Developments in Precambrian Geology 10
PROTEROZOIC CRUSTAL EVOLUTION
DEVELOPMENTS IN PRECAMBRIAN GEOLOGY Advisory Ed...
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Developments in Precambrian Geology 10
PROTEROZOIC CRUSTAL EVOLUTION
DEVELOPMENTS IN PRECAMBRIAN GEOLOGY Advisory Editor B.F. Windley Further titles in this series 1. B.F. WINDLEY and S.M. NAQVl (Editors) Archaean Geochemistry 2. D.R. HUNTER (Editor) Precambrian of the Southern Hemisphere 3. K.C. CONDIE Archean Greenstone Belts 4. A. KRONER (Editor) Precambrian Plate Tectonics 5. Y.P. MEL'NIK Precambrian Banded Iron-formations. Physicochemical Conditions of Formation 6. A.F. TRENDALL and R.C. MORRIS (Editors) Iron-Formation: Facts and Problems 7 . B. NAGY, R. WEBER, J.C. GUERRERO and M. SCHIDLOWSKI (Editors) Developments and Interactions of the Precambrian Atmosphere, Lithosphere and Biosphere 8. S.M. NAQVI (Editor) Precambrian Continental Crust and its Economic Resources 9. D.V. RUNDQVIST and F.P. MITROFANOV (Editors) Precambrian Geology of the USSR
DEVELOPMENTS IN PRECAMBRIAN GEOLOGY 10
PROTEROZOIC CRUSTAL EVOLUTION
Edited by
K.C.CONDIE Department of Geosciences, New Mexico Institute of Mining & Technology, Socorro, New Mexico 87801, U.S.A.
ELSEVIER, Amsterdam - New York - Tokyo
1992
ELSEVIER SCIENCE PUBLISHERS B.V. Sara Burgerhartstraat 25 P.O. Box 21 1.1000 AE Amsterdam, The Netherlands
Library o f Congress Cataloging-in-Publication
Data
Proterozoic crustal evolution / edited by K.C. Conoie. p. cm. -- (Developments in Precambrian geology ; 10) Includes bibliographicai references and indexes. ISBN 0-444-88782-2 (alk. paper) 1 . Earth--Crust. 2. Geology, S t r a t i g r a p h i c - - P r o t e r o Z o i C . I. Condie, Kent C. 11. Series. QE5 1 1 . P79 1992 551.7'15--dc20 92-34776 CIP
ISBN: 0-444-887822
0 1992 Elsevier Science Publishers B.V. All rights reserved. No part of this publication may be reproduced, stored i n a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without the prior written permission of the publisher, Elsevier Science Publishers B.V., Copyright and Permissions Department, P.O. Box 521,1000 A M Amsterdam, The Netherlands. Special regulations for readers in the U.S.A. - This publication has been registered with the Copyright Clearance Center Inc. (CCC), Salem, Massachusetts. Information can be obtained from the CCC about conditions under which photocopies of parts of this publication may be made in the U.S.A. All other copyright questions, including photocopying outside of the U.S.A., should be referred to the publisher. No responsibility is assumed by the publisher for any injury and/or damage t o persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein. This book is printed on acid-free paper. Printed in The Netherlands
V
CONTRIBUTING AUTHORS
J. LAWFORD ANDERSON Department of Geological Sciences, Universityof Southern California, Los Angeles, CA 90089, U.S.A. NICOLAS J. BEUKES Deparhnent of Geology, Rand Afiikaam University, Johannesburg, 2000 South Afiica KENT C. CONDIE Department of Geoscience, New Mexico Institute of Mining & Echnoloa, Socorro, NM 87801, U.S.A. JOHN C. GREEN Department of Geology, University of Minnesota, Duhith, MN 55812, U.S.A. SIMON L. HARLEY Department of Geology and Geophysics, University of Edinburgh, Edinburgh, Scotland EN9 3Ju! UK R.E. HARMER Institute for Geological Research on the Bushveld Cornplq Universiteit van Pretoria, 0001 Pretoria, South Afiica H.H. HELMSTAEDT Deparhnent of Geological Sciences, Queens University,Kingston, Ontario K7L 3N6, Canada CORNELIS KLEIN, Jr. Department of Geology, University of New Mexico, Albuquerque, NM 87131, U S.A. JEAN MORRISON Department of Geological Sciences, University of Southern California, Los Angeles, CA 90089, USA. P. JONATHAN PATCHETT Department of Geosciences, University of Arizona, Tucson, AZ 85721, U.S.A. JOAQUIN RUIZ Department of Geosciences, University of Arizona, Tucson, AZ 85721, l% S.A. DAVID J. SCOTT Department of Geological Sciences, Queens Universitj;Kingston, Ontario K7L 3N6, Canada Present address: GEOTOe Universitt du Que'bec (iMontrtal, C.R 8888, SuccursaleA, MontrkaJ Qutbec H3C 3P8, Canada
VI
Contributing authors
TERRY E. SMITH Deparhnent of Geology and Geological Engineering, University of Wino304 Winhor; Ontario N9B 3P3, Canada JOHN TARNEY Department of Geology, Universig of Leicester, Leicestel; LEI 7RH, England, U K . G. von GRUENEWALDT
Institute for Geological Research on the Bushveld Complq Universiteit van Pretoria, 0001 Pretoria, South Africa ROBERT A. WIEBE Departnient of Geology, Franklin and Marshall College, Lancaster, PA 17604, U S.A. BRIAN E WINDLEY Department of Geology, University of Leicestel; Leicestel; L E I 7 N , England, U K
v1I
CONTENTS
Contributing authors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
V
INTRODUCTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . K.C. Condie Reference . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
1
6
Chapter 1. VOLCANIC ROCKS OF EARLY PROTEROZOIC GREENSTONE BELTS . . . . T.E. Smith
1
....... ................................ Introduction . . . . . . . . . ....... Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Early Proterozoic greenstone belts ............................ . . .. . . . . . . The Baltic Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .......... ....... ....... Lewisian Complex, northwest Scotland . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ....... The Laurentian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ....... Central and southwestern North America ................................. ....... The South American Shield ......................... .... . . . . . . The African Shield ............................... ....... The Indian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ....... China . ................................... ....... The Australian Shield ...................................... .......... Discussion and conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ....... Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References ................................ ........ .......
1
Chapter 2 . THE PROTEROZOIC OPHIOLITE PROBLEM H.H. Helmstaedt and D.J. Scott
8 10 10 21 22 31 34
36 38 39 40 44 46 46
.......
55
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. . . . . . . Problem of preservation of Proterozoic ophiolites . . .. . . . . . . Pan-African ophiolites ........................ .... . . . . . . Trans-Saharan suture . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Bou Azzer. Morocco. 59 - Pharusian belt. 63 - Dahomey belt. 64 Damara belt. southern Africa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Arabian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Yanbu suture. 67 . Bir Umq suture. 69 . Nabitah suture zone. 69 . Urd and Al Amar sutures. 70 Early Proterozoic ophiolites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . BalticShield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Jormua complex. 73
55 51 59 59 64 65
71 71
Contents
VIII
... Canadian Shield . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Cape Smith belt, 75 - The Purtuniq ophiolite, 77 .............. Proterozoic accreted terranes of southwestern United States The Payson ophiolite, 84 Diversity of Proterozoic ophiolites ............................................. ............... Ophiolite analogues in Archean greenstone belts? . . . . . . . ........................................................ Acknowledgements ......................................... References . . . . . . . . . . Chapter 3. PROTEROZOIC RIFTS J.C. Green
,
. .. . .. .. . . . . . . . . . . . . . . . . .. .. . . . . . . . . . . . . . . . . . .
