 1972. A unified stratigraphic scale of the Precambrian. Int. Geol. Congr., 24th Sess., I, Precambrian Geol., pp. 253-259. Schidlowski, M., 1965. Probable life-forms from the Precambrian of the Witwatersrand System (South Africa). Nature, 205: 895-896. Schopf, J.W., 1970. Precambrian microorganisms and evolutionary events prior to the origin of vascular plants. Biol. Rev. Cambridge Philos. Soc., 45: 319-352. Schopf, J.W., 1975. Precambrian paleobiology: problems and perspectives. Annu. Rev. Earth Planet. Sci., 3: 212-249. Schopf, J.W., 1976. Evidence of Archaean life: a brief appraisal. In: B.F. Windley (Editor), The Early History of the Earth. Wiley Interscience, New York, N.Y., pp. 589-593. Schopf, J.W., Oehler, D.Z., Horodyski, R.J. and Kvenvolden, K.A., 1971. Biogenicity and the significance of the oldest known stromatolites. J. Paleontol., 45: 477-485. Stasek, C.R., 1972. The molluscan framework. Chem. Zool., 6: 1-44. Truswell, J.F. and Eriksson, K.A., 1972. The morphology of stromatolites from the Transvaal Dolomite northwest of Johannesburg, South Africa. Trans. Geol. SOC.S. Afr., 75: 99-110. Truswell, J.F. and Eriksson, K.A., 1973. Stromatolitic associations and their palaeoenvironmental significance: a re-appraisal of a Lower Proterozoic locality from the north Cape Province, South Africa. Sediment. Geol., 10: 1-23. Truswell, J.F. and Eriksson, K.A., 1975. A palaeoenvironmental interpretation of the early Proterozoic Malmani Dolomite from Zwartkops, South Africa. Precambrian Res., 2: 277-303. Tyrrwhitt, D.S., 1966. Petrology, mineralogy and genesis of some copper deposits in the Middle Sabi Valley, Rhodesia. Ph.D. Thesis, Univ. Leeds (unpubl.). Vail, J.R. and Dodson, M.M., 1969. Geochronology of Rhodesia. Trans. Geol. SOC.S. Afr., 72: 79-113. Van Niekerk, C.B. and Burger, A.J., 1964. The age of the Ventersdorp System. Ann. Geol. Surv. S. Afr., 3: 75-86. Van Niekerk, C.B. and Burger, A.J., 1969. A note on the minimum age of the acid lava of the Onverwacht Series of the Swaziland System. Trans. Geol. SOC.S. Afr., 72: 9-21. Van Niekerk, C.B. and Burger, A.J., 1977. A review of the age of the Ventersdorp Supergroup. Abstr., 1 7 t h Annu. Congr., Geol. SOC.s. Afr., pp. 132-135. Viljoen, M.J. and Viljoen, R.P., 1969. An introduction to the geology of the Barberton granite-reenstone terrain. Geol. SOC.S. Afr., Spec. Publ., 2: 9-28. Vos, R.G., 1975. An alluvial plain and lacustrine model for the Precambrian Witwatersrand deposits of South Africa. J. Sediment. Petrol., 45: 480-493. Von Brunn, V. and Mason, T.R., 1977. Siliciclastic-carbonate tidal deposits from the 3000 m.y. Pongola Supergroup, South Africa. Sediment. Geol., 18: 245-255. Walter, M.R., 1972a. Siliceous algal and bacterial stromatolites in hot spring and geyser effluents of Yellowstone National Park. Science, 178: 402-405. Walter, M.R., 1972b. Stromatolites and the biostratigraphy of the Australian Precambrian and Cambrian. Palaeontol. Assoc. London, Spec. Paper 11, 1 9 0 pp.

833 Walter, M.R., 1975. Morphogenetic analysis: a unifying approach to the s t d y of stromatolites. Abstr. International Symposium on Fossil Algae, Erlangen, Germany. Walter, M.R., 1976. Palaeobiological information from Precambrian and Cambrian stromatolites. Abstr. Int. Geol. Congr., 25th Sess., 1: 33-34. Welke, H.J., Allsopp, H.L., Kostlin, E., Corner, B. and Kroner, A., 1973. Review of age measurements from the Richtersveld Complex and its environs. Abstr. 15th Annu. Congr., Geol. SOC.S. Afr., p. 92. Winnall, N.J., 1971. Some occurrences and aspects of stromatolites in Rhodesia. B S c . Hons. project., Univ. Rhodesia, (unpubl.). Winter, H. de la R., 1963. Algal stromatolites in the sediments of the Ventersdorp System. Trans. Geol. SOC.S. Afr., 65: 115-121. Young, R.B., 1934. A comparison of certain stromatolitic rocks in the Dolomite Series of South Africa with marine algal sediments in the Bahamas. Trans. Geol. SOC.S. Afr., 37: 153-162. Zhuraleva, I.T., 1970. Marine faunas and Lower Cambrian stratigraphy. Am. J. Sci., 269: 4 17-445,

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Section C

SOUTH AMERICA

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835

Chapter 13

STRUCTURAL FRAMEWORK

E. SUSZCZYNSKI

INTRODUCTION

The South American continent comprises three major tectonic environments: (1)the marginal Andean orogenic belt (i.e. the Andean Cordillera); (2) the Braziliaw-Guyanan or South American platform; and (3) the elongated intracontinental depression. The latter represents the tectonic environment between the Andean Cordillera and the South American platform within which a lithic fill of several geological ages accumulated in the Andean “avant-fosse”. The total number of tectonic units within the South American platform is not yet known, but, in a broad sense, two major tectonic domains may be recognized, namely, the Amazonian shield or craton and the Brazilides tectoorogenic belt (Fig. 13.1). Whereas no precise geological data are yet available to permit subdivision of the older Amazonian shield, first identified by Suszczynski (1967), the Brazilides orogenic cycle is well-defined in time and space. The Brazilian-Guyanan platform is regarded as continuous beneath the large transverse aulacogenic structure within which the Amazonian sedimentary basin developed during post-Cambrian times. Similarities in lithologies and structural trends between the Guyana craton t o the north of the Amazon river and the Guapore craton t o the south indicate that tectonic continuity prevailed prior t o the formation of the aulacogenic structure. In the present state of geological knowledge of the South American platform subdivision into: ( 1)deeply eroded and folded crystalline metamorphic basement; (2) folded volcano-sedimentary cover; and (3) unfolded volcano-sedimentary cover, ranging in age from Devonian to Tertiary, constitutes the most reliable classification, but a provisional attempt at further subdivision is made in the following paragraphs.

836

BRAZILIAN PLATFORM MAIN TECTONIC UNITS

-

0

150

300

450

600

750

Kilometres

Limit of the Amazonian Platform Non-Folded Sedimentary Cover (Phanerozoic) Folded Sedimentary Cover (Late Precambrian)

0Metamorphic Crystalline Basement Limit of the Guyanan Orogenic Belt Limit of the Rio Negro-Roraima

Orogenic Belt

E l Limit of the Rondonides Orogenic Belt Direction of the Tectonic Evolution : Brazilides Tecto-orogenic Belt

If-] Synform Fig. 13.1. The main tectonic units of Brazilian-Guyanan platform.

B

R

A

Z

I

L

GEOGRAPHICAL DIVISIONS 0

100 200 300 400 500 600 700 S e O 9OC lOLl0 Kilometres

,

LEGEND

$ 6?

1 Iron Ouadrangle

’

DO SUL L. *.

x vvv

I

3c 3c o

Fig. 13.2.Principal geographic areas of Brazil and adjacent countries showing the loca tions of some important geological areas and features.

2 Jacobina Hiils 3 Caraja’s Plateau 4 Chapada Copper Ore Deposit (INCO-ELUMA) 5 Vitoria da Conquista-ltambe’ - ltapetinga Area 6 Caraiba Copper Ore Deposit 7 Jacurici Vallev Chromite Ore DeDosit 8 Amorindpolis-lpord Sedimentary Uranium Deposits 9 Poxordu 10 Macau 11 Maceid

841 THE AMAZONIAN SHIELD

The Amazonian shield should be regarded as a structural province that was established after the termination of the Middle Precambrian, and that existed before the Brazilides tecto-orogenic cycle. Although it is probable that during the Upper Precambrian the whole of the Amazonian shield acted as a platform to the Brazilides cycle, this concept must be considered to be theoretical at this stage, because only a tectonic discordance is observed between the two structural provinces. All the regional fault systems, tectonic lineaments, elongated graben-like sedimentary basins, volcanic foci with apparently associated kimberlites, and long, narrow superimposed mobile belts are located within a zone bordering the Amazonian shield and the Brazilides. Abrupt changes in structural trends and metamorphic grade are other features that permit separation of these domains. Insufficient geological data militate against the identification of subprovinces within the Amazonian shield. As a provisional statement four tectonic environments are tentatively recognized (Fig. 13.1): (1)The Bacajh-Pan? tectonic block in the northeast. (2) The Rondonides orogenic cycle occupying a roughly n o r t h s o u t h trending area in Rondonia Federal Teritory. (3) The Guyana orogenic cycle encircling the northern side of the Amazonian shield. (4)The Rio Negro-Roraima orogenic cycle (also known as the Negroides or Rio Negro tectonic zone) recognized in western Colombia, southern Venezuela and large areas of Brazil in northwestern Amazonas state and the western parts of the Roraima Federal Territory. The Bacaja-Paru tectonic block This tectonic block, also referred to as the Transamazonic cycle, and the Curua-Xin@ and Paru tectonic zones, is approximately 1600 km long and 900 to 1100 km wide. The Geochronological Research Institute of Sa"o Paulo University report radiometric ages between 2000 and 2200 Ma from this area but amphibolite from the Tapiraph hills has yielded the greatest Amazonian date of 3280 Ma. Taking into account that similar ages of about 3000 Ma have been reported for the crystalline basement at Natividade-Dianopolis in northeastern Goik, it seems to be reasonable to suggest a greater antiquity for the BacajA-Pafi tectonic province, large areas of the western part of which were partly "re-generated" during the Upper Precambrian. Major structural trends within the Bacaji-ParG province are aligned northwestwards. The well-known Caraj5s plateau, built by a major volcano-sedimentary sequence resting with marked discordance on the crystalline basement, occupies the central portion of the Bacaji-Paru tectonic block.

842

The Rondonides orogenic cycle This structural province has a preserved length of 900 km, extending in a northerly or northwesterly direction, and is 500 t o 600 km wide. It is possible that it may extend beneath younger formations into southern Mato Grosso and the Republic of Paraguay, increasing its length to over 1900 km (Fig. 13.2). To the northwest this structural province extends across the Madeira River into Bolivia. Granitic rocks form a belt over 450 km long associated with migmatites extending through the Rondonia Federal Territory and are partly related t o the largest Brazilian tin belt through Jamari, Alto Candeias and SZo Miguel. A sub-parallel, but greatly eroded granite migmatite belt is present in the east-central part of the Rondonides near the Roosevelt and Aripuana" rivers. It is over 600 km long and also has cassiterite mineralization associated with it. Geochronological studies indicate ages between 920 and 970 Ma for the last granitic phases of the Rondonia belt and 1100 to 1450 Ma for the Roosevelt-AripuanZ belt. Regressive metasedimentary sequences overlie the western parts of the Rondonides.

The Guyana orogenic cycle This major province encircles the northern side of the Amazonian shield, extending for 1850 km through Amapa Territory of Brazil and the .three Guyanas into eastern Venezuela. It is possible that this 300 km-wide belt may represent the northern and northwestern continuation of the Brazilides cycle beyond the Amazon aulacogen. The topographically most elevated areas are found in the southern parts of the belt but deeply eroded granitic and migmatitic rocks giving rise to only moderate relief are located towards the Atlantic side. These rocks continue into the high-grade metamorphic and metasomatic rocks associated with granites in northern Brazil. Manganese, iron, graphite and gold deposits are the main known mineral occurrences within this structural province.