................................... Introduction . . . . . . . . . . . . . . . . . . . . . . . . ... ...... Early Proterozoic rifts: 2.5-1.9 Ga . . . . . . . . . . . . . . . . . . , ., . . . . . . . . . . . . . . . South African basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ...... . . . Circum-Superior belt ......................................... Animikie rift ..................................... . . . _ . . _ _ Svecokarelian belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Wopmay orogen . . . . . . .............................................. Early to Middle Proterozoic a ............................................ Canadian Shield sequences .............................................. SouthernAfrican basins ............................... .. ............... Australian basins ................................................... Middle Proterozoic rifts: 1.5-1.0 Ga . . .................................... . . . . .. . . . ........................ North American Midcontinent Rift Belt-Purcell-Wernecke basins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Riphean aulacogens of the former U.S.S.R. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ....... ..... ..... ... Grenville and Telemark supracrustals . . . . . . . . . . . . . . .......................................... Kalahari copper belt ......................... Kibaran belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .................................... Middle to Late Proterozoic rifts: 1.0-0.6 Ga Damara-Ribiera Province . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Dahomeyan-Pharusian (Trans-Sahara) belt . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ...._..._............ Iapetus rift . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Australo-American Trough, Proto-Arctic and Proto-Pacific rifting . . . . . . . . . . . . . . . . . . . . . . .................................... Proterozoic rifts and plate tectonics . . . . . . . . . . . . References ........ . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ...................
74 83 84 86 88 88
97
97 99 100 101 104 106 109 110 110 111 113 116 116 120 122 122 123 124 125 125 126 127 132 135 136
GEOCHEMISTRY AND SIGNIFICANCE O F MAFIC DYKE SWARMS IN THE PROTEROZOIC . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J. Tarney
151
Introduction ........................................ ............................. Form and features . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chronology . . . . ..................................... Petrological cliarac .................................... ..................... Geocliemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .....
151 152 154 154 156
Chapter 4
Contents
IX
..................... Later Proterozoic dykes . . . . . . . . . . . . . . . . . . . . . . . . ................................... Mantle evolution . . . . . . . . ................................. PREMA mantle and Proterozoic dykes .............. Thermal problems in dyke generation . . . . . . . . . . . . . . . . . . . . . . .
162 164 167 169 169 170 172 173 174 174
................................... .......... .................................. ............. Comparison with continental flood basalts . . . . . . . . . . . . . . . . . Conclusions . . . . ......................................... ........................... .................................... References . . . . . . . . . . . . . . Chapter 5. TECTONIC SETTING O F PROTEROZOIC LAYERED INTRUSIONS WITH SPECIAL REFERENCE TO THE BUSHVELD COMPLEX . . . . . . . . . . . . . . . . . . G. von Gruenewaldt and R.E. Harmer
181
........... Introduction ......................................... Tectonic setting and parental magmas of rift-related Proterozoic layered complexes . . . . . . . . . . ................................ The Great Dyke of Zimbabwe . . . Proterozoic layered intrusio The Jimberlana intrusion, Western Australia FoxRiversill .............................................. .................................. Kiglapait . . . . . . . . . . . . . . . Muskox intrusion ...................................... .......... The Duluth Complex . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ........................... The Bushveld Complex and related magmatic events ............. Geological setting and age relations . . . . . . . . . . . . . . . . . . . . . The Dullstroom Formation ................................ The Rooiberg Group . . . . . . . . . Pre-Bushveld sills . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Syn-Bushveld sills and marg Rustenburg Layered Suite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ............................ The Lebowa Granite Suite . . . . . . ........... The tectonic setting of the Bushveld Complex . . . . . . . . . . . . . . . Discussionsand conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Acknowledgements ............................... .................. References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .............
181 182 182 184 187 188 189 190 192 194 194 198 198 199 200 201 202 203 206 208 208
Chapter 6 . PROTEROZOIC ANORTHOSITE COMPLEXES . . . . . . . . . R.A. Wiebe
215
Introduction ................................................................ Composition and rock nomenclature . . . . . . . . . . . . . . . . . . . . Size and shape of anorthosite massifs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Comparison with Archean and other anorthosites ........................... Distribution ................................................................. Geologic setting ... ......................................... Isotopicages ........................................................................
215 216 217 218 218 219 220
x
Contents
Internal constitution of massif anorthosite complexes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Characteristics of anorthositic plutons and associated dikes . ................... Types of plutons . . . . .............................. . . . . .. . . . Petrographic characteristics of anorthosite plutons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Diapirs .................................. ...................
.............. ........................................ ................... ................. Mineralogy . . . . . . ... .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. .................................. Plagioclase . . . . . . . . . .......... ................................... Olivineand pyroxenes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . High-Al orthopyroxene megacrysts . ... .. . . . Otherminerals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Nd, Sr, and Pb isotopic compositions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Massiveplutons
.......... .......... .......... ................ .............. ........................... ................ ............. .. . .... . . ................ .............. Late basalticdikes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The generation of massif-type anorthosites . . . . . . . ........... .......... Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References ..... ... ......................................................
220 225 225 226 228 230 232 234 235 235 236 231 238 239 244 246 246 251 252 253 253 255 255
Chapter 7 . THE ROLE OF ANOROGENIC GRANITES IN THE PROTEROZOIC CRUSTAL DEVELOPMENTOFNORTHAMERICA ................................. 263 J.L. Anderson and J. Morrison Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Anorogenic magmatism of North America Compositional variations amo Metaluminous, ilmenite-series granites . . . . . . .......... Metaluminous, magnetite Peraluminous granites . . . . . Abundances of U, Th, Zr, an ................................................. InitialSrand Ndisotopes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Oxygen isotopic compositions ......... ............... ........... Conditions of crystallization . . . . . . . . . . . . . . , . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Temperature . . . . . . . . . . . . . . . . ......... Depth of emplacement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Water and oxygen fugacity . . . . .. . . . . Assessment of liquidus temperature based on zircon saturation . . . . . . . . . . . . . . . . . . . . . . . . Anorogenic or epizonal? . . . . . . ..... .. ...... .. ... ...... .. Source of granitic melts .................................................. Origin of anorogenic magmatism: rifts, plumes, and superswells . . . . . . Acknowledgemcn ts ....................................
..........................
..................
263 264 267 268 211 212 213 215 279 282 282 283
284 285 287 288 290 291 291
Contents
XI
Chapter 8. PROTEROZOIC GRANULITE TERRANES . . . . . . S.L. Harley
301
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Modes of occurrence and time-space distribution .................................... Structural styles and sequences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Complexity in structural and metamorphic evolution: the role of reworking . . . . . . . . . . . . . . . . . . Pressure-temperature conditions of Proterozoic granulite terranes ......................... ......... P-T estimation: geothermobarometric approaches and uncertainties The P-T spectrum from GTB . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Some assemblage constraints on Proterozoic granulites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pressure-temperature paths . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Methodology and uncertainties in P-T path determination . . . . . . . . . . . . . . . . . . . . . . . . . . . . Transitional granulite terranes and prograde P-T trajectories .......................... Types of retrograde P-T paths . . . . . . . . . . Rates of metamorphism and cooling: P-T-time constraints ............................ Roles and significance of fluids and melts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fluids in granulite metamorphism: the case from Proterozoic granulites . . . . . . . . . . . . . . . . Migmatites and the role of partial melting in Proterozoic granulites .................... Some remarks on tectonic models for Proterozoic granulites Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References ........
301 302 313 318 320 320 322 324 327 327 329 331 333 335 335 340 344 347 347
Chapter 9. XENOLITHS IN PROTEROZOIC CRUST EVIDENCE FOR REWORKlNG OFTHELOWER CRUST . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . J. Ruiz
361
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Lower crustal xenolith descriptions and localities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Effectsoftransport . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geochemistlyofxenoliths . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geochronologyandisotopicdata . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Comparison of granulite xenoliths and exposed granulite facies Proterozoic crust Evolution of the lower crust in Proterozoic crustal blocks ................................. Acknowledgements .................................................................. References .......... .................
361 362 366 367 371 375 376 377 377
Chapter 10. PROTEROZOIC IRON-FORMATIONS C. Klein and N.J. Beukes
383
.................
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Distribution of iron-formations throughout the Precambrian . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Metamorphism of iron-formation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Major element chemistry of several major iron-formations ..................... Rare earth and trace element chemistry of several major iron-formations . . . . . . . . . . . . . . . . . . . Stratigraphy and sedimentology of depositional basins of iron-formation ....................