The Rio Negro-Roraima

orogenic cycle

The dominant structural trend in this province is northeast which can be seen from ERTS and Radar imageries to be superimposed upon, and to be transverse to the older Precambrian northwest-trend which is the dominant direction in the Amazonian shield. This cycle is younger than the Rondonide and Bacaja-Paru provinces and may possibly be younger than the Guyana orogenic cycle. Mountain chains aligned parallel to the structural grain are prominent along the southern and eastern borders of the province where the highest elevations of the Amazonian shield are found (Pic0 de Neblina, 3014 m). The southern limits of the mountain chains along the frontier between Brazil and Venezuela are terminated by a series of regional faults

843 striking N5OoE, south of which lies the plain of the Amazon basin. The most characteristic feature of this structural province is the pronounced, sub-rectilinear fault pattern along its southern and eastern boundaries. THE BRAZILIDES TECTO-OROGENIC CYCLE

The concept of one major tectonic cycle occupying all the continental areas bordering the Atlantic Ocean and extending: from Uruguay to northeastern Brazil was introduced by the author in 1965 and named the Brazilides by him in 1968. In a preliminary tectonic sketch map of this structural unit, three sub-provinces, namely southeastern, northeastern and eastern, were recognized. Names such as Minas orogenic cycle and Minas-Uruaquano cycle are still used to define tectonic units within the Brazilides but the author considers that their use masks the structural unity of this province and should not be encouraged. Theoretical delimitation of geosynclinal environments led t o the recognition of the Paraguayan, Brasilia, Propri6 and Espinhaqo geosynclines, based on an attempt t o identify subsiding, elongated troughs formed during the Upper Precambrian, and possibly extending into the early Phanerozoic. All these environments are presently considered t o be integral parts of the last stages of evolution of the Brazilides cycle. Except for the Atlantic coast, large areas of Mato Grosso and Goias states were not included in the Brazilides structural province until 1972-1973 when the first metallogenic map of Brazil (Suszczynski, 1973) indicated a probable continuity of these areas with the Atlantic coastal regions. As no elongated belts of granitic rocks exist in Goias and Mato Grosso, such as those typical of the Atlantic coast, the meta-volcanic and meta-sedimentary rocks were interpreted as the roots of a partially preserved geosyncline, to which the name preserved Goianian-Matogrossense geosyncline was applied. This portion of the Brazilides has an elongated sigmoidal shape covering over 3300 km in length and 250 t o 300 km in width. The structural trends within this area are concordant t o those in the previously recognized western area. Red schist belts and large stratiform ultramafic bodies, over 350 km in total length, form the main geological features of this environment. The elongated central Goi& ultramafic belt is considered t o represent a major flexure, rocks within which have given ages from 502 to 627 Ma. As a consequence of the enlargement of the areal extent of the Brazilides by the inclusion of the Goianian-Matogrossense sub-province, it now is in juxtaposition to the Amazonian shield. The Brazilides structural province now extends for 4750 km from Uruguay northwards and has a regular width of 1300 t o 1500 km. In the vicinity of Rio de Janeiro, Minas Gerais and Bahia the Brazilides belt shows a distinct curvature within the central parts of which the Bambui System or SZo Francisco basin, the largest sedimentary remnant in the Brazilides province, is preserved. In the northeastern portion of the Brazilides province, structural trends

844 perpendicular t o the Atlantic coast are developed, the presence of which has led some geologists to propose the existence of another structural province called the Caririan orogeny. Although local tectonic disruption does occur along the lower Sao Francisco river valley and in the Tabatinga hills in southeastern Piaui, no major erosive or tectonic discordances have been recognised by the author, t o support the separation of this northeastern orogeny. A series of sedimentary basins aligned north-south developed prior to tectonism, each successive basin migrating westwards onto the Amazonian shield. Each of the sedimentary sequences in these basins exhibits transgressive characteristics on the east side and regressive ones on the west. Migmatization and granitization processes preferentially affect the eastern portions of the sedimentary sequences with the result that parallel belts of successively younger granitic and migmatitic rocks are developed from east t o west. Preserved between these granitic and migmatitic zones are belts of metasedimentary rocks. Metavolcanic rocks are commonly found on the western sides of every metasedimentary belt. Geochronologic data have revealed ages greater than 2000 Ma for the preserved metasedimentary sequences on the eastern side of the Brazilides province and less than 1000 Ma for non-Sanitized rocks such as in G o i h and Mato Grosso states. In the intervening central area as at Minas Gerais, ages for the “Minas Series” are around 1350 Ma. The granitic and migmatitic belts have yielded ages ranging from 475 Ma t o 720 Ma. In the northeast the youngest granitic events are dated at 475 t o 550 Ma, whereas in the soythern and central parts migmatites which here are predominant have ages between 600 and 720 Ma depending on their geographical situation within the Brazilides province. Volcanicity is more common in the west of the Brazilides, the best example being the “Red Schist Belts’’ of G o i h which extend for over 2000 km in a predominantly northerly direction. (References are listed at the end of Chapter 18).

845

Chapter 1 4

THE GREENSTONE-GRANITIC GNEISS SHIELD E. SUSZCZYNSKI

INTRODUCTION

The greenstone or greenschist belt concept is not used commonly by South American geologists, the French terminology, “roches vertes”, being preferred because of its broader meaning. Two decades ago systematic mapping in the Iron Quadrangle region revealed the presence of a sea of granitic rocks preserving remnants of greenstone belts apparently forming an old nucleus surrounded by young cycles of metasedimentary rocks. An age of 2510 Ma using the K/Ar method, was obtained for the Engenheiro Correa biotite granite of the BacQo Complex that forms part of the basement in the Iron quadrangle and is overlain by the “Minas Series”. Three different types of greenstone and mafic rocks were recognized (Van Dorr, 1963): (1)post-metamorphic gabbroic intrusions; (2) ultramafic intrusive rocks, metamorphosed to serpentinite, talc schist and steatite rock, considered to be almost certainly pre-“Minas Series” in age ; (3) metamorphosed mafic to ultramafic intrusions, probably post-“Minas Series” in age and represented by greenstone and chlorite schists. De Mascarenhas (1976) examined the east-central portion of Bahia state and reported the presence of greenstone belts associated with granitic and migmatite rocks that belong t o the “Granulite Complex of Jequih” with ages ranging from 2600 to 2800 Ma. This author recognized three lithologic groups within the greenstone belts, namely, ultramafic, greenstone and sedimentary. “Serpentine” belts and ultrabasic complexes have been recognized, the origin of these rocks being attributed to early episodes of ophiolitic magmatism, related t o the beginning of geosyncline formation (de Almeida, 1967) or to “Alpine”-type magmatic activity, when they are found as fault-controlled, isolated or discontinuous massifs. More detailed mapping has been or is being undertaken in Rio Grande do Sul state (Goni, 1962) and in the Goiis region of east-central Brazil to obtain a better understanding of the

846

serpentine belts and large stratiform ultramafic bodies. In Goi6s the ultramafic rocks appear to be cross-cut by a regional system of tear faults producing elongated tectonic blocks with a westward-dipping imbricate structure surrounded by the geologically younger Volcanogenic Schist Series. According to the author’s studies these massifs were derived from primitive metasedimentary, pure and impure thick carbonatic sequences, because they show no evidence of old ultrabasic submarine volcanism. The Amazonian greenstone belts in southwestern Amap6, middle Parh River, western Roraima and west-central Rondonia are imperfectly understood and, as a consequence, those located in the eastern portion of the Brazilian platform will be discussed in this chapter. Considering only their lithological nature, several types of greenstones can be recognized in the Brazilian shield: (a) amphibolitic belts of metasedimentary origin; (b) hybrid or impure amphibolite units developed on top of the Volcanogenic Schist Series; (c) chloritic-talc schist and amphibolitic association; (d) serpentine belts; (e) red schist belts; (f) pyroxenite belts; (g) fault-controlled ultramafic bodies; (h) large stratiform ultramafic bodies; (i) greenstone belts, undivided. The three main geological environments in which greenstones are believed to have developed include: (1)submarine ultramafic volcanism constituting the ophiolite suite; (2) volcano-sedimentary ; and (3) metasedimentary. Greenstones of metasedimentary origin have distinctive chemical characteristics; those at the top of transgressive marine sedimentary sequences have a special calcium-iron composition whereas those in the lower parts of regressive sedimentary sequences are magnesian-rich. Although there are undoubted examples of ophiolite suites in the Brazilian Precambrian, many others formerly included in this classification are regarded now as products of high-grade metamorphism of original carbonate facies in regressive sedimentary sequences. Previous workers have frequently made no distinction between the small, isolated ultramafic massifs or the elongated stratiform bodies and the ultrabasic rocks of meta-sedimentary origin. Amphibolite rocks are very characteristic of the Brazilian Precambrian and amphibolite belts are common. The recognition of the nature and origin of individual amphibolites is an important factor in mineral exploration. Many workers classify all of them as metabasites but the amphibolites are of both pure metasedimentary and impure or hybrid origins.

847 It is important t o note that no unequivocal evidence favouring a volcanic origin can be found for many greenstones in the Brazilian Precambrian. This writer believes that this lithology is derived commonly from magnesian carbonatic facies developed in the lower stratigraphic levels of regressive marine sedimentary cycles. The questionable status of these and the uncertainty surrounding the so-called ophiolite suites implies that types (1)and (2) above do not represent examples of classic greenstone belts such as those found on other continents. However, the most common environment in which greenstones are developed in the Brazilian Precambrian is the volcano-sedimentary environment similar to those in Africa and Australia: (a) the Red Schist belts, found particularly overlying the western parts of the Brazilides tectonic province; and (b) the green schist belts found principally on the eastern parts of the Brazilides tectonic province. The latter include ultramafic sequences that constitute the lowermost stratigraphic levels of thick volcano-sedimentary piles. Good examples of such sequences are found mainly near the Atlantic coast in several areas in Espirito Santo and Sergipe as well as in the southern parts of Brazil in Rio Grande do Sul, Santa Catarina and ParanL. Greenstone belts have been described from eastern Bahia in the Patamute-Uauii belt (Suszczynski, 1972 (de Mascarenand 1975) and around Santa Luz-Queimadas-Uaui-Serrinha has, 1976). In the Goiiis region the development of submarine ultramafic volcanics at the base of volcano-sedimentary piles is much reduced. Geochronological data indicate that the greenstone belts near the Atlantic coast are over 2400 Ma old, whereas the Red schist belts of Goiiis have given ages of around 600 to 700 Ma. These younger schist belts are regarded as having developed during the closing stages of the Brazilides tectonoorogenic cycle. Some ages of over 3000 Ma have been reported from sequences near the Atlantic coast, suggesting development of greenstone belts from early Archaean to late Proterozoic and migrating from east t o west with time. RED SCHIST BELTS, GOIAS AND SURROUNDING AREAS

Large areas in the Goiis region and adjacent parts of western Minas Gerais are underlain by distinctive volcano-sedimentary sequences, deposited predominantly in a marine environment. Many geologists classify these metamorphosed rocks as belonging principally t o the Arax6 Series which this writer considers to be inappropriate on the grounds that a variety of different lithologies, reflecting changing depositional environments, are present . The base of the sequence is rarely exposed but in central Goi6s between Inhumus and Itauqu it is represented by sediments finely intercalated with rocks of magmatic origin, some tens of metres thick. Overlying this basal

848

group is a volcano-sedimentary succession that is overlain by the thick Volcanogenic Schist Series, the characteristic feature of which is its reddish colour that contrasts with the green or grey hues of the underlying rocks. Among the most characteristic lithologies of the Volcanogenic Schist Series are: (1)The presence of several thin layers of amphibolite of which three types can be recognized. At the base the amphibolites are of clear volcanic origin but include minor intercalations of sedimentary material. At intermediate stratigraphic levels the amphibolite is described as hybrid reflecting an intermixture of amphibolites of both metavolcanic and metasedimentary origins. Amphibolites at higher stratigraphic levels are entirely of pure metasedimentary origin. (2) Pure quartz or micaceous, sometimes garnetiferous and kyanite-bearing, saccharoidal t o glassy quartzitic formations that represent deposition under regressive conditions. (3) Some intermediate to acidic volcanic beds at the top of the Volcanogenic Schist Series are associated with pyroclastic material including ash beds. These volcanics were apparently extruded under continental conditions and have the characteristic red or purple hues from which the term Red Schist belt is derived. Kaolin-rich layers are common, and metamorphism has given rise to quartz-muscovite-feldspar segregation veins and dykes, as may be seen at southern Goias, north of Pires do Rio. The Volcanogenic Schist Series comprises 60% of the total thickness of the volcano-sedimentary pile and 10 to 15% is composed of lithologic units that have the appearance of greenstones, only the lowermost of which being derived from basaltic rocks. GREENSTONE BELTS ADJACENT T O T H E ATLANTIC COAST

Ultrabasic rock types are important in this environment, being represented by pyroxenites, talc schists, peridotites, amphibolites and chloritic schists. Biotite-rich layers higher in the stratigraphic pile are overlain by thick amphibolites or micaceous amphibolites and mica gneisses, schists and metaarkosic sequences. Glassy orthoquartzites associated with massive magnetiterich layers are found commonly in eastern Bahia and in the south-central part of Minas Gerais. Amphibolites are well-represented, hornblendite and pyroxenite being common in Bahia. Towards the west in Minas Gerais hornblendite is less abundant being replaced by tremolitic and actinolitic schists with local intercalations of chlorite and talcose schists. SIALIC PLUTONISM, METAMORPHISM AND STRUCTURE

Granitic rocks are best developed near the central parts of migmatized schist series or very close to the lowest stratigraphic levels of greenstone

849

belts. Migmatization has affected only the volcano-sedimentary sequences towards the Atlantic coast. In the east-central parts of Bahia deep erosion exposes large areas of migmatites within which are preserved small remnants of greenstones (De Mascarenhas, 1976). Westward of central Minas Gerais volcanogenic schists and greenstones intercalated with them did not suffer any migmatization. Granitic intrusions are found in east-central Bahia, for example, the Cansancao, Santa Luz and Conceiqao de Coit6 areas. The Baqiio granitic-migmatitic complex in the central part of the Iron Quadrangle, particularly its Engenheiro Correa facies, may be also considered as the product of partial granitization of the lower levels of a thick volcanogenic schist sequence of pre-Minas age. In Rio Grande do Sul state, north of Bag6 and in central Bahia between Jequi6 and Milagres migrnatization has affected thick volcanogenic schist sequences. N o extensive migmatization is observed in the GoiLs area, but nevertheless small granitized areas are reported, giving rise to sub-alkaline granites, that occur in the cores of small domical structures as, for example, east of Damolandia in central Goifis. This particular granitic cupola is 2 to 3 km in diameter. The grade of regional metamorphism decreases from east to west so that the volcano-sedimentary piles in Goias have been affected by only low-grade epizonal metamorphism without the development of strong schistosity. In the stratigraphically higher levels original sedimentary and volcanic textures are preserved, emphasizing the low grade of metamorphism that has affected these rocks. The volcano-sedimentary sequences and associated greenstones are mainly confined to structurally controlled depressions occupying parallel belts, for example, in the Rio Pardo and eastern Minas Gerais areas. In Minas Gerais these alternate belts are aligned n o r t h s o u t h with widths of 30 to 70 km and hundreds of kilometres in length. Individual belts are separated by migmatitic or granitic terranes. Westward from central Minas Gerais to Goi6s, the volcano-sedimentary piles build low plateaux resulting from block faulting. MINERALIZATION

The Red Schist belts are largely devoid of mineralization. Thin beds of massive magnetite or titanomagnetite are found intercalated in the stratigraphically lower levels in central Goias, and towards the north sub-economic graphitic layers are developed. Although the amphibolites contain scattered and disseminated pyrite with minor pyrrhotite, no sulphide orebodies have been located. At Passa Tempo-Desterro-Entre Rios, near the Iron Quadrangle, three massive hematite beds are intercalated in the Volcanogenic Schist Series. Manganese occurrences are found between Goi6s and west-central Minas

850 Gerais as, for example, at Conselheiro Lafaiete. The presence of iron and manganese characterizes the older greenstone belts. The Ubaira-Santa In& magnetite deposit is associated with glassy, quartzites. The manganese deposits at MaraG, Coaraci and Santo Antonio de Jesus are typical examples of such deposits in eastern Bahia. The high-magnesium greenstones that the writer believes t o be derived from original magnesium-rich carbonate sediments are associated with deposits of chromite, vermiculite and asbestos, but the writer does not preclude in some few cases the possibility that similar mineralization may also be associated with magmatic ultrabasic intrusions. However, it is believed that these latter rocks are much less widely developed than previously supposed. Phosphatic minerals are found in rocks of pyroxenitic composition such as at Rio do Peixe-Ipir6 and Itambk. Massive magnetite and titanium-rich minerals are associated with apatite in this environment. Nickel and copper sulphides in economic concentrations as at the Caraiba mine in east-central Bahia are found as disseminations, stringers and small massive lenses within ultramafic stratified rocks metamorphosed at granulite grade, but nevertheless in close relationship with old metasedimentary sequences. (References are listed at the end of Chapter 18).

851

Chapter 15

THE CRATONIC ENVIRONMENT E. SUSZCZYNSKI

INTRODUCTION

The crystalline basement in both the Amazonian and Brazilides structural provinces is overlain by sedimentary and volcanic cover sequences, ranging in age from about 2000 Ma to Phanerozoic, that are now confined t o a great number of isolated, remnant basins. The main areas covered by these rocks in the Amazonian shield are: (1)Roraima plateau, and extending through northern Brazil into Venezuela, Guyana and Colombia (locality 31, Fig. 15.1). (2) Cubencranqubm-Gorotire plateau covering the mid-Xing6 River (Fig. 13.2). ( 3 ) Tapaj6s basin and Cachimbo plateau covering a large area south of the Amazonian aulacogen (localities 26', 26'' and 35, Fig. 15.1) and Prosperanca north of the Amazon River (locality 29, Fig. 15.1). (4)Mapuera-Uatuma area north of the Amazon River (locality 36, Fig. 15.1). (5) Scattered outcrops of the Palmeiral Formation and other sedimentary sequences in Rondonia (locality 30, Fig. 15.1). (6) Caiabis-Apiacas, Serra Formosa-Minissau6 and Rio Liberdade basins in northern Mato Grosso (localities 25, 24 and 23 respectively, Fig. 15.1). (7) Carajas plateau in Para state (locality 26, Fig. 15.1). (8) Amazonian aulacogen covering the Amazon basin (locality 28, Fig. 15.1). The age of the cover sequences is often uncertain, but those of the Roraima plateau have a minimum age of about 2000 Ma based on dating of intrusive dolerite dykes and sills (McDougal and Compston, 1963). A minimum age of about 920 Ma is indicated for the cover rocks in Rondonia on the basis of their approximate contemporaneity with granitic ring dykes and rhyolites of that age. The age of the sediments in the Cubencranqubm and Gorotire plateau in the mid-Xingii River, is older than 1350 Ma. Lithologic similarities between these rocks and those in the Caiabis-Apiacas, Serra

852

Formosa-MinissauA and Rio Liberdade basins leads t o the provisional conclusion that the latter are of a similar age. An Upper Precambrian age has been proposed for the cover rocks on the Caraj6s plateau. The folded cover rocks in the Mapuera-Uatuma area have yielded fossils of Lower Silurian age, and younger sequences placed in the Lower Devonian. The Amazonian aulacogen has a lithic fill that accumulated during Ordovician and Silurian times. The Roraima cover rocks have been subdivided into three formations from the base up consisting of:

853 (1)Kaieteur Formation has a thickness in excess of 500 m and consists of conglomerate and sandstone, from which detrital diamonds have been reported in lensoid mglomerate beds. ( 2 ) Surumu Formation consists of rhyolites and ignimbrites with a thickness greater than 1000 m. (3) R o r i m a Formation, 900 m thick, composed of arkosic sandstones with minor intercalations of conglomerate, argillaceous schists, jaspilites and basic lavas. Sheets of gabbro and norite, several hundred metres thick, together with dolerite dykes and sills intrude the cover sequence causing it t o be domed and folded. On the Cubencranqukm and Gorotire plateaux Barbosa (1966) described from the base to the top the following formations: Rio Fresco, Gorotire, and Cubencranqukm. Subsequently the author has changed the stratigraphic order as follows from the base upwards: (1) Gorotire Formation consists of conglomeratic sandstones that are micaceous and koalinized with some intercalated blue and black marine shales. (2) Rio Fresco Formation is a volcano-sedimentary sequence with carbonaceous beds, micaceous arkosic sandstones and andesites intruded by dolerite. ( 3) Cubencranqukm Formation comprises sandstones and fine-grained arkoses with lenses of chert and jaspilite.

Fig. 15.1. The structural framework of Brazil. Numbers in circles refer to the following areas: 1 = Santa Barbara; 2 = Camaqua; 20 = Upper Balsas; 3 = Itajai; 21 = Tocantins basin; 4 = CampoLargo; 22 = Paraguayan-Araguaian; 5 = Gurutubinha; 23 = Rio Liberdade; 6 = Vale d o Ribeira, including 24 = Serra Formose-Minicua; 25 = ApiacasCaiabis; Cerro Azul-Malo Preto; 7 = Castro-Iapb; 26l = Cachimbo plateau; 8 = EleutBrio; 2611 = Old Tapajbs basin; 9 = Carandai-Veloso; 26 = Carajas plateau; 27 = Jamanxim Area; 10 = Rio CuitC-Tumirintinga; 28 = Amazonian aulacogen; 11 = “Bambui System’‘ 29 = Prosperanca; -SZo Francisco basin; 30 = “Palmeiral Formation” 12 = Jacar&-Lenqois; and others; 13 = Rio Pardo; 31 = Roraima; 14 = Miaba-Canudos ; 32 = Serra da Borracha; 1 5 = Iara ; 33 = Itaberai-Rio Bonito; 16 = Cococi; 17 = Jaibaras; 34 = Estlncia ; 35 = Sucunduri area; i a = Rio Paraim; 1 9 = Pindorama; 36 = Mapuera-VatumT Area.

854 Certain of the sedimentary features have been interpreted as the consequence of deposition in a prograding deltaic environment. Further south, in Mato Grosso state the preserved cover rocks are lithologically similar t o the Cubencranqukm Formation with which a correlation is proposed. The Cachimbo plateau is formed by sandstone sequences with basal conglomerates, locally containing placer gold. In the Tapaj6s basin two formations, called the Juruena-Teles Pires and Middle Tapajbs, have been recognized, each consisting of sediments reflecting transgressive conditions followed by regression. No volcanic rocks are known but the sedimentary rocks are cut by basic dykes and sills. In the mid-Aripuaiia valley, de Almeida (1959) described the Beneficente Group consisting of orthoquartzites with basal conglomerate lenses. The sequence is distinguished by the presence of rhyolites and ignimbrites. The Caraj6s plateau is built by the Carajas series consisting of conglomerates, quartzites, phyllites, manganiferous slates, mica schists and basaltic volcanic rocks overlain by banded siliceous rocks with high-grade iron-ore concentrations. These later have been described as itabirites but the author considers that they owe their origin t o differentiation within the magma chamber that was the source of the early basaltic flows coupled with later deuteric and telethermal processes. The Palmeiral Formation and other sedimentary remnants in Rondonia consist of arkosic sandstone, locally conglomeratic, that are red in colour. Conglomerates are also present, that may contain fossil cassiterite placers. Manganese and iron ore are associated with the arkosic facies. Carbonate sequences are found in the Sucunduri area (locality 35, Fig. 15.1), and along the Tapaj6s River. At the last named locality these rocks constitute the Acari Group, which is over 350 m thick. Stromatolite structures have been discovered in the carbonate sequence known as the Sucunduri Group. Two types of preserved basins of cover rocks can be distinguished within the Brazilides domain. The first and most prominent type are the elongated basins in which subsidence occurred, and the second type are small tectonic basins with the lithic fill preserved in old graben structures. The principal basins of the first type are: (1)The SZo Francisco basin which contains the Bambui System, the maximum age of which is 1350 Ma, but an age of 600 Ma is suggested (Cordani, 1968) (locality 11, Fig. 15.1). (2) The Tocantins basin lying close t o the boundary between the Amazonian shield and the Brazilides structural province. The basin is filled mainly by the Tocantins Series. It is open northward (locality 21, Fig. 15.1). (3) The Paraguayan-Araguaian basin whose lithic fill is subdivided into the Upper Paraguay, Caacupk, Araras and Bodoquena groups. I t is open southward (locality 22, Fig. 15.1).