383 384 385 392 395 397
XI1
Contents
... Review of recently published iron-formation models . . . . . . . . . . . Paleoenvironmental interpretation of iron-formation deposition in the Transvaal Supergroup, ............................................................ South Africa ...., Paleoenvironmental interpretation of iron-formation deposition throughout Precambrian time . Acknowledgements ..................................................... References . . . . . . . . . . .. ..............................
.
Chapter 11. PROTEROZOIC COLLISIONAL AND ACCRETIONARY OROGENS . . . . B. Windley
Collisional orogens (CO) . . . . . . . . . . . . . . . . . . . . . . . The Grenville
.............................
Accretionary orogens (AO)
The Ketilidian The Penokean . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
.... ... ...... .. Pan-African of the Arabian-Nubian Shield Discussion ........................................................... Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
........................................
401 403 409 411 412
419
419 419 419 422 423 424 427 428 431 433 434 436 437 439 440
Chapter 12. PROTEROZOICTERRANES AND CONTINENTAL ACCRETION IN SOUTHWESTERN NORTH AMERICA . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 447 K.C. Condie
.
Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . TheMojaveProvince . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Yavapai Province . . .................. Dubois terrane Hualapai terrane
........
Pecos terrane
The Grenville Province . . Overlapassemblages . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Idaho Springs-Black Canyon assemblage . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. Wet Mountains assemblage . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cochetopa-Salida assemblage . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Manzano assemblage ...................
441 448 451 452 452 453 454 455 455 455 457 458 459 459 460 462 463
XI11
Contents Franklin Mountains assemblage . . . . . . . . . . . . . . . . . . Discussion . . . . . . Province boundaries . Summary of tectonic settings . . . . . . . . . . . . . . . . . . . . . . . .
............................. Acknowledgments
. .. . .
............................................
...............................................
Chapter 13. ISOTOPIC STUDIES OF PROTEROZOIC CRUSTAL GROWTH AND EVOLUTION . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . P.J. Patchett
............... Introduction ......................................... Global coverage of Nd isotopic data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . ................. Crustal growth curves . . . . . . . . . . . . . . . . Isotopic system stability and reliability of initial Nd isotopic parameters . . . . . . . . . . . . . . . . . Meaning of Nd model ages . . . . . . . . . . . . 1. Granitoids and felsic volcanics, 490 491 ..... .............. Interpretation of ENd values of Proterozoic rocks . . . . . . . Origins of Proterozoic crust . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Proterozoic terranes consisting mainly of recycled older continental crust, 496 - Proterozoic terranes consisting mainly of subduction-related igneous rocks or their derivatives, 498 - Proterozoic terranes consisting of oceanic plateaux and their derivatives?, 499 ............... More rapid crustal genesis in the Proterozoic? . . . . . . . . Acknowledgements . . . . . ................................ References ................................................ ... .................................................................... Subject Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References Index
464 464 464 466 466 469 469 470 471
481
481 481 484 485 488
492 496
500 502 502 509 531
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1
INTRODUCTION KENT C. CONDIE
In the last ten years it has become increasingly clear that plate tectonics has operated on the earth at least since the Late Archean. When I first constructed a version of Figure 1for my Plate Tectonics and Crustal Evolution book in 1980, many of the lines representing lithologic associations in geologic time were shorter or nonexistent. Greenstones were thought to be an Archean phenomenon, the craton association didn’t appear until the Proterozoic, and the collision and ophiolite assemblages were recognized only in the last 1 Gy. We now know that most modern lithologic assemblages appeared first in the Archean and strongly suggest that similar tectonic settings were present on the earth a t that time. However, there are some enigmas: In what tectonic setting did the granite-anorthosite association form and is it really restricted in time to the Proterozoic? Are there different types of greenstones and if so what tectonic regimes do they represent? What processes lead to cratonization? Are there different types of collisional mountain belts and if so, how has their relative abundances varied with time?
-
I ARC
Greenstone
-Quartzite- Pelite- Carbonate
CRATON
- Bimodal ---- Volcanic-Arkose-Conglomerate Mafic Dike Swarms
---
RIFT
High- P Granulite
I COLLISION
Anorogenic Granite - Anorthosite
--
7
4
Ophiolite
--
3
2
I
OCEAN RIDGE 0
AGE ( G o )
Fig. 1. Distribution of lithologic associations with geologic time (after Condie, 1989).
2
K. C.Condie
Why have not any pre-2.0-Ga ophiolites been described yet? Why are massive anorthosites and banded iron formations relatively abundant in the Proterozoic? Is there really a time gap in the generation of continental crust in the Early Proterozoic or a t other times? Clearly the Proterozoic seems to be a key period of time to seek answers to these questions. These and related questions have been the focus of International Geological Correlation Program (IGCP) Project 217, Proterozoic Geochemisq, which came to an end in 1990 after seven years of exciting meetings and field trips. The project brought scientists together from many different disciplines and provided the impetus for a great deal of collaborative research. Between 1984 and 1990, the project sponsored (or partially sponsored) 12 meetings in 11 different countries. At the inaugural meeting in Moscow in 1984, it was decided to focus on nine research objectives as follows: (1) Comparison of Archean and Proterozoic supracrustal assemblages to more fully understand differences between Archean and post-Archean tectonic regimes; (2) To more fully understand the geochemical differences between Archean and post-Archean sediments and to evaluate the various factors that control sediment composition; (3) From combined U/Pb zircon and whole-rock Sm/Nd studies, to see if the apparent 2.4-2.0 Ga continental crust “generation gap” is real; (4) To employ new techniques in the dating of individual zircons to more fully understand Proterozoic tectonic history and the role of crustal reworking; (5) From trace element ratios and Nd isotopic data from basalts, to better understand Proterozoic mantle evolution; (6) To encourage more detailed studies of the anorogenic granite-anorthosite association to better understand its origin and significance in terms of crustal evolution; (7) From combined Nd, Pb, and Sr isotopic data, to more precisely estimate the amount of new continental crust formed during the Proterozoic; (8) To encourage joint P-T and geochronological studies of Proterozoic highgrade terranes to better understand Proterozoic orogenesis; and (9) To try and understand why hydrothermal precious metal deposits are relatively rare in the Proterozoic compared to both the Archean and the Phanerozoic. As a final product of IGCP Project 217, this volume brings together significant advances in our understanding of Proterozoic crustal evolution. Leading experts in their fields have been asked to write each of the chapters. Many of the results are also relevant to other recent IGCP projects such as IGCP 215, Proterozoic Fold Belts and IGCP 257, Precambrian Dyke Swarms. Over the last 10-15 years, it has become increasingly clear that greenstones, once thought to be the earmark of the Archean, are common features in both the Proterozoic and the Phanerozoic. In Chapter 1, Terry Smith summarizes our current knowledge of volcanic rocks from Early Proterozoic greenstones from many locations on five continents. He contrasts the composition of Proterozoic and Archean greenstone volcanics and compares them with modern volcanics
Introduction
3