855 (4) The Lenqois-Jacar6 basin located in north-central Bahia which is regarded as a portion of the Sgo Francisco basin, from which separation occurred during the closing stages of sedimentation. Several lithologic groups such as the Paraguaqu, Tombador and Morro do Chap& have been identified in this basin. There are twenty of the second type of structurally preserved basins: (1) Camaqua and Santa Barbara basins, distinguished by the presence of cupriferous ores (localities 1 and 2, Fig. 15.1). (2) Itajai basin filled by clastic sediments and volcanics (locality 3, Fig. 15.1). ( 3 ) Castro-Iap6 basin filled by clastic sediments and volcanics (locality 7, Fig. 15.1). (4) Vale d o Ribeira in which low-grade metamorphism is superimposed on a volcano-sedimentary sequence (locality 6 , Fig. 15.1). (5) Rio Bonito-Itaberai basin filled by a volcano-sedimentary sequence (locality 33, Fig. 15.1). (6) Carandai-Barroso basin that contains limestone and black shales (locality 9, Fig. 15.1). ( 7 ) Rio Pardo basin in which conglomerates contain detrital diamonds and the early Palaeozoic sediments include evaporite deposits (locality 13, Fig. 15.1). (8) Estgncia formation which contains red beds (locality 34, Fig. 15.1). (9) Miaba-Canudos and Itabaiana groups of low metamorphic grade (locality 14, Fig. 15.1). (10) Serra da Borracha (locality 32, Fig. 15.1). (11)Rio Paraim basin containing limestones and black shale (locality 18, Fig. 15.1). (12) Cococi basin of clastic sediments and volcanics (locality 16, Fig. 15.1). (13) Jaibaras basin consisting of clastic and carbonate sediments, black shales and volcanics. Stratiform copper deposits are associated with the sediments (locality 17, Fig. 15.1). (14) Pindorama basin (locality 19, Fig. 15.1). The following small structurally preserved basins comprise clastic sediments with or without intercalated volcanics: (15) Campo Largo (locality 4 , Fig. 15.1). (16) Guaratubinha (locality 5, Fig. 15.1). (17) Eleutkrio (locality 8, Fig. 15.1). (18) Rio Cuit6-Tumiritinga (locality 10, Fig. 15.1). (19) Iara (locality 15, Fig. 15.1). (20) Upper Balsas area (locality 20, Fig. 15.1). It is difficult to be precise about the absolute ages of isolated basins containing similar lithologies. It seems that the western remnants are younger in age than those t o the east. Thus the Paraguayan-Araguian basin is younger

856 than the Bambui System in the Sgo Francisco basin. A further difficulty arises from the fact that close lithological similarities exist between the Bambui System in the SZo Francisco basin and the Araras, Corumba and Bodoquena Groups in the Paraguay-Araguaian basin. An Eocambrian age is regarded as probable for certain of the coarse clastic sedimentary formations, some of which were deposited under palaeodeltaic conditions. Fundamental differences can be distinguished between the nature of the lithic fill of the basins resting on the Amazonian shield and those within the Brazilides structural province. The characteristic feature of the folded sedimentary cover rocks in the Brazilides structural province is the widespread presence of carbonate rocks that are scarce in the sedimentary cover sequences resting on the Amazonian shield. In the Jacar&-Lenqois basin carbonate sequences are over 1300 m thick. The cover rocks on the Amazonian shield are characterized by the presence of thick volcanic flows, usually rhyolitic in composition but also including dacites and andesites. Ignimbrites and granophyres are also present and several types of granitic stocks intrude the cover. In contrast volcanic rocks play only a minor role in the cover sequences within the Brazilides province and have a restricted development when they are present. With the exception of the cover rocks in the UatumZ-Trombetas area which are confined t o a tectonic depression within the Amazonian shield all the remaining cover sequences build prominent plateaux, bounded by steep scarps, that rest with a clear unconformable relationship on the crystalline basement. The folded cover rocks in the Brazilides province are found principally along regonal tectonic axes. Nearly all the outcrops of cover rocks are confined to the area of tectonic downwarping known as the P&ran&Sa"o Francisco-Parnaiba axis which has a general north-northeasterly or northeasterly trend. Similarly the cover rocks in the Paraguay-Araguaian and Tocantin basins occupy clearly defined structural depressions, the latter having a general northward strike. There are significant age differences between the sedimentary cover of the two structural provinces, those resting on the Amazonian shield being older. The volcano-sedimentary sequences forming the folded cover to the Amazonian shield appear to represent at least four cycles, starting with an initial period of marine transgression and terminating with regression during which sandstones and conglomerates were deposited. The final stages of regression were marked by periods of continental volcanism, there being no evidence of submarine volcanics. Attention must be drawn to the lithologic similarities between the Bambui System or the Araras and Bodoquena Groups with the Katanga System of central Africa which is emphasized by the occurrence of stratiform copper mineralization in some areas. In Minas Gerais, the following succession has been recognized from the base upwards, forming collectively the classic "Minas Series":

857 (1) Caraqa Group: with the Moeda Quartzitic Formation, about 200 m thick, rests on the crystalline basement that is invaded by granites at -1350 Ma and by dykes at -1110 Ma; and the Batatal Formation of slates with minor cherts, about 150 m thick. (2) Itabira Group consisting of the CauE Formation with banded ironstones, and the Gandarela Formation with dolomites, about 300 m thick. (3) Piracicaba Group, 3000 m thick, comprising several geological formations with quartzites, ferruginous quartzites, pelites and dolomites. (4)Itacolomi Group overlies the “Minas Series” unconformably and consists of cross-bedded quartzites with conglomerates, phyllites and an upper cross-bedded quartzite. It is in this series that the flexible sandstone, itacolomite, is found. A palaeodeltaic regressive environment is here clearly represented, with some scattered gold occurrences. The crystalline basement and older cover sequences are overlain in both the Amazonian and Brazilides provinces by largely unfolded sediments and volcanics of Phanerozoic age. In the Jacobina area, thin-bedded and cross-bedded quartzites, containing about 27 auriferous conglomerates and cross-bedded deltaic quartzites overlie micaceous quartzites, schists and phyllites. The conglomerates that consist of white quartz pebbles set in an arenaceous matrix are largely barren and low gold values are associated with pyrite-bearing conglomerates. There are many similarities between the Jacobina Series and the Witwatersrand succession of South Africa. The Jacobina Series is considered t o be Early Proterozoic in age, and it has a net .regressive metasedimentary character. Only regressive clastic sequences have developed gold-pyrite mineralization in the Brazilian platform. Primitive transgressive sequences are barren. STRUCTURE

Although the preserved cover sequences on the Amazonian shield are found now as isolated fragments it is possible t o re-construct their probable original distribution and identify three major structural trends south of the Amazon River. The oldest one is bounded by the upper X i n d and the mid-TapajGs rivers and includes the sedimentary-volcanic sequences on the Cubencranqu6m, Gorotire, and CarajAs plateaux, where volcanic rocks are well-preserved. T‘ne oldest rocks are found along the Xing% River. A prograding deltaic sequence at the end of a major sedimentary cycle is represented by the sediments on the Cubencranqukm-Gorotire plateau. One hundred kilometres south of this plateau, marine sediments are preserved but imperfectly known in the Liberdade-Minissaui area (localities 23 and 24, Fig. 15.1), extending into northern Mato Grosso state. The dominant structural trend is northwest-southeast.

858 Post-dating this trend is one that affects the remnants of lower rocks in Rondonia. It is possible that old sediments preserved in the Caiabis basin (locality 25, Fig. 15.1) also belong t o this event, which, if proved correct, would extend considerably the area affected. The main characteristic of this group of rocks is their consistent westward dip so that the youngest members of the stratigraphic column are found along the western border of Rondonia Federal Territory. Superimposed on these earlier trends is the west-northwesterly aligned Cachimbo-Tapaj6sProsperanca axis (localities 26l, 2611and 29, Fig. 15.1) the cover sequence associated with which rests with strong discordance on older rocks involved in the earlier events. The sediments contained within this structure have been subdivided into two series, both consisting of regressive and transgressive sequences. The southernmost comprises the JuruemaTeles Pires series and the younger, Middle Tapaj6s series lies towards the northwest. The Apiachs sedimentary sequence (locality 25, Fig. 15.1) represents an isolated southeasterly extension of this event, which is represented north of the Amazon aulacogen by the outcrops in the Prosperanqa area. The main structural expression is similar to a megasynclinorium plunging to the west-north west. North of the Amazon River one of the most prominent structural trends is that which involves the rocks of the Roraima plateau. In this area clastic sedimentary and volcanic sequences present arcuate structures with northwesterly directed plunges on the southern side of the Amazon aulacogen. The author suggests that the Roraima cover rocks young t o the south. If this interpretation is correct it implies that the central core or nucleus of the Amazonian shield was the site of successive volcano-sedimentary depositories the depositional axes of which migrated radially away from the central nucleus. The major feature of the Amazonian shield is the early Phanerozoic Amazonian aulacogen, trending N60" E and cutting transversely across all the Precambrian structures. Significant geophysical anomalies, attributed to the presence of ultrabasic rocks, have been found along the central axis of this megastructure in addition to large basic intrusions of late Mesozoic age. The structural patterns within the Brazilides province are quite different from those just described for the Amazonian shield. The identification of continuous structural trends associated with basins of preserved sedimentaryvolcanic sequences, aligned generally northwards, can be made readily in the Brazilides province. The preserved basins now have the form of synclinoria that consistently plunge northwards throughout the Brazilides domain. The lithologic sequences in each preserved basin reflect an upward coarsening along their western flanks that is interpreted as the response to regression. In the larger basins regressive cycles are commonly terminated by volcanism. The sediments accumulated in a series of basins, the depositional axes of which migrated towards the west with the result that lithic sequences found

in the east are older than those in the west. The most extensive outcrops of the cover rocks are found in the Siio Francisco basin where there is arcuation of the Brazilides trend, concave to the east. Like all the other preserved basins the SZo Francisco basin was originally open to the west. The absence of carbonates in certain of the smaller preserved basins is attributed to the lack of preservation of marine environments during regressive phases, which each time finished in continental environments. The similarity in lithologic and structural features in successive basins which migrated from east to west emphasizes the essential tectonic unity that prevailed during the evolution of the Brazilides tectono-orogenic cycle. MINERALIZATION

Some Brazilides manganese ores may be similar to those found in India. The classic primary manganese protores of Minais Gerais (for example, that from Conselheiro Lafaiete) were developed within volcano-sedimentary sequences spatially associated with amphibolitic-garnetiferous and calc-silicate horizons. To the west towards the G o i h region the manganese deposits are exclusively associated with thin amphibolitic zones on top of the Volcanogenic Schist Series, i.e. the so-called Red Schist belts. Specific carbonatic manganese mineralization is found only in some facies of the Bambui System, near the Iron Quadrangle, in Minais Gerais state. Toward the top of the Bambui System in the Goi6s region the manganiferous sedimentary facies grades into slate, quartzite and metaconglomerate, strongly developed at the top of the regressive sequences, as in the case of the manganiferous area around SZo JoZo da Alianqa. Near the Atlantic coast in Bahia, Espirito Santo and Rio de Janeiro the manganese protore is associated with highgrade metamorphic rocks (leptites and anorthosites) derived from impure metasedimentary sequences. The itabirites are found with rocks that were originally limestones and dolomites, calc-pelitic rocks and chert-quartzites giving rise to three types of itabirites, namely, dolomitic, amphibolitic and quartzitic. All of these lithologies were formed under regressive sedimentary conditions. No itabirites related to transgressive sedimentary sequences have been found in the Brazilian platform. Another type of manganese mineralization is that represented by iiie mixed iron-manganese ore found especially in the Iron Quadrangle. This type, predominantly uneconomic, is developed in fine detrital metasedimentary sequences which cover the lower, most iron-rich sequences formed by chemical precipitation. The classic iron-manganese association in the Iron Quadrangle represents one of the branches of the well-known “Jacutinga” gold mineralization. In this case impure, manganese-bearing strata are developed within ferruginous quartzite. Graphite at low stratigraphic levels,