from various tectonic settings. He concludes from similarities in trace element distributions in continental within-plate basalts, that the processes of formation of the subcontinental lithosphere have been similar since the Late Archean. With exception of low-Mg komatiites, which occur in some Proterozoic greenstones, he shows that lithologic associations and chemical compositions of Proterozoic greenstones are consistent with the existence of modern plate-tectonic regimes. In Chapter 2, Herb Helmstaedt and David Scott address the Proterozoic ophiolite problem and discuss preservation and recognition of ophiolites in Proterozoic terranes. They review and describe the principal occurrences and geochemistry of the well known Pan-African ophiolites, as well as the three known examples of Early Proterozoic ophiolites (Jormua, Purtuniq, and Payson). The authors conclude that Proterozoic ophiolites represent a spectrum of types ranging from small fragments of oceanic lithosphere trapped between juvenile arcs and microcontinents to more extensive ophiolite thrust sheets obducted onto deformed passive continental margins. The diversity of Proterozoic ophiolites is comparable to Phanerozoic ophiolites and probably reflects a similar range of tectonic settings. With the rapid growth of continents in the Late Archean, continental rifts became an important tectonic setting on the continents. In Chapter 3 John Green presents a comprehensive review of Proterozoic cratonic rifts, including a summary of volcanic and sedimentary rocks formed in rifts. He reviews the geochemistry of rift volcanics and their use and misuse in constraining tectonic settings and mantle source compositions. He concludes that rifting culminated at three major times in the Proterozoic, 2.0-1.8, 1.2-0.9, and 0.8-0.6 Ga, with at least the last period resulting in the fragmentation of a supercontinent. Although most Proterozoic rifts seem to be readily accommodated in a modern plate tectonic framework, the mobile belts in Africa and Australia present a problem in terms of their long durations (many 100 My). Mafic dyke swarms seem to be unusually common in Proterozoic cratons and require a large amount of thermal energy to be focused in a small region in the upper mantle for a given swarm. As pointed out by John Tarney in Chapter 4, this energy must be delivered rapidly and then shut-off rapidly to accommodate the short duration of dyke swarm intrusion. In this chapter, Tarney summarizes the occurrence, timing, and petrologic and geochemical constraints on magma production and mantle source composition. He concludes there are two major types of Proterozoic mafic dyke swarms, dolerites and norites. The dolerites may have been derived from oceanic plateau crust that was earlier added to the subcontinental lithosphere, whereas the norites probably come from metasomatized or refractory harzburgite in the subcontinental lithosphere. The Proterozoic and in particular the Early Proterozoic is also the time of intrusion of major layered igneous complexes, of which the grand-dad is the gigantic Bushveld Complex in South Africa. In Chapter 5, Gerhard von Gruenewaldt and Jack Harmer review the major features of Proterozoic layered complexes with emphasis on the Bushveld Complex. They discuss recent data related to the origin and source of parent magmas, relations of layered complexes
4
K.C. Condie
to pre- and post-complex magmatism and associated granites, and the problem of a subduction-zone geochemical signature in some layered-complex magmas. They present a convincing case that most Proterozoic layered complexes were emplaced in rift-like tectonic settings. In Chapter 6 Bob Wiebe presents an excellent review of field, petrographic, structural, and compositional characteristics of Proterozoic anorthosites. Included are state-of-the-art discussions of anorthosite emplacement, contrasts between Proterozoic and Archean anorthosites, characteristics and origin of related rocks (dikes, granites, etc.), and various constraints on magma crystallization and source. One well-established result is that clinopyroxene is an early crystallizing phase from most anorthositic magmas. Isotopic and geochemical results are also reviewed and, when considered with field and petrographic results, seem to indicate that although parent magmas are produced in the mantle, they are commonly crustally contaminated and that associated granites are crustal melts. Also, Wiebe concludes that massive anorthosites are emplaced chiefly in an anorogenic tectonic setting. Although anorogenic granites are found from the Archean onwards, they are particularly characteristic of the Proterozoic. In Chapter 7 Lawford Anderson and Jean Morrison review the role of anorogenic granites in the evolution of Proterozoic crust in North America. Their chapter includes a discussion of both major and trace element compositional variations in Proterozoic anorogenic granites as well as the interpretation of Sr, Nd, and oxygen isotopic data bearing on the origin of these rocks. They also include a summary of temperatures, water and oxygen fugacities, and depths of emplacement of plutons. They conclude that anorogenic granite magmas are produced by small degrees of melting of intermediate meta-igneous rocks or metasediments in the lower crust, and present a model whereby the heat source for melting is mantle plumes. In Chapter 8, Simon Harley presents a very thorough review of Proterozoic granulites. Included is a summary of major occurrences with their structural characteristics and important mineral assemblages from both low- and highpressure assemblages. Approaches and uncertainties in geothermobarometry are discussed together with the role of reworking in granulite terranes. Both prograde and retrograde pressure-temperature-time paths are summarized as are methods and uncertainties in deducing such paths. Other topics included are rates of metamorphism and cooling, the role of fluids, partial melting, and tectonic models. All Proterozoic granulites seem to require an input of extra heat over and above that available from thickened crust and lithosphere during continental collisions perhaps from underplated basaltic magmas. Harley concludes from considering the remarkable diversity and range in structural characteristics, pressure-temperature paths, cooling histories, and fluid/melt relationships in Proterozoic granulites, that it is unlikely a single tectonic model can explain all of them. In Chapter 9, Joaquin Ruiz reviews Proterozoic crustal xenoliths, which occur in both rift-related alkali basalts and in kimberlites and related rocks. The most abundant lower crustal xenoliths are mafic granulites with lesser and variable
Introduction
5
amounts of intermediate to felsic granulites and metasediments. Compared to exposed Proterozoic lower crust, Proterozoic lower crustal xenolith populations are more mafic, record greater equilibration pressures, and generally have greater LIL element depletions. Isotopic studies indicate that many mafic xenoliths represent mantle-derived melts underplated beneath Proterozoic crust during the Phanerozoic. Xenoliths most representative of the lower crust during the Proterozoic appear to be those found in kimberlites. Proterozoic and Late Archean iron formations are discussed in Chapter 10 by Cornelis Klein and Nick Beukes. Included in this rather exhaustive review is the time distribution of the various types of iron formations, a discussion of metamorphism and both major and trace element geochemistry of banded iron formations (BIFs), and a review of stratigraphy and sedimentology of BIFs. As an example of the depositional environment of Early Proterozoic BIFs, the Transvaal Supergroup in South Africa is discussed in detail. The authors present an interesting secular model for the origin of iron formations based on changing hydrothermal input into the oceans and its effect on stratification of seawater. The Late Proterozoic BIFs may be related to episodic glaciations at that time. In Chapter 11, Brian Windley proposes two types of Proterozoic orogens. The first characterized by narrow belts, thickened and highly uplifted crust, extensive reworking of older crust, and in some orogens foreland deformation, results from orthogonal collision of two continental blocks (like India and Tibet in the Tertiary). Oblique collision still produces narrow orogens, but reworking and foreland deformation are minor. The second type, which is produced by multiple collisions of juvenile island arcs (and other oceanic terranes), results in wide orogens with little or no reworking of older crust (like the terranes in Alaska and British Columbia during the late Mesozoic). As an example of Early Proterozoic continental accretion, in Chapter 12 I review the Proterozoic accretional history of the Southwestern United States. Beginning at about 1780 Ma and continuing until 1200 Ma, most of the Southwestern United States (southern Wyoming to West Texas and southern California) was accreted to the Archean Wyoming Craton. Included in the chapter is a summary of the major features of Proterozoic terranes, overlap assemblages, and terrane boundaries. Also, I discuss Proterozoic cratonization and briefly compare the Proterozoic terranes to Phanerozoic Cordilleran terranes, which in many respects are quite different from their Proterozoic ancestors. In the last chapter, Jon Patchett reviews methods and pitfalls of the Sm-Nd isotopic approach to constraining continental growth rates during the Proterozoic. He also clearly discusses ambiguities in the interpretation of Nd model ages and emphasizes that they should be used with caution or, if at all possible, completely avoided in dating rocks. He gives examples of how 6Nd values can be interpreted in terms of remelting of older crust or by mixing of older crustal material with negative E values with mantle-derived magmas. The chapter concludes with a discussion of Proterozoic continental growth rates and the possible role of accretion of oceanic plateaus in explaining periods of rapid continental growth.
6
K.C. Condie
lb improve the quality and completeness of each chapter, the original manuscripts were reviewed by a minimum of three reviewers. The final chapters reflect important revisions suggested by the reviewers. In particular, I want to thank the following individuals who spent considerable time reviewing one or more chapters: Nicholas Arndt, J.H. Berg, Eric Christiansen, Robert Cullers, Yildirim Dilek, Robert Dymek, Lang Farmer, Paul Hoffman, Don Hunter, Yrjo Kahkonen, Randy Keller, Pamela Kempton, Mike Lesher, Suzanne Nicholson, Norman Page, John Pallister, Tim Pharaoh, WC. Phinney, Mary Reid, R.N. Shackleton, and John Valley. REFERENCE Condie, K.C., 1989. Plate Tectonics and Crustal Evolution. Pergamon Press, New York, N.Y., 476 pp.