860

manganese at intermediate levels and iron in the higher stratigraphic levels form a very common metallogenetic association in Precambrian areas of the Brazilian Shield. The Brazilian iron-ore deposits are located in several kinds of metasedimentary series where thin layers of haematite are intercalated rhythmically with quartzite of which itabirite is the classic type. The “Minas Series”, now subdivided by the author into several other “groups” with similar lithologic characteristics due t o the cyclic nature of the old marine environments, in the Iron Quadrangle, is well-known for this type of deposit. They are also found in many other areas on the Atlantic side of the Brazilian shield. A correlation with the iron ore deposits in the Transvaal Supergroup of South Africa may be possible. Within the Amazonian shield important deposits of manganese and iron have been discovered since 1967 on the Carajas plateau. Supergene enrichment of siliceous, magnetite-rich protore related directly t o old basic and hybrid volcanism, has resulted in the formation of large and economically important deposits in this area. Manganiferous slates and shales stratigraphically below the iron protore have been prospected with success in recent years. Large amounts of primary, banded magnetite and titaniferous magnetite ore in basic volcanic rocks have not yet been taken into consideration in the Carajas plateau. The gold deposits of the Jacobina district are similar to those of the Witwatersrand. Gold placers associated with pyritic conglomerates have been mined from the lower stratigraphic levels of a thick, clastic, regressive sequence which in the west overlie various schists and micaceous quartzites. In GoiAs, Mato Grosso, Iron Quadrangle where similar geologic environments are found further gold mineralization is known. Recent discoveries of economic lead and zinc sulphide ores have been made in the western outcrops of the Bambui System in the Vazante-Morro Agudo-Paracutu area. The mineralization is associated with brecciated calcarenites in the lowest stratigraphic levels of a 350-400 m-thick dolomite sequence which is overlain by phillites, slates and quartzites reflecting a low metamorphic grade, representing a classic lithologic profile of a regressive metasedimentary environment open toward the west. The many significant features of mineralization in sedimentary-volcanic sequences in Brazil indicate their association with regressive mega-cycles of sedimentation. It must be recognized, however, that detailed studies and exploration of transgressive cycles are rare because they are almost completely unknown within the Brazilian Precambrian. In very few known cases only poor, uneconomic and scattered mineralization is found in this environment which still needs more studies t o confirm that it is in fact a primitive transgressive sedimentary environment. It is important t o emphasize this point because it leads to a better understanding of metallogenesis in the Brazilian platform. This part of the Earth’s crust is almost entirely formed by

861 rocks types characteristic of regressive, metasedimentary sequences, and associated volcanism. In addition to authigenic placer gold deposits derived from the precipitation of colloidal gold under specific, Recent conditions with pH lower than 5 . 3 , mainly in the Amazonian areas, detrital tin is also mined in Rondonia and other regions. Placer tin and gold deposits are also known on the Amazonian shield north of the Amazon River in AmapA state but these deposits have been superseded in importance by the recent discovery of extensive cassiterite placers in northeastern Goi6s. (References are listed at the end of Chapter 18).

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863

Chapter 16 ZONES OF REACTIVATION E. SUSZCZYNSKI

Reworking of older crust is a feature of the South American Precambrian, the most clearly understood area constituting the Brazilides tectono-orogenic domain. On the African continent, the Mozambiquian, West-African and Madagascan rejuvenation belts were identified to account for the concentrations of radiometric dates, around 475 t o 550 Ma, that were reported from wide geographical areas. These ages represent a period of reactivation of older crustal material. In 1967, a geochronological programme, conducted in support of the second international Symposium on Granitic Rocks, led to the recognition in northeastern Brazil of a similar concentration of dates between 450 and 550 Ma. Soon afterwards a direct correlation was established between this area of Brazil and those on the West-African continent, thereby drawing attention t o the possible existence of a widespread mobile belt affecting both continents contemporaneously. In northeastern Brazil, four parallel granitic chains separated by belts of preserved metasedimentary rocks reflect an evolutionary sequence of events typical of other areas within the Brazilides structural province. Radiometric ages of the granitic rocks decrease from east to west and it has been shown subsequently that ages decrease from south to north. The intervening belts of unmigmatized, preserved metasedimentary belts have yielded radiometric ages older than 750 Ma and, locally, greater than 1500 Ma. The Jabitaca basic metamorphic rock in the central areas of Pemambuco has been dated at over 1800 Ma. The well-established concentration of ages of 450 to 550 Ma in northeastern Brazil was considered t o be evidence for a distinct Caririan orogeny affecting this part of the Brazilian shield. In the author’s view there is no evidence to support the recognition of a distinct and discrete orogenic event. This area is regarded as an integral part of the Brazilides domain, although it is admitted that granitic rocks are more extensively developed in the northeast. Within the last decade, additional geochronologic data, predominantly

864 from migmatitic areas near the Atlantic coast in the central-southern and southern parts of Brazil, have yielded ages of between 620 and 750 Ma. These areas are still regarded by some geologists as forming part of an older and separate domain, termed the Paraibides tectonic belt. The author regards these areas as the southern tectonic domain of the major, widespread Brazilides structural province. The relatively older dates for the southern and eastern areas are interpreted as the earliest events of the Brazilides cycle, metamorphism, migmatization and the intrusion of granites being initiated at an earlier time and affecting old meta-sedimentary sequences than those in the northeast. In a similar manner, radiometric dates decrease towards the west, implying that the evolution of the Brazilides tectono-orogenic event involved a migration northwards and westwards reworking progressively younger cover sequences as it did so. In Bahia ages of 3150 Ma have been recorded along the Atlantic coast. Rocks of similar lithologic type have ages less than 2500 Ma farther westwards. The typical sequence of events in the reactivation of the older crust involves initial regional metamorphism, locally of granulite grade, followed by migmatization and finally t o the production of various granites. These three major successive stages in the process of reactivation can be distinguished geochronologically. The granitic rocks yield the youngest ages, whereas the migmatites give intermediate ages and the preserved cover sequences affected by regional metamorphism only are the oldest. Migmatization and granitization are most intense in the lower stratigraphic levels and decrease upwards resulting in the preservation of regionally metamorphosed volcano-sedimentary remnants. Accompanying this stratigraphic control is the progressive decrease in intensity of metamorphism, migmatization and granitization in successive belts from east t o west. This feature makes the correlation of the cover sequences from belt t o belt not so difficult, particularly because similar lithologies are also repeated in time as well as space. The Brazilides domain displays distinct zonation with more intense reworking in the east than the west. Similarly there is a geochronologic zonation that is reflected by the younger ages of granitic rocks formed during the terminal phase of reactivation. Thus younger ages are found towards the west and northeast, presumably reflecting the development and growth of this tectono-orogenic cycle as it extended onto the older shield. Reactivation zones within the Amazonian shield are imperfectly known and understood, largely due t o the thick zone of weathering and dense jungle that covers so much of this area. One of the better-known areas is that between the Iriri and Curu5 rivers (300 km southeast of locality 27, Fig. 15.1) and also affecting parts of the middle Xin@ River valley (northeast of locality 25, Fig. 15.1). Rheomorphic phenomena accompanied by extensive magmatic plutonism are found in the Jamanxim area (locality 27, Fig. 15.1), a northwestern continuation of the Iriri-CuruA area. The high temperatures attained in this area are associated with gold mineralization

86 5 only. The reactivated zone in the vicinity of Jamanxim has a northerly or northwesterly trend extending over 450 km in length and 100 t o 220 km in width. It constitutes the central portion or nucleus of the major Iriri-Cum6 area in which preserved old sedimentary sequences, elongated granitic intrusives and structurally discordant volcanic belts are found. Transverse structural trends and superimposed lithologic belts are a feature of the Jamanxim reactivated area in which there are pronounced tectonic lineaments. The sequence and nature of geological events are not clearly understood but it is possible that the Jamanxim area may represent the last stage in the development of the Iriri-Curua region. The reactivated zone is separated from the old crystalline region to the east, the so-called Bacaj6-ParG tectonic block, by a major tectonic discordance as well as a system of regional faults striking northwestwards across the middle Xingfi River valley. It is possible t o regard the crystalline basement of the Amazonian platform as one unit, the Jamanxim and Iriri-Curua areas occupying the more central portions. If these central regions have been more deeply eroded than the peripheral areas, due t o a greater rate of, or high uplift of the nucleus, it is probable that the greater development of rheomorphic phenomenon may be related t o deeper crustal levels. One of the most important features is the extensive magmatism that has affected the area. Fragments of the old sedimentary cover are intruded and metamorphosed creating belts of hornfelsic rocks. The large granitic or granodioritic intrusions give rise t o more prominent relief but in general dense forest and deep weathering masks the geol-

owThe western and southeastern margins of Iriri-Cum; area have not yet been studied. The Tapaj6s and Apiacas-Caiabis basins overlap part of the western side. Pronounced tectonic lineaments and a major structural discontinuity that controlled the development of the Tapaj6s basin may mark the western boundary of the Iriri-Curu6 zone. It seems probable that a continuation of the Jamanxim area reappears north of the Amazonian aulacogen in the Mapucra-Uatumg area. The Bacaja-ParG tectonic block has yielded radiometric ages >2.5 Ga. Reactivation within the Iriri-Curui zone appears t o have occurred around 1350 Ma ago which implies a Late Precambrian age for the Tapaj6s basin. The culmination of reactivation in the Jamanxim area probably lies between 750 and 900 Ma but little dating has been undertaken on rocks from this area.

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867

Chapter 1 7 KIMBERLITES, ALKALINE COMPLEXES AND CARBONATES E. SUSZCZYNSKI

KIMBERLITES

Diamonds have been recovered from three Precambrian environments but although their presence implies the existence of Precambrian kimberlite pipes, none has so far been identified. Several small diamond occurrences scattered through the crystalline basement of the Brazilides domain and the Amazonian shield have been exploited. In Pernambuco, Bahia, Minas Gerais and Goi6s the presence of diamonds associated with deeply eroded parts of the Precambrian basement raises an enigma concerning their origin and geological age. Remnants of basic or ultrabasic intrusions are always found adjacent t o these diamond occurrences. In north-central Minas Gerais state a clastic sequence of low metamorphic grade constitutes a source of diamonds that has been exploited for more than two centuries. The clastic sediments belong t o the Espinhaco Group which crops out along strike for 500 km, the majority of the important deposits being located in the northern half of its outcrop area. Diamonds are recovered by dredging placer deposits along the upper reaches of the Jequitinhonha River. The sources of the diamonds are quartzites and metaconglomerates with a phyllitic matrix. The precise age of this group is not known but it is at least Upper Precambrian. The Itacolomi Group, consisting of a thick succession of quartzites, phyllites and metaconglomerates, crops out in the southeastern corner of the Iron Quadrangle. Recent reports of diamonds from this area have been made but no details are presently available. Detrital diamonds in the Roraima Group (>2.0 Ga old) provide the source for placer deposits north of the Amazon River, and similar occurrences are known on the Gorotire and Cubencranqukm plateaux in the upper XingG River valley. Magmatic intrusions into these old sedimentary successions have been dated between 1350 and 1500 Ma. In some of the examples mentioned the source of the diamonds in clastic rocks has been identified.

868

Upper Precambrian and Eocambrian clastic, paleo-deltaic and continental successions in the Brazilides province contain beds rich in detrital diamonds, all of them always linked to regressive sedimentary sequences, at lower levels. Examples are the upper Paraguaian Group in Mato Grosso and the Paraguaqu Group within the Lenq6is-Jacar6 basin (locality 12, Fig. 15.1). These areas have been exploited for diamonds for more than a century. Similar conditions prevail in the Rio Pardo basin (locality 13, Fig. 15.1) where conglomerates contain detrital diamonds. In certain Phanerozoic sedimentary units diamonds are also conspicuous, which implies that diamondiferous pipes must have been exposed t o erosion over a long period of time. Although searches for kimberlites have been conducted in areas adjacent t o the richer Phanerozoic deposits, the results have been disappointing but minerals diagnostic of kimberlite have been found. Evidence of Phanerozoic kimberlites have been reported from the northeastern corner of the Phanerozoic Paran6 basin but no diamondiferous kimberlites have been located. It is significant that every coarse, clastic regressive sedimentary succession from Early Precambrian t o Cretaceous contains greater or lesses amounts of detrital diamonds, erosion of which has given rise t o recent placer deposits. Nothing is known about the geologic controls of the primary diamondiferous sources in the Precambrian. No preserved kimberlitic or related intrusions are known. In the younger Phanerozoic environments, small ultrabasic volcanic foci, including kimberlites are found in belts associated in space with sedimentary successions that exhibit a transition from transgressive to regressive sedimentary conditions. Whether kimberlites in the Precambrian successions have a similar spatial relationship is not known. ALKALINE COMPLEXES

Although no Precambrian intrusive or volcanic alkaline complexes are known, there are examples of stratiform alkaline bodies which the writer believes to have been formed by metamorphic or metasomatic processes from suitable Precambrian metasedimentary lithologic units. Many of these alkaline bodies have been described as Mesozoic intrusive complexes, the evidence for which is not unequivocal in the writer’s opinion. At a number of localities, such as near Ipora in southern Goias, northwest of Canabrara in northern Goias, at Tunas in Parana and in the PotiraguaItarantim areas of Bahia, the so-called alkaline complexes are clearly metasomatized zones of original carbonate composition interstratified with other metasedimentary rocks. The alkaline complex cropping out between Canabrara and Rio d o Peixe in northern Goias is 14 km long and 2 km wide. Metamorphic rocks crop out along its southern margin and on the west side there is a carbonatite-like

869 rock, originally mapped as a younger intrusion, in which alkali-amphiboles are finely disseminated. Southwards the banded carbonatite passes into a stratified, calc-silicate rock, called rodingite, which forms the top of a large stratiform ultrabasic body. The complete paragenesis of this complex has not been studied, but the author regards the carbonatite complex as the northwestern continuation of a sedimentary succession that at Canabrara was metamorphosed to a rock of ultrabasic composition. Facies changes within the sedimentary succession resulted in the formation of carbonatite-like rocks following metamorphism. Several elongated outcrops of syenitic rocks, such as those at Itiiiba and Bom Jesus da Lapa in Bahia, Itoritama in Pernambuco and Piquiri in Rio Grande do Sul, are regarded by the author as the products of metamorphism of sedimentary rocks. In many cases marbles and dolomites are found structurally and stratigraphically concordant to these stratified alkaline metamorphites. In the alkaline massifs of Bahia, for example, some carbonate layers can be traced into a sodalite facies through an intermediate stage of sodalite marble. The author cannot support the opinion that these occurrences represent Precambrian alkaline intrusives. Field relationships suggest that they are of metamorphic or metasomatic origin related t o thick stratified carbonatic metasediments, dated at between 720 and 760 Ma. Other syenitic complexes in the Brazilides structural province are related to alkaline granites formed during the final stages of the tectono-metasomatic transformation event. The large Itatiaia massif, between Rio de Janeiro and SZo Paulo, and the Princesa massif between Pernambuco and ParGba are examples of syenitic rocks crowning the tops of prominent metasomatic granitic hills. In the Amazonian shield similar occurrences are known within the Sucunduri area.

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Chapter 18

THE PALAEONTOLOGICAL RECORD E. SUSZCZYNSKI

The principal and better known organic fossils have been found in carbonate beds at the top of the Bambui System in the SZo Francisco valley. Systematic studies t o establish the exact stratigraphic position of several forms of stromatolites found in the Bambui System have been undertaken in recent years in order to establish a new subdivision of this widespread system. Among stromatolites the Collenia forms are found in the stratigraphically higher beds of the Bambui System where they may be used to define distinct palaeogeographic zones within the epicontinental sea. The presence of fossils resembling a primitive echinoderm has also been reported at even higher stratigraphic levels in the Bambui System in the Jacark-Lenqois basin. These fossils are probably of lower Palaeozoic age as the upper stratigraphic limits of the Bambui System are regarded as Lower Silurian. Stromatolite-like structures have recently been discovered in dolomites of low metamorphic grade at the top of the “Minas Series” in the Iron Quadrangle. If these structures are in fact organic, it would imply that the Bambui System was deposited immediately following the Middle Precambrian “Minas Series”. In northeastern Rondonia Federal Territory and in the southwestern portions of Amazonas state, several sandstone, siltstone and conglomeratic sequences contain flat discoidal structures that are a few millimetres in diameter. They are found particularly in specific sandstone facies which are at least early Palaeozoic in age. These structures have been interpreted as ancient algae. In the Fresco River basin, an upper tributary of the XingG River, there are several occurrences of coal in Precambrian sedimentary sequences. Individual coal seams are less than one metre thick and are intercalated with grey siltstones and white sandstones. Precambrian volcanism and granitic intrusions have caused extensive jointing and faulting in this sequence of deltaic sediments of the Cubencroqu6m-Gorotire region. It is considered that this type of coal was formed as the result of accumulation of filliform algae, possibly in

872 restricted environments under marsh conditions with an indirect contribution of carbon from volcanic sources. No characteristic fossil structures have been found in the coals but filliform aggregates are preserved. PRECAMBRIAN-CAMBRIAN BOUNDARY

No Cambrian sequences have been positively identified throughout Brazil, nor have equivalents of the Ordovician and Silurian sediments in the Amazonian aulacogen been identified in other parts of the continent. The oldest, wellestablished palaeontological record is found in Devonian sediments. This apparent absence of early Palaeozoic fossils has led some investigators to believe that the whole South American platform was subjected to continuous uplift after the termination of the Upper Precambrian. The Bambui System, however, suggests that there was no major break between Precambrian and Cambrian times. The lowest stratigraphic units in the SZo Francisco and Paraguay-Araguaia basins are folded and slightly metamorphosed, and intruded by a few scattered granitic stocks of late- or postorogenic character. Stratigraphically higher formations have not suffered deformation or metamorphism, and have yielded fossil-like forms attributed to the Silurian. Nevertheless away from the preserved sedimentary basins situated along the eastern border of the Amazonian shield or within the Brazilides domain, there is clear evidence of a major stratigraphic gap between the Precambrian crystalline basement and the oldest Palaeozoic sediments overlying this basement. REFERENCES (Chapters 13-18) Barbosa, O., 1966. Geologia bhica e econ6mica da regigo do Mddico Tapajbs, Estado do Para. Bull. 126, DGM/DNPM, Rio de Janeiro. Cordani, V., 1968. Esb6qo da geocronologia Precambriana da America do Sol. Ann, Acad. Bras. Cienc., 4 : 47-51. De Almeida, F.F.M., 1959. Reconhecimento geolbgico d o Rio AripuanZ Bull. 199, DGM/DNPM, Rio de Janeiro. De Almeida, F.F.M., 1967. Origin and evolution of the Brazilian platform. Bull. 241, DGM/DNPM, Rio de Janeiro. De Mascarenhas, J.F., 1976. Structures of greenstone-belt type in eastern Bahia, Brazil. Geol. Congr., 29th, p. 185. Dorr, J. van E., 1963. Geologic map and sections of Quadriliitero Ferrifero, 1 : 150,000. DNPM/USGS paper, Rio de Janeiro. Dorr, J.V., Soares, I. and Horen, A., 1956. The manganese deposits of Minas Gerais, Brazil. Int. Geol. Congr., ZOth, Mexico, Manganese Symp., Vol. 111, pp. 279-346. Goni, J.C., 1962. Origine des roches ultrabasiques et serpentineuses du Precambrien de Rio Grande d o Sul, B r e d : mode ge gisement et mineralisation. Bull. 12, Esc. Geologia, U.F.R.G.S., Port Allegre. McDougall, I. and Compston, W.,1963. Leakage of radiogenic argon and strontium from minerals in Proterozoic dolerites from British Guiana. Nature, 198, p. 564. Suszczynski, E., 1967. La tectonique et la geologie de la partie orientale du Bouclier Brdsilian. Bull. SOC.GBol. Fr. 7th Ser.; Vol. IX, pp. 147-152.

873 Suszczynski, E., 1 9 7 0 . La g6ologie et la tectonique de la Platforme Amazoniene. Geol. Rundsch., 59( 3 ) : 1232-1 252. Suszczynski, E., 1 9 7 3 . Metalogenetic map o f Brazil: explanatory note of legend. D W M / MME, Brasilia. 1 5 pp. Suszczynski, E., 1 9 7 5 . 0 s recursos minerais reais e potenciais do B r ad e sua metalogenia. Livraria Intersciencia Ltda., Rio de Janeiro, 533 pp.

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875

SUBJECT INDEX

Abbabis Formation, 757 - inlier, 757-758 Abbotshall Beds, 41 Acaciella, 3 79 Acari Group, 854 Adelaide Geosyncline, 327-354 - -, Delamerian Orogeny, 349-354 _ _ , geochronology, 334,345 - _ , problems, 330-331 _ _ , sedimentation, 335-345 - - , stratigraphy, 331-334 - - , tectonic framework, 327-330 - -, tectono-stratigraphic, 345-349 Adelaidean, 3,216,283-296,327-349 - age, 287,294 - glacials, 289-294 - paleogeography, 287-288,294-295 Albert Edward Group, 292 Alcheringa, 366 -36 7 Alkaline complexes, 188-189,803-810,868869 Amadeus Basin, 283-289 - Transverse Zone, 20-25 Amosite, 536 Angola-Kasai craton, 488-492,721 Annelida, 387 Antimony, 63,91,441-442 Araras Group, 856 Archaean, -, Angola, 488-490 -, Australia, 36-81,220-221,314-315 -, Namibia (S.W. Africa), 756-765 -, South Africa, 424-453,501-510 -, South America, 845 -, Zambia-Zaire, 721-723 -, Zimbabwe, 454-488 Archaean chelogen, 12-16 Arthropoda, 386 Arunta Inlier, 215-216, 274-283 _ - , geochronology, 279-281 _ - , metamorphism, 281 _ _ , stratigraphy, 275-278 - -, tectonic models, 282-283

- - , tectonic setting, 274-276 - -, tectonic zones, 278-281 Ashburton Formation, 120-121 Auvergne Group, 253,254

Bacaji-Par6 Block, 841 BaGo Complex, 845 Backdoor Formation, 164 Badgerrada Group, 177 Baicala, 818 Bambui System, 854,871 Bandolierskop Formation, 661 Bangemall Basin (Group), 158-1 77 - _ , eastern facies, 164 - -, northern facies, 163 - _ , sedimentation, 169-170 - - , stratigraphy, 158-165,170-173 - - , structure, 165-168, 173-175 - - , western facies, 160-163 Barberton greenstone belt (Barberton Mountain Land), 432-438 - - - , geochronology, 436-437 - _ _ , metamorphism, 436 - _ _ , mineralization, 437-438 - _ _ , stratigraphy, 432-435 - - - , structure, 435 Basement re-activation, 19 -, see also Mobile Belts Basins, sedimentary, 109-134,134-152, 156, 157,158-177,179,235-239,239249,250-251,251-255,269,283289,511-513,529-530,532,536541,854,855,856 - relationship between the Hamersley and Nabberu, 152-153 -, structurally preserved, 855 Bastion Group, 236,238 Bauxite, 66 Beaton’s Creek Conglomerate, 1 1 3 Beazley River Quartzite, 120-121 Beltanella, 385 Beneficente Group, 854 Bentley Supergroup, 175-177 Beryl, 66,91,442,622