7
Chapter 1
VOLCANIC ROCKS OF EARLY PROTEROZOIC GREENSTONE BELTS TE. SMITH
INTRODUCTION
Dewey and Windley (1981) suggest that during the Archean up to 85% of the continental crust had formed by 2500 Ma, by accretion of volcanic arcs. However, the structure of Early Precambrian rocks is difficult to interpret because they are commonly highly deformed and metamorphosed and there is a lack of oceanic paleomagnetic and paleontological data to help decipher stratigraphic and structural relationships (Park, 1988). Burke et al. (1976) and Thrney and Windley (1977), believe that Early Precambrian plate tectonic processes differ only in rate, size of plates, and other minor aspects, from those operating at present. However, Baer (1977), citing the inability of eclogite to form in the thinner and warmer oceanic lithosphere, and the apparent lack of ophiolites and low temperaturelhigh pressure metamorphism in the Early Proterozoic record, proposed that modern plate tectonic processes were not operational until the Late Precambrian (see also Wyborn et al., 1987). Continental paleomagnetic studies provide some evidence supporting the concept of plate movement, and the existence of a supercontinent, in the Early Proterozoic, but are inconclusive (Irving and McGlynn, 1981; Morgan and Briden, 1981; Piper, 1982). Past and present heat loss and geothermal gradients (McKenzie and Weiss, 1975; Bickle, 1980; Cartwright and Barnicot, 1987) suggest that the Archean oceanic crust comprised numerous, small, thin, largely komatiitic plates (Amdt, 1983; Nisbet and Fowler, 1983) which are rapidly recycled into the mantle (Bickle, 1980). Continental crust would be created, and the average rate of heat flow would decline, as this process continued, until suitable conditions for the operation of modern plate tectonic processes were established. It has been suggested that there is a fundamental change in the tectonic processes taking place at the Archean/Proterozoic boundary (Xiylor, 1987; McClennan and 'Bylor, 1991). However, current views favour a process similar to Phanerozoic plate tectonics, involving accretion of exotic terranes having a variety of origins, including island-arcs and back-arc basins and microcontinental blocks, in the evolution of the Archean crust (Condie, 1989, 1990; Kerr, 1991). Many recent studies of the geochemistry of Early Proterozoic volcanic suites have been interpreted'using a plate tectonic scenario (see references cited below), but mantle plumes may play a more prominent role in Precambrian tectonic processes than they do in Phanerozoic plate tectonics (Kerr, 1991).
8
TE. Smith
METHODS
This review is not meant to be a comprehensive description of Early Proterozoic volcanic activity. The geology and geochemistry of selected suites are used to demonstrate that they erupted in thermotectonic environments similar to those of modern volcanic suites, implying that Early Proterozoic and Phanerozoic tectonic processes are not fundamentally different in nature. Mid-ocean ridge basalt (M0RB)-normalized multi-element plots (Wood et al., 1979) have been used by many authors (see below) to identify the thermotectonic environment of eruption of Early Proterozoic (2500-1600 Ma) volcanic suites by comparison with the characteristic signatures of modern suites. These diagrams are preferred to binary and ternary discriminant diagrams (Pearce and Norry, 1979; Pearce, 1982; Pearce et al., 1984; Perfit et al., 1981; Sun et al., 1979; Sun, 1980; Wood, 1980; Wood et al., 1979) because they apply to a range of rock types not just basalts, are commonly available in the literature, facilitate visual comparison with the standard suites, allow examination of numerous inter-element ratios simultaneously, contain petrogenetic information and can be adequately described in terms of elemental concentrations and ratios. In this study the diagrams have been used to compare the trace element compositions of Early Proterozoic volcanics with the average compositions of modern suites compiled from the largest and most complete data sets available, (Ewart, 1979, 1982; Bailey, 1981; Condie, 1985; Holm, 1985; Marsh 1987; B b l e 1). All enrichments and depletions of various elements referred to in this paper are relative to their concentrations in MORB. Some of the important geochemical parameters of the Early Proterozoic volcanic suites discussed are summarized in Tables 2, 3 and 4. The multi-element plots are reproduced where they are not available in the original publications and/or where the conclusions drawn here differ from those of the original study. The majority of Early Proterozoic volcanic suites have undergone metamorphism and/or metasomatism, involving chemical alteration. ,5202, FeO* (total Fe expressed as FeO), Ti02, P2O5, high field strength elements (HFSE) (Zr, Nb, Y), and the rare-earth elements (REE) are generally considered to be relatively immobile, and Al203, MgO, CaO, N a2 0 , K20, and the large ion lithophile elements (LILE, Rb, Ba, U) are considered to be relatively mobile (Condie, 1981; Beswick and Soucie, 1978; Gelinas et al., 1982; Ludden et al., 1982; Pearce, 1983; Pharaoh et al., 1987). Thus, in this paper emphasis is placed on the use of the light rare-earth elements (LREE), HFSE, and heavy rare-earth elements (HREE) to classify Proterozoic volcanic suites. The use of the average concentrations and ratios of LILE, especially Th, to identify suites of altered subduction-related volcanics (Pearce, 1983; Pharaoh and Pearce 1984; Condie, 1986, 1987, 1989; Pharaoh et al., 1987) is considered inappropriate because most of the LILE are mobile. In addition, enrichments in LILE (and Th) may occur in continental tholeiites and may be absent in island-arc tholeiites. Ba, K, and Sr spikes on multi-element diagrams are diagnostic of
9
Volcanic rocks of Early Proterozoic greenstone belts TABLE 1
Average morb-normalized elemental ratios and contents of selected Phanerozoic basalts and andesites
Orogenic basalts Low-Ka Calcalkalinea High-Ka Shoshonitic a BAT BAO BAC
(b/Yb)N
(La/Nb)N
Nb (ppm)
1.31 1.46 1.17
1.20 4.64 15.18 11.62 2.01 2.20 1.84
2.01 2.29 3.25 4.16 2.13 3.53 1.55
1.7 5.5 19.5 1.3 3.2 4.5 2.3
17 21 26 20 27 29 24
1.66 1.89 2.69 3.45 1.76 1.29 2.95
1.62 6.55 21.53 51.14 1.07 5.73 7.59 17.31
2.82 3.93 5.76 9.99 4.48 2.83 2.16
1.4 5.3 12.9 16.4 0.8
24 26 21 26 29 24 22 12
3.30 3.25 4.76 8.26 3.70 2.34 1.79
2.43 0.91
4.52 1.00
2.01 1.00
2.01 2.04 2.18 3.86
6.34 6.34 5.00
0.80 0.95 1.10 2.03
7.8 4.3 3.5 20.7 20.2 14.6 12
30 35 30 29 31 31 24
1.66 0.83 1.24 0.66 0.79 0.91 2.38
2.10 2.90 3.11 3.78
5.69 5.53
0.48 1.10 2.38 1.53 0.47 0.67
65 47 28 23 57 110
0.57 1.32 2.19 1.79 0.60 0.78
Andesites Low-Ka Calwkaline a High-Ka Shoshonitic a hw-KC OA CIA MC
Nonorogenic basalts CT OFTb
N-MORB E-MORB OIT
IRT Albin
Andesites Iceland Grande Ronde Karoo Tusas Anorogenica Ethiopia8
Y (ppm)
(La/Sm)N
12.59
5.1 9.6
84 17 5 14 40
La/Nb
a Ewart (1982); Holm (1985); Bailey (1981); Pearce (1983); Marsh (1987); Gill (1981); Basaltic Volcanism Study Project (1981).