876 Billeranga Group, 178 Biocatenoides, 372 Burindudu Group, 250,251 Bismuth, 66 Black Flag Metasediments, 40 Black Range Dyke Suite, 188 Bodoquena Group, 856,867 Brazilides orogeny, 843-844,864-865 Brockman Iron Formation, 114-118 Broken Hill subdomain, 313,317-320 Brown Lake Metasediments, 40 Bulawayan Group, 463-470,816,820 Bullita Group, 254 Burra Group, 332,333,336-338 Bushveld Complex, 599-618 - -, acid phase, 611-612 - -, geochronology, 616 - -, mafic sequence, 603-610 - -,tectonic setting, 614-618 CalIanna Beds, 332,333,335-336 Calyie Sandstone, 164 Cambrian, base of, 389 Campbell Group, 531-534,816-817,820821 Capricorn Formation, 120-121 Caraca Group, 856-857 Carbonatites, 188-189,803-810,868-869 Cardup Group, 178 Carpentarian, 2 Carr Boyd Group, 254-255 Caudiculophycus, 380 Causeway Beds, 41 Cave Rock Beds, 40 - - , ophiolite belt, 40 Central Rand Group, 5 17-5 18 Cephalophytanan, 380 Chelogenic cycle, 9-12 - - , Archaean, 12-16 - _ , Proterozoic, 16-18,28 Chondroplon, 385 Chrome, 188,482,577,604,606-607,608 Chrysotile, 91,437,451,481-482 Chuniespoort Group, 531434,816-817, 820-821 Cliff Springs Formation, 112 Cloudina, 824 Cobalt, 62,738 Coelenterata, 386 Coobarra Formation, 164 Coodardoo Formation, 163 Coolgardie Ophiolites, 41

Copper, 62,65-66,90,134,152,177,272, 274,492,561-562,621,626,663,664 702-703,719,723,727,737-739, 781-784,805,850 Coramba Group, 856 Cratons, Angola-Kasai, 405 -, Gabon, 721 -, Kmpvaal, 397-399 -, Rhodesian, 399-404 - , Tanzanian, 405 -, Zambian, 404-405 Crocidolite, 133,536 Crowhurst Group, 236,238 Cubencranquin Formation, 853-854 Curnamoona Cratonic nucleus, 330 Curran Formation, 163 Damara Mobile Belt, 754-788 - - -, Abbabis inlier, 757-758 - - _ , basement, 756-765 - - _ , granitoids, 778-781 _ - - , geochronology, 778-779 _ - - , Kamanjab inlier, 758-763 - - -, Kaokoveld inlier, 763-765 - - _ , lithologies, 756-765,766-770, 770-777 - - - , metamorphism, 778-781 - - -, mineralization, 781-786 - - _ , Rehoboth inlier, 757 - - _ , structure, 778-781 - - - , tectonic framework, 786-788 Damara Supergroup, 689,770-777,817-819 Delamerian Orogeny, 330,349-354 - _ , deformational style, 350-354 _ - , subdivisions, 330 - - , timing, 349-350 Devil’s Creek Formation, 162 Deweras Group, 618-627 Diamonds, 867-868 Dickinsonia, 384 Discovery Chert, 161, 169 Domains, orogenic, 5 , 6 Dominion Group, 516 Doornpoort Formation, see Ghanzi Formation Duck Creek Dolomite, 120-121 Dudawa Beds, 178 Duerdin Group, 292 Earaheedy Group, 134-143 Ediacara, Australia, 383-389 -, Southern Africa, 822-825 Elim Formation, 757

877 Emeralds, 442,664 Eoentophysalis, 371 Epupa Formation, 764 Ernietta, 822 -8 23 Espinhaqo Group, 867 Felsic volcanic association, 52-54 Fem‘monolis, 372 Fickling Group, 258 Fig Tree Group, 434-435 Fitzmaurice Group, 255,256 “Flinders Gneiss”, 314 - Series”, 314 Fluorite, 536,720, 784 Fold belt, Delamarian, 349-354 - -, Opthalmian, 153-154 - _ , Stanley, 153-154 Ford’s Creek Shale, 16 I, 163 Fortescue Group, 2,111 - 113 Fraser Complex, 181-182 Frere Formation, 138-141 Gabon Craton, 721 Gariep Group, 689,818 Gawler Block, 25,26 - Province, 309-327 - Range Volcanics, 310-322,324-325 Geochronology, 60-62,88-90,126-127, 148-149,158,181-183,184,206207,224,228,278-281,287,314315,322,325-326,334,436-437, 441,450,453,455-457,463,488489,490,501-503,513,520-522, 547-548,561,616,620,699-702, 841,842,843-844 Ghanzi Formation, 768-769 Gindalbie Group, 4 1 Glacial sequences, 289-294 Glaessnerina, 385 Glengarry Group, 143-144 Glenobotrydion, 380 Glidden Group, 255 Gold,62,63-65,90-91,113,134,152,177, 274,437,441-442,445,451,480481,492,510,519-520,527,621,664, 723,860 Gorotire Formation, 853 Granitoid rocks, 58-62,85-88,181,182183,220-221,310,312,316,424430,449-453,457-460,463,471473,488-489,562-568,658-659, 683-686,712-713,718,763,766, 769,778-781,848-849

Gravity data, 5,107,411-419 Great “Dyke”, 572-578 Greenstone belts, 37-55,7341,403-438, 438-442,442-445,445-449,463470,845-846 Griquatown Group, 531-534,816-817,820821 Gundockerta Group, 40 Gunga meta-argillites, 41 Guyana orogeny, 842 Hallidaya, 384 Hall’s Creek Group, 221-225 - - -, age,224 - - _ , sedimentation, 224-225 Hall’s Creek Inlier, 221-226 Hall’s Creek-Fitzmaurice Mobile Zone, 25525 6 Hamersley Basin, 109-134 - -, geochronology, 126-127 - -, metamorphism, 124-126 - -, mineralization, 130-134,152-153 - -, relationship to Nabberu Basin, 152-153 - -, sedimentation and tectonics, 127-130 - -, stratigraphy, 109-124 - -, structure, 124-126 Hamersley Group, 113-119,127-129 Hart Dolerite, 238 Haslingden Group, 259 Huab Formation, 763 Hutchinson Group, 310-311,315-316

llgarari Bed$, 164 Iron formation, 46-47,317-320 - -, depositional models, 154-156 Iron ore, 65,91,131-133,152-273,437438,492,536,783,859-860 lrregully Formation, 160, 169 lrumide Orogeny see Kibaran Itabira Group, 857 Itacolomi Group, 857 Jacobina Series, 857 Jerrinah Formation, 112 Jillawarra Formation, 169 Kaffirskraal pyroxenite, 612 Kaieteur Formation, 85 3 Kalpini Group, 40 Kamanjab inlier, 758-763 Kaokoveld inlier, 763-765 Katangan orogeny, 727-736,741,742 - -, see also Pan-African

878 Katangan Supergroup, 727-739 _ _ , distribution, 727-728, _ _ , mineralization, 737-739 _ _ , stratigraphy, 728-729 _ _ , structure, 728-736 Keepera Group, 284 Khoabendus Formation, 763 Khomas Trough, 776 KiangiCreek Formation, 161,164, 169 Kibaran orogeny, 723,750-753,763,765769 Kibaran Supergroup, 7 2 3 -72 7 _ _ , geochronology, 724,726 - -, mineralization, 727 _ _ , stratigraphy, 723-727 Kimberley Block, 25 - Group,238 Kimberlite, 809-810,867-868 Klein Aub Formation, 767-768 Klipriviersberg Group, 516,522-525 Koras Formation, 688 Kulele Creek Formation, 138, 143 Kunene Complex, 489-490,578-599,764 _ _ , genesis, 590-597 _ _ , geochronology, 597-599 _ _ , lithology, 579-589 _ _ , setting, 597-599 _ _ , structure, 589-590 Kuniandi Formation, 292 Kurabuku Formation, 161,163 Kurrawang Beds, 40 - conglomerate, 40 - Group,40 Kwebe Formation, 768-769 Labouchere Formation, 144 Lake Dam Beds, 40 Lamboo Complex, 225-226 Lawn Hill Platform, 250 _ _ _ , see also Mount Isa orogen Lead, 65-66,133,271,274,317-318,702703,720,738-739,782-783,860 Leeuwin Block, 184 Limbunya Group, 250-25 1 Limpopo Belt, 641-671 _ _ , anorthosite, 657-658 _ - , basement, 654-656 _ _ , craton margins, 642-647,661 _ _ , central zone, 652-660 _ _ , granitoids, 642-647,658-659 _ _ , granulite, 647-649,661 _ _ , geochronology, 665-666 _ - , rnafic complexes 649-650

- _

, Mahalapye, 660 _ _ , metamorphism, 650,659-660,661662 _ _ , mineralization, 663-664 _ _ , north marginal zone, 642-650 _ _ , south marginal zone, 660-662 _ _ , structure, 650,659-660,661-662 _ _ , synthesis, 664-671 - _ , Tuli-Sabi shear belt, 650-652 Limpopo Group, 656-657 Lincoln Complex, 310,316 Lithium, 66,91,442 Lomagundi Group, 618-627,741 Losberg Complex, 612. Louisa Downs Group, 292 Lower greenstones, Zimbabwe, 465-466 Lower Kundelungu Group, 728-729,821 Lower sedimentary association, 49-5 1 L u f i a n orogeny, see Katangan orogeny Macuville Group, 654-656 Mafic/ultramafic association Yilgarn, 43-49, 51-52 Mafingi Group, 752-753 Magondi orogeny, 741 Magnesite, 66 Makuti Group, 742 Malay Road Group, 242,244,248 Malbon Group, 259 Mandilla Beds, 40 Manganese, 65, 133,152,492,536,783,850, 859,860 Manganese Group, 122-124,129-130 Marienhof Formation, 757 - -, see also Abbabis and Elim Formations Marinoan, 329,332, 333,340-344 Mary Kathleen Group, 259 Mapumolo Group, 712 Mashaba Igneous Complex, 570-572 - Ultramafic Suite, 475 Mashonaland Dolerites, 5 6 1 Matchless belt, 776 Matsap trough, 541 McArthur Group, 242-243 . -, palaeogeography, 247-248 Mchinji Group, 746 McNamara Group, 257 Merougil Creek Beds, 40 Metamorphic complex, _ _ , Albany-Fraser, 179-183 - _ , Gascoyne, 183-184 - - , Leeuwin, 184 _ _ , Lincoln, 310,316

879 , Mullingarra, 185

-

.-

_ _ _ _

_ , Namaqua, 671-705 _ , Northampton, 184 _ , Rudall, 185-186 _ , West Nile, 772

- - , Willyama, 312 Miami Event, 742-743 _ _ , see also Pan-African Mica, 719 Microfossils, 363 -366,368- 3 7 7 , 3 79- 382, 819-822 Millidie Creek Formation, 144 “Minas Series”, 856-857,871 Mineralization, 62-66,90-91,130-134, 152, 177,219,249,270-273,437-438, 441-442,445,480-482,491-492, 510,519-520,527,536,561-562, 575-576,607,608,609,663--664, 702-703,718-720,737-739,781786,805,849-850,859-861 Mine (Roan) Group, 728-729,822 Miningarra Subgroup, 142-143 Mobile Belts and Zones, _ _ , Damara, 754-788 - _ , Hall’s Creek, 214 _ _ , Irumide, 407 _ _ , see also Kibaran _ - , Kibaran, 407-408,746,748,749-752, 765-769 _ _ , King Leopold, 214 _ _ , Limpopo, 405-406,641-671 - -, Mocambique,410 - -, see also Pan-African - -, Namaque-Natal, 408,671-715 - -, Pan-African, 408-410,729-736,73975 2 - -, South America, 863-865 _ _ , Ubendian, 406-407,747,748 _ _ , Zambezi, 715-720,739-743 Mobile Belts, Africa, origins of, 410-41 1 Mocambique, see Mobile Belts and Zones Moodies Group, 435 Moora Group, 178 Mopunga Group, 284 Morelands Group, 41 Mosquito Creek Group, 12,73-78,91 Mount Albert Group, 258 Mount Bruce Supergroup, 35,109-124 _ - -, geochronology, 126-127 _ _ - , metamorphism, 124-126 _ _ - , mineralization, 130-134 _ _ _ , stratigraphy, 109-124 ~- - - , structure, 124-126

Mount House Group, 292 Mount McGrath Formation, 120-121 Mount Robert Volcanic Group, 442-444 Mount Robinson Ophiolites, 40 Mount Roe Basalt, 112 Mount Isa Group, 257-263 Mount Isa Orogen, 215,256,273 _ - _ , age, 262-263 - - _ , basement, 265-266 _ _ _ , Carpentian cover, 266-268 _ _ _ , igneous rocks, 269 _ _ _ , metamorphism, 269-270 _ _ _ , mineralization, 270-273 _ _ _ , palaeogeography, 265 - - _ , stratigraphy, 257-262 _ _ _ , structure, 270 - _ _ , tectonic elements, 263-265 - _ _ , tectonic setting, 260-261 Mount Jope Volcanis, 112 Mount Thirsty Beds, 4 1 Mount Vernon Sandstones, 161, 163 Mulden Group, 777 Mulgabbie Group, 41 Mulgarra Sandstone, 138, 143 Mullingara Inlier, 185 Mungari Beds, 40 Murchison greenstone belt, 438-442 - - _ , geochronology, 441 _ _ _ , metamorphism, 441 - - _ , mineralization, 441-442 _ _ - , stratigraphy, 438-440 - - _ , structure, 440-441 Murphy Inlier, 234 Mwami Group, 746,750,752 Myxomorphia, 369 Nabberu Basin, 134-152,363 _ _ , geochronology, 148-149 _ _ , metamorphism, 145-148 _ _ , mineralization, 152 _ _ , sedimentation and tectonism, 149-151 _ _ , stratigraphy, 134-144 - _ , structure, 145-148 - _ , sub-basins, 134-145 Nama Group, 689,777,817-819,821-822, 822-825 Namaqua-Natal belt, 671 -715 Namaqua Metamorphic Complex, 671-704 _ _ -, craton margin, 678 _ _ -, dykes, 686-687 _ _ -, geochronology, 699-702 _ _ -, granitoids, 683-686 - _ -, lithology, 672-677

- - - , mafic intrusions, 686,688 _ - _ , metamorphism, 694-699 - - -, mineralization, 702-703 - - -, pegmatites, 687-688 - - - , structure, 690-694 - - - , synthesis, 703-705 Nasepia, 8 2 3 Natal sector Namaque-Natal belt, 705-715 _ - , geochronology, 714 _ - , granitic zone, 712-713 _ - , metamorphism, 711-712 - - , migmatite zone, 712 - -, northern frontal zone, 707-712 _ - , southern granulite zone, 713-714 - -, structure, 708-711 Nickel,62-63,66,90,482,663,723,850 Nilling Beds, 177 Niobium, 66 Noganyer Group, 41 Northampton Block, 184 Nosib Group, 771-773 Nullaginian, 2 Nullagine Succession, 74 Obconicophycus, 380 Oendeolongo System, 490 Olary subdomain, 312,316,320-322 Onverwacht Group, 432-434,819 Otavi-Swakop Groups, 773-775 Outjo facies, 771

Padbury Group, 144 Palabora complex, 803-806 Palaeobiology -, Archaean, 362-363,813-816,819-820 -, Archaean-Proterozoic boundary, 826-827 -, Australia, 361-391 -, Proterozoic, 363-369,816-825,827-828 -, Southern Africa, 813-828 -, South America, 871-872 Palmeiral Formation, 854 Pan-African orogeny (event), 408-410,729736,739-788 Penneshaw Beds, 4 1 Phycodes, 8 24 Pietersburg greenstone belt, 442-445 _ - -, metamorphism, 445 _ - -, mineralization, 445 _ _ -, stratigraphy, 442-444 - _ -, structure, 444-445 Pietersburg Sequence, 442-444 Pilanesberg alkaline complex, 807-810

Pilbara Block, 73-96 - -, evolution, 91-94 - -, geochronology, 88-90 - - , granitoids, 85-88 - -, metamorphism, 81-85 - -, mineralization, 90-91 - -, stratigraphy, 73-81 - -,structure, 81-85 Pilbaria. 8 17 Pine Creek Inlier, 227-234 - - - , age, 228 - _ _ , igneous intrusion, 232 - _ _ , metamorphism, 232 _ _ - , mineralization, 233-234 - _ - , sedimentation, 231-232 - _ - , tectonic setting, 227-228 Pinjan Chert Breccia, 122,123 Piracicaba Group, 85 7 Piriwiri Group, 618,627,741 Platberg Group, 5 25 -526 Platform cover, Australia, 26-27 213, 216217 Platinum, 62,706-608 Pniel Group, 526-527 Pongola Group,501-510,813 - -, distribution, 501 - _ , geochronology, 501,503 - -, mineralization,510 - -, stratigraphy, 503-510 - -, structure,503 Postmasburg Group, 5 35 -5 36 Precambrian-Cambrian boundary, 389,827828,872 Pretoria Group, 535,8 17 Princess Ranges Quartzite, 138, 142 Proterozoic - Angola, 490-492 - chelogen Australia, 16-18 - cover Australia, 213,216-217,283-295 - crustal structure, 189-193 - North. Australia, 205-307 - South Africa, 501-618,671-714,770788,803-810 - South America, 851-861,863-865 - South Australia, 309-360 - West Australia, 105 -203 - Zambia-ZaBe, 721-739,743,754 - Zimbabwe, 618-627,715-720,739-743 Provinces, structural, 1-29,208-209,397411 -, Albany-Fraser, 179- 183 -, Amazonian, 835,841-843

881 -, Andean, 835 - Brazilian-Guyanan, 835 -, in central Australia, 214-216 -, Eastern Goldfields, 38 -, Gascoyne, 8 , 1 8 3 -, Gawler, 26,309-327 -, Kimberley, 8,280 -, Murchison, 38 -, Musgrave-Fraser, 7-8,19,326 -, in N. Australia, 211-213 - in N.E. Queensland, 215,273-283 -, PiIbara, 20 - recognition of, 3-9,208-209 -, South African, 397-411 -, Southwestern, 38 -, Tasman, 7 - in W. Australia, 107-109,210-211 -, Yilgarn, 20 Pteridium, 822 Rangea, 822,823 Red Lake ophiolites, 40 Red schist belts, 847-850 Rehoboth Group, 767-768 Rhodesian (Zimbabwe) craton, 399-404 - - , evolutionary models, 454-488 _ - , metamorphism, 479-480 - - , mineralization, 480-482 - - , structure, 476-479 Rio Fresco Formation, 853 Rio Negro-Roraima orogeny, 842-843 Robinson Range Formation, 144 Rondonides orogeny, 842 Roper Group, 242,244,248 Roraima Formation, 853 Rudall Inlier, 185-186 Rushinga Group, 742 Ruzizian (Ubendian) fold belt, 722-723

Scorpion Group, 157 Sebakwian Group, 460-463 Shamvaian Group, 470-471 Silver, 703,781,783 Sinclair Group, 688-689,756,766-767 Sleaford Complex, 315 Soldiers Gap Group, 259 Soutpansberg Group, 536-556 - _ , palaeogeography, 547-555 _ - , stratigraphy, 543 - -, structure, 543-545 - -, tectonic evolution, 547-556 - -, volcanicity, 545-546

Speewah Group, 235-238 “Stirling-Barren Series”, 178-179 Stratigraphy, 37-55,73-81,109-124,134144,156-157,158-165,170-173, 177-179,235,242,253,261-262, 286,289-293,316-322,331-334, 432,435,438-440,442-444,447, 461,467,488-491,503-510,515518,522-527,531-536,541-543, 559-560,618-627,656-657,661, 672-677,728-729,770-777 Stromatolites, 367,374,375, 377-379,383, 813-819,871 Structural framework, 1-29,397-411,708710,786-788,835-844 Structure, 55-58,81-85,124-126,145-148, 238-239,248-249,251,256,288289,322-327,435,440-441,444445,447-449,475 -479,503,5 13515,522,529,543-545,560-561, 650,659-660,661-662,690-694, 711-712,729-736,746-754,857859 Stuart Shelf, 327-354 - - , see also Adelaide Geosyncline Sturtian, 329, 332, 333,338-340 Sub-basins -, Earaheedy, 134-143 -, Glengarry, 143-144 -, metamorphism of, 145-148 -, Padbury, 144 -, relationships of, 145 -, sedimentation in, 149-151 -, structure of, 145-148 Sub-provinces, 7 , 8 , 1 4 , 1 7 , 3 4 , 3 8 , 2 2 1 - 2 2 6 , 227-234 Surumua Formation, 853 Sutherland greenstone belt, 445 -449 - _ _ , geochronology, 449 - _ _ , metamorphism, 449 - - _ , mineralization, 449 - _ - , stratigraphy, 447 - _ - , structure, 447-449 Swakop facies, 770 Swaziland Supergroup, 432-438 Tantalum, 622,727 Tawallah Group, 243-244 - -, palaeogeography, 245-246 Tectonic setting and tectonism, 127-130, 149-151,239-244,260-261,274276,278,286-327,614-618

Tennant Creek Inlier, 226 Tewinga Group, 260 Tin,66,91,274,622,719-727,784-785,861 Tolmer Group. 254 Tooloo Group, 137-142 Top Camp Dolomite, 163, 169 Torrens Hinge Zone (lineament) 310,327 Torrensian, 329,332, 333,336-338 Trace Fossils, 388-822-825 Trainor Inlier, 157 Transvaal Supergroup, 527-536 - - , distribution, 527-528 - - , geochronology, 529 _ - , mineralization, 536 - -, stratigraphy, 531-536 - - , structure, 529-530 Triangle Islands Beds, 40 Trompsberg Complex, 612 Tungsten, 66,274,719,727 Turee Creek Group, 119-120,127-129 Umberatana Group, 332,333,339-346 Umkondo Dolerites, 561 Umkondo Group, 556-562,742 - - , deposition, 560 _ _ , distribution, 556-559 - -, geochronology, 561 - - , metamorphism, 560-561 _ - , mineralization, 561-562 - - , stratigraphy, 559-560 _ - , structure, 560-561 Upper Greenstones, Zimbabwe, 466-470 Upper Kundelungu Group, 728-729 Uranium, 134,272-273,274,519-520,737, 785-786 Usushwana Complex, 568-570 Ventersdorp Supergroup, 520-5 27 - - , distribution, 520 - - , geochronology, 520-522 _ - , mineralization, 527 _ _ , stratigraphy, 522-527 _ _ , structure, 522 Veteronostacale. 380 Wanda Wanda Beds, 40 Wandiwarra Formation, 138, 142 Wangawol Formation, 138,142-143 Warrawoona Beds, 73 Waterberg Group, 536-556 - -, palaeogeography, 547-555 - -, stratigraphy, 541-543 - -, structure, 543-545

- _ , tectonic evolution, 555-556 _ _ , volcanicity, 545-546 Wattie Group, 253,254 Wenmillia Formation, 177 West Congo - - syncline, 721,726 - - System, 491 West Nile Complex, 722 West Rand Group, 516-517 Widgiemooltha Dyke Suite, 186-187 Willourian, 329,332,333,335-336 WiUyama domain, 312 Wilpena Group, 332,333,342-344 Windidda Formation, 138,141-142 Witwatersrand Supergroup, 5 11-520 - -, distribution,511-513 - -, geochronology, 513 - -, metamorphism, 5 19 - -, mineralization, 519-520 - -, stratigraphy, 515-518 - -, structure,513-515 Wolkberg Group, 531 Woodline Beds, 179 Woodroffe thrust, 310 Woolyeenyer Group, 4 1 Wyloo Group, 120-121,129-130 Yandanooka Group, 177-178 Yelma Formation, 137-138 Yeneena Group, 170,174, 175 Yilgarn Block, 36-73 - -, evolution, 66-73 - -, granitoids, 58-62 - -, metamorphism, 55-58 - -, mineralization, 62-66 - -, stratigraphy, 37-55 - -, structure, 55-58 - -, subdivision of, 34, 36-39 Yilmia ophiolite belt, 40 Zambezi metamorphic belt, 715-720,739743 - - -, geochronology, 720,739,741-742 _ - -, granitoids, 718 - - -, Karnativi Formation, 718 - _ -, Malaputese Formation, 717-718 - - -, mineralization, 718-720 - _ -, stratigraphy, 741-743 - _ -, structure, 742-743 _ - -, Tshontanda Formation, 718 Zandian orogeny, 490 Zinc, 65-66,133,271,274,317-318,702703,723,738-739,781-783,860