orogenic basalts (Sun, 1980) but are not generally seen in Proterozoic basalts. Subduction-related basalts generally have La/Nb > 2.0, and Y < 20 ppm, and basalts erupted in an extensional environment generally have La/Nb < 1.0 and Y > 20 ppm (Thompson et al., 1983; Winchester et al., 1987; Lees et al., 1987), but there are many exceptions to this observation (Table 1).The chemical data base available for Early Proterozoic volcanic suites is limited, and often incomplete. In addition, both Early Proterozoic and modern volcanic suites have complex petrogenetic histories which affect their elemental concentrations and
10
TE. Smith
ratios (Ewart, 1979, 1982; Le Roex et al., 1983; Holm, 1985). Thus it is not possible to show any systematic differences between Early Proterozoic and modern trace element signatures of volcanic rocks. The geochemical data should be supplemented by geological information, when attempting to identify the eruptive environment of Early Proterozoic volcanic rock suites, because of the difficulties in distinguishing chemically between sequences of continental flood basalts (CFB) and the calcalkaline basalts (CAB) of volcanic arcs (Thompson et al., 1983; Arculus, 1987; Duncan, 1987; Marsh, 1987). The range of volcanic compositions, dominant flow type, presence or absence of an iron enrichment trend, phenocryst assemblages, facies of associated sediments, nature of basement and its relationship to the supracrustal sequence, can all be helpful when identifying the nature of a volcanic suite and the tectonic environment in which it erupted (particularly in the Precambrian) (Pearce, 1987; Smith and Holm, 1987). For example, continental tholeiitic suites (CT) may be entirely mafic or bimodal (mafic-felsic) and rarely have associated andesites (Hall, 1987), whereas arc-related sequences generally contain a complete range of rock types from basalt to rhyolite, including abundant andesites (Ewart, 1982). In general, there are few basalts and abundant dacites and rhyolites in the continental-margin arc suites and the opposite is true of the oceanic arc suites (Ewart, 1982). Although commonly regarded as a single tectonomagmatic entity, CFB may erupt in a variety of tectonic settings. For example, basalts erupt in the Karoo Province where continental rifting leads to ocean floor spreading (Etendeka suite, Namibia), where stretching and rifting of the crust does not lead directly to ocean floor formation (Lebombo suite, southeastern Africa), and in an a-tectonic within-plate setting (Lesotho suite). The basalt geochemistry is not diagnostic of the tectonic setting but the rift-related environments may be identified by the petrology and geochemistry of the whole igneous suite. A comparison (Marsh, 1987) of the Karoo CFB with other well known volcanic suites shows that uncontaminated CFB may be chemically similar to either ocean island tholeiite (OIT) or to average CT/Type 2 CFB (Holm, 1985; Thompson et al., 1983). The Type 2 CFB’s are distinguishable from modern arc-related basalts by their low Sr content (Marsh, 1987). The association of flood basalts (Type 2 CFE3 or OIT) with abundant silicic magmatism and late stage dykeshavas having more MORB-like geochemistry, is indicative of a rift regime. The earlier enriched basalts (Type 2 CFB or OIT) may originate in the subcontinental lithosphere and the later depleted basalts (MORB) are of asthenospheric origin (Marsh, 1987).
EARLY PROTEROZOIC GREENSTONE BELTS
The Baltic Shield
The majority of current views suggest that most of the Baltic Shield developed, during the period from 3.5 to 1.5 Ga, in four westward younging orogenic events.
11
Volcanic rocks of Early Proterozoic greenstone belts
The central part of the Baltic Shield was formed by growth of new continental crust during the Svecofennian (2.0-1.75 Ga) and Gothian (1.75-1.55 Ga) orogenies (GaAl and Gorbatschev, 1987). These two events are gradational but each is characterised by large volumes of mantle-derived tonalite-granodiorite intrusions which were emplaced 100-150 Ma apart. Nd isotope studies (Patchettt and Arndt, 1986) indicate that >SO% of the 1.9-1.7 Ga crust in Scandinavia is newly derived from the mantle, and almost certainly represents a major subduction-related mantle-to-crust differentiation. Later orogenies contributed IittIe new material to the shield but involved intense crustal reworking especially during the Hallandian (1.5-1.4 Ga), the complex multi-stage Sveconorwegian-Grenvillian (1.25-0.9 Ga) and Caledonian (0.6-0.4 Ga) orogenies (Gorbatschev and GaAl, 1987; GaAl and Gorbatschev, 1987; Wilson et al., 1987). The principal subdivisions of the Central Baltic Shield, within which three Early Proterozoic tectonostratigraphic packages can be recognized (Pharaoh and Brewer, 1990, table l), are shown in Fig. 1. The oldest tectonostratigraphic package is made up of the Lapponian Supergroup plus the Jatulian (Pharaoh and Brewer, 1990). It is characterized by submarine volcanic sequences, associated with, or enclosed in, epicontinental sedimentary sequences which are exposed in several structural windows and greenstone belts within the Karelian Province. Those for which chemical data are
AN
It
POST-SVECOKARELIAN PROTEROZOIC TRANSSCAM)WVUN GRANITE-PORPHYRY BELT AM) RAPAKIVI MASSIFS
$=-
5SOUTHWESTSCANQHAVUN OROGEN
.*
SVECOKARELIAN DOMAIN ! A M (UIANULrTE BELT AND ASSOCUTED ROCKS
@KARELIAN PROVINCE m N O R l H SVECOFENNIAN SUBPROVINCE
@SOUTH
SVECOFENNIAN SUBPROVINCE
‘W
.“y Ez,
I
YO_ OLL
CENTRAL SVECOFENNUNSUBPROVINCE
2
ARCHAEAN OOMAIN
0
KARELIA AND KOLA PENINSULA PROVINCES
‘
BELOWORIAN PROVINCE
-
LATE ARCHAEAN OF THE LAPPONIA GROUP (REWSTONE BELTS IN KARELIA AN0 LAPLAND
M
300 KM
a
MAJMI FAULTSAND WCTLE SHEAR ZONES
PZ
Fig. 1. Schematic geological map of the Baltic Shield (from Gorbatschev and Gail, 1987). LBB = Ladoga-Bothnian Bay Tectonic Zone, LWST = Lapland-White Sea Thrust Fault, MZ = Mylonite Zone, PBTZ = Pechenga-Varzuga Tectonic Zone, PZ = “Protogine” Zone.
TE. Smith
1-
Caledonian cover and post-Svecokarelian units
0
Metasediments etc
Z E : z a n i c suite
a
:ZSt:es
0
Gneiss etc
G;anulite
1
Karelian domain
1
Arzhaean craton
n
Svecotennian domaln
Arc volcanlc suite
-
>Major thrust
J
-
Shear zone Southern limit of isOtopiCallY defined Archaean crust
Fig. 2. Sketch map of the northern part of the Baltic Shield showing the distribution of Early Proterozoic greenstone belts (from Pharaoh and Pearce, 1984). KuB = Kautokino belt, AKW = Alta-Kvaenangen window, AW = Altenes window, RW = Repparfjord window, KrB = Karasjok belt, KiB = KirunaVittangi greenstone belt, PG = Petsamo Group, PC = Pechenga Complex, SG3 = Kola Superdeep Borehole.
Volcanic rods of Early Proterozoic greenstone belts
13
available include the volcanic rocks of the Lower Holmvatn Group (RW, Fig. 2) and the Kviby Group (AW, Fig. 2) which are assigned on lithostratigraphic grounds (Pharaoh et al., 1987, fig. 2) to the Lower Lapponian (2.50-2.45 Ga), and the Upper Holmvatn Group (RW, Fig. 2) which occurs at the base of the Upper Lapponian (ca. 2.40 Ga) (Pharaoh and Brewer, 1990, table 1). The Nussir Group (RW, Fig. 2), the Cas'kejas Group (KaB, Fig. 2), the Bakkilvarri Formation (KrB, Fig. 2), and the Kiruna and Vittangi Greenstone Groups (KiB, Fig. 2) also occur in the Upper Lapponian (>2.0 Ga) (Pharaoh and Brewer, 1990, table 1). The Pechenga Complex (PC, Fig. 2) and Petsamo Group (PG, Fig. 2) are separated from the Karelian suites by the Lapland Granulite Complex (Fig. 2), but are of approximately the same age ( ~ 2 . 0Ga), (Hanskii et al., 1990; Pharaoh and Brewer, 1990, table 1). The Lower Holmvatn Group consists of mafic, intermediate, and acid metavolcanics interbedded with clastic sediments of mixed volcanigenic and continental provenance. It is overlain by the Upper Holmvatn Group which is comprised of metamorphosed tholeiitic pillow basalts. The Lower Holmvatn Group, the Upper Holmvatn and Kviby Groups have LREE enrichment and negative Nb anomalies (Tdble 2). They are similar to calcalkaline suites formed in continental margin arcs (Pharaoh and Pearce, 1984, table 1, fig. 5; Pharaoh et al., 1987, table 1, fig. 4), and Albin basalts and intermediate rocks of the Etendeka area in the Karoo (Erlank et al., 1984, table 1, averages 4 and 5; Fig. 3, Table 1).Their age relationships are uncertain and they may represent either a sequence of arc volcanics (Pharaoh et al., 1987, fig. 10; Pharaoh and Brewer, 1990) or a sequence of within-plate basalts similar in composition to CFB erupted on rifted continental crust. The Nussir Group comprises a monotonous series of subalkaline basalts, with no andesites or felsic rocks. They show strong iron enrichment and have a considerable range of Ti02 content (0.74-3.52). The majority of the basalts have negative Nb anomalies and LREE-enriched multi-element patterns (Pharaoh and Pearce, 1984, table 1, fig. 7; Pharaoh, 1985; Table 2). These are identified as within-plate basalts (Pharaoh et al., 1987) and are similar to Type 2 CFB (see Marsh, 1987).
0
.
1
A829 "
>
Rb Ba Th
+
,
K
8
R464
-
r , , Nb La Ce
4
8
*5
Sr P Zr Srn
Ti
Y Yb
Fig. 3. MORB-normalized multi-element plot of Lower Holmvatn basalt (R4674) and andesite (A829), and average Albin basalt ( 4 ) and intermediate rock (5). Normalization values from Pearce (1983).
TE. Smith
14 TABLE 2
Typical morb-normalized elemental ratios and contents of selected Early Proterozoic basalts and andesites from the Baltic Shield (b/Sm)N L. Holmvatn a L. Holmvatn Kviby U. Holmvatna Nussir a Cas’jekas a Bakillvarri 1 Bakillvarri 1a Bakillvarri 2‘ Bakillvarri 2a Vittangi a Luostarin Luostarin Nikel a Nikel Ostrobothnia a Jormua a Outokumpu Skellefte a Skellefte Ulangera Nagu-Korpo a Kiruna a Arvidsjaur a Loch Maree A a Loch Maree B a
&a/Yb)N
4.32 4.87
5.32 12.46
5.71 4.5
7.45 3.35
0.68 0.85 2.79 2.30 1.50 4.11 4.85 1.99 5.46 1.46 zl.0 2.58 3.24 0.89 1.61 3.18 3.06
0.69 0.86 5.63 5.56 1.55 8.69 10.55 2.78 11.85 0.58 1.92 z1.0 5.41 7.21 1.33 4.34 6.56 9.37
1.22 2.35
1.80 4.31
0.99
(b/Nb)N 1.74 2.14 2.17 2.68 1.23 0.15 0.32 0.64 1.17 0.90 8.94 23.66 2.51 9.76 0.43
Nb (ppm) 5 12 13 7 6 7.5 7 6 8 15.5 5 1.9 1.7 3.2 3.5 4.0
z1.0 2.96
z 2 10
0.68 0.59 2.02 2.00
5 12 7 7
0.86 2.29
5 6
Y (ppm) 14 24 22 25 19 25 19 24 11 34 22
22 z21 24 18 22 21 27 17 21 30
La/Nb 2.04 2.82 2.60 3.14 1.44 0.18 0.38 0.75 1.37 1.06 10.47 27.68 2.94 11.42 0.50
z1.2 3.46 0.80 0.85 2.37 2.34
1.01 2.68
basalt; andesite; komatiite; trachyandesite. 1 = depleted; 2 = enriched. a
The Cas’kejas and Bakkilvarri Formations are underlain by shallow marine
or fluviatile quartzo-feldspathic clastic sediments and overlain by shallow marine dolomitic, pelitic and coarse clastic sediments. They are lithostratigraphically correlated with the Nussir Group. The Cas’kejas Formation is a monotonous series of metabasaltic lavas, diabase dykes and tuffs. Very little geochemical information is available for this unit but the incomplete data which are available (Pharaoh et al., 1987, table 1) suggest that the unit is similar in composition to the Nussir Group (Tmble 2). The sedimentary environment and chemistry of the volcanic rocks suggest that this suite also erupted in a within-plate environment (see also Pharaoh et al., 1987, p. 50). The Bakkilvarri Formation, can be divided into two groups, each comprising subordinate komatiites (< 10%) and tholeiites, on the basis of their trace element signatures (Pharaoh et al., 1987, table 1, figs. 6, 7, and 8). The rocks within each group may be related by different
Volcanic rocks of Early Proterozoic greenstone belts
15
100
g
10
0
I
5 $
1
0.1
Rb Ba Th K Nb La Ce Sr
P Zr Sm Ti Y Yb
Fig. 4. MORB-normalized multi-element plots of the enriched komatiite (83/76) and high-Fe tholeiite (83/74) and depleted komatiite (81/73) and high-Mg tholeiite (81/75) of the Bakkilvarri Formation.
degrees of partial melting and/or fractional crystallization (Pharaoh et al., 1987). Thin quartzo-feldspathic schists, associated with the komatiites, may represent former felsic volcanics. The first group comprises incompatible-element depleted komatiitic basalts and Mg-tholeiites with low concentrations of trace elements, positive Nb anomalies, and LREE-depleted multi-element patterns (Fig. 4, Table 2). The second group, comprising basaltic komatiites and Fe-tholeiites, has higher overall concentrations of trace elements, generally without Nb anomalies, and LREE-enriched multi-element patterns (Fig. 4, B b l e 2 ) . The signature of the Mg-tholeiites of the depleted group is very similar to typical N-MORB, while that of the enriched tholeiites lies in the range shown by CFB which lack negative Nb anomalies (OIT type, Marsh, 1987). The Karelian supracrustal sequences of North Norbotten include the metabasalts of the Vittangi/Kiruna greenstone belt. By analogy with the Cas'kejas Formation the Vittangi/Kiruna greenstone belt is inferred to lie unconformably on Archean basement, which occurs north and east of Kiruna but has not been detected further south. The basalts of the Vittangi sequence have no significant Nb anomalies, and are characterized by slightly LREE-enriched patterns (Pharaoh and Pearce, 1984, table 1, fig. 5; Qble 2) and may be compared to the OIT flood basalt suite of Marsh (1987). Pharaoh and Brewer (1990) assign the Vittangi/Kiruna greenstone belt to the Nordkalott continental tholeiitic province, which they compare to the Karoo. The narrow Pechenga-Varzuga and Lapland Granulite belts (Fig. 2) form part of the Belomoride belt, which may be a collision orogen resulting from the convergence the Archean Inari-Kola Craton to the east and the Archean South Lapland-Karelia Craton to the west ca. 2.0-1.9 Ga (GaAl and Gorbatschev, 1987; Pharaoh et al., 1987). The Pechenga Complex (Petsamo Group of eastern Finnmark) comprises two metavolcanic series (the Luostarin Series and the Nikel Series) and associated metasediments including conglomerates, marbles, shales, and sandstones, deposited on rifted Archean crust of the Inari-Kola Craton (Kozlovsky, 1984; Pharaoh et al., 1987). The age of the complex is poorly established (Luostarin Series K-Ar ages 2130-1570 Ma, most 252100 Ma, Pharaoh
TE. Smith
16
-
0 . 1 ' I Rb Ba Th 8
1
I
N10
+
I
I
CAB
I
+
I
I
LES
I
1
I
K Nb La Ce Sr P Zr Sm Ti Y Yb Fig. 5. MORB-normalized multi-element plot of average Nikel basalt (N10) and average calcalkaline basalt (CAB) and average central Lesotho basalt (LES). ~
et al., 1987), and the Pilgujarvi Suite (Upper Nikel Series, Sm-Nd isochron age of 1990 k 66 Ma, Hanski et al., 1990), and its correlation with Nussir and Cas'kejas Formations of W. Finnmark is uncertain (Pharaoh and Brewer, 1990, table 1). The Luostarin Series comprises interbedded, abundant basaltic-andesite and trachy-andesite lavas with moderately LREE-enriched patterns and strong negative Nb anomalies (Table 2). The overlying Nikel Series is made up of predominantly, massive and pillowed picritic metabasalts and tuffs, with relatively few basaltic andesites, which all have strongly enriched LREE patterns and deep negative Nb anomalies (Fig. 5). The Nikel basalts have low Sr like CFB (Marsh, 1987) but also show Z r depletion like arc related volcanics. These lavas were initially interpreted as arc-related volcanics (Pharaoh et al., 1987, fig. 9; Pharaoh and Brewer, 1990, fig. 2) but Hanski and Smolkin (1989) considered that they were products of intra-continental rift eruption. The metasedimentary sequence, eruptive environment and presence of ferropicrites and komatiites in the Pechenga-Varzuga belt (Hanski and Smolkin, 1989) favours an intracontinental rift environment. In addition, the basalts of the Nikel series have trace element signatures similar to those the average basalt of central Lesotho in the Karoo (Marsh, 1987, table 1, analysis 1; Tible 1). Also, with the exception of the large negative Nb anomalies, the trachybasalts and trachyandesites of the Luostarin Series are chemically similar to the intermediate rocks erupted in the Snake River area and in the Ethiopian rift at Boina (Table 1). Pharaoh and Brewer (personal communication, 1991) now believe that this group is dominantly a continental tholeiitic suite. The Kalevian tectonostratigraphic suites (z1.96 Ga) occur, together with the younger Svecofennian suites (1.92-1.87 Ga), in the Svecofennian Province, along the southwest edge of the Archean craton (Fig. 1). The shelf clastics of the Jatulian type occur to the northeast and younger, turbiditic, deep water sediments occur to the southwest of this boundary. The Kalevian volcanics are mainly tholeiitic nietabasalts, and those for which chemical data is available occur in the Kiiminki belt (Ostrobothnia), the Jormua Ophiolite and Outokumpu Allochthon, of north central Finland (Fig. 1, Pharaoh and Brewer, 1990, table 1).
Volcanic roch of Early Proterozoic greenstone belts
17
Marine metasediments and metavolcanic rocks, the youngest of which are 1.9 to 2.1 Ga (Honkamo, 1987), overlie the Archean basement unconformably in the Ostrobothnia, Finland. Coarse metarkoses and conglomerates predominate to the southeast, and metavolcanics associated with turbiditic metagreywackes and micaschists occur further north. The dominantly basaltic rocks, which include pillow lavas, pillow breccias and hyaloclastite in the east (Martimo) and centre (Kimiinki), and massive basalt lavas and tuffs in the west (Haukipudas), all have very similar geochemical signatures, although the Haukipudas rocks have higher trace element abundances (Honkamo, 1987, table 1). They are characterized by and LREE enrichment relative to typical N-MORB (Honkamo, 1987 slight Th, B, fig. 9; Table 2) but are chemically similar to OFT (Bble 1). The Ti/V contents of the metavolcanics, and the nature of the associated sedimentary rock facies, suggest that they may have been emplaced in a back-arc basin (Honkamo, 1987). The 1.96 Ga old Jormua Complex (Kontinen, 1987) is a convincing example of an Early Proterozoic ophiolite. It occurs in a suture zone along which the Svecofennian Orogen was accreted to the craton (Park, 1985), and is tectonically associated with Jatulian meta-arenite and Lower Kalevian conglomerates, micaschists, phyllites and iron formation (Kontinen, 1987). The complex consists of a basal unit of serpentinites, intruded by gabbros, which contain minor trondhjemite segregations, a middle unit of mafic dykes and an uppermost unit of basaltic pillow lavas. The gabbros are cumulative and probably co-genetic with the dykes and lavas some of which have slight LREE enrichment (Bble 2) and are chemically similar to BAT The chemistry of the trondhjemites is very similar to the granitic rocks of mid-ocean ridges and high-Ti ophiolites (Kontinen, 1987, table 1). The Outokumpu Allochthon is similar to the Jormua Complex (Park, 19SS), however, the mafic volcanics are less abundant at Outokumpu and have been strongly deformed and hydrothermally altered. Unlike the mafic rocks of the Jormua Complex they are strongly enriched in LILE, including Th. They have lower HFSE and higher Cr contents than MORB and may be more comparable to island-arc or back-arc basalts (Pharaoh and Brewer, 1990, fig. 4). However, their (La/Sm)N, (La/Yb)N, (La/Nb)N ratios are very similar to those of N-MORB (lhble 2). The third tectonostratigraphic package is represented by the Svecofennian volcanics which erupted between 1.92-1.87 Ga, (Welin, 1987). The volcanics are associated with abundant of turbiditic sediments and they range in composition from early primitive oceanic arc-like rocks, to felsic calcalkaline volcanics and plutonics (Pharaoh and Brewer, 1990). The Svecofennian volcanics are concentrated in two areas, the Skellefte Field in the North Svecofennian Subprovince and the Bergslagen Field in the South Svecofennian Subprovince. These two areas are separated by the Bothnian Basin, a wide area of metasedimentary rocks which forms much of the Central Svecofennian Subprovince (Figs. 1 and 2). The Skellefte Field is the smaller of these two areas and occurs adjacent to the Archean craton margin, but no evidence of pre-existing gneissose basement has been found. The larger Bergslagen Field is represented by a series of “leptite” volcanics outcropping in
TE. Smith
18
south central Sweden and S. Finland (Pharaoh and Brewer, 1990). The Norbotten Porphyry Arc occurs in the North Svecofennian Subprovince and is comprised of the Arvidsjaur and Arjeplog porphyries and the Kiruna Porphyry Group (Fig. 2), which are younger (1.89-1.87 Ga) and less deformed than the Skellefte Group. These porphyries are probably co-magmatic with synorogenic intrusive suites such as the Haparanda Series (Pharaoh and Brewer, 1990). In the North Svecofennian Subprovince detailed geochemical data is available for the volcanic rocks of the Skellefte Group, a deformed and metamorphosed, Early Proterozoic volcano-sedimentary sequence which contains pyriterich, stratabound, massive sulphide deposits (Vivallo and Claesson, 1987). The sequence was intruded by granites in the period 1890-1760 Ma, and may be co-magmatic with those of the Jorn group (1890 Ma, Wilson et al., 1985). It comprises abundant basalts, including some MgCr-rich basalts in the Boliden-Langdal area thought to have affinities to komatiites or boninites, moderately abundant andesites, less abundant submarine pyroclastic dacites, and very abundant welded, submarine pyroclastic rhyolites (Vivallo and Claesson, 1987, fig. 3, table 1). The mafic rocks are believed to be younger than the majority of the felsic rocks (Pharaoh and Brewer, 1990). The basalts (Vivallo and Claesson, 1987, fig. 6) are LREE enriched and are characterized by negative Nb anomalies (Pharaoh and Pearce, 1984, table 1, fig. 5; Vivallo and Claesson, 1987; Table 2). The MgCr-rich basalts of the Boliden-Langdal area have higher Si02 and Al203, and lower MgO than komatiites, and lower SiOz than most typical boninites. They are similar in major and trace element chemistry to picritic basalts from oceanic island-arcs (see olivine microphyric basalts, Caldwell et al., 1984, table 2; Fig. 6 ) . The andesites and dacites all show multi-element patterns which are very similar to those of the basalts (Xtble 2). The andesites have the general compositional features and geochemical signatures of orogenic andesites formed at continental margins (CIA, Bailey, 1981). The rhyolitic rocks have similar trace element patterns to the andesites and dacites, but notably higher concentrations of incompatible elements. Vivallo and Claesson (1987) suggest that most of the mafic volcanics erupted dur-
-
0.1
MgCr bas
Rb Ba Th K Nb La Ce S r
-
OMB
P Zr Sm Ti
Y Yb
Fig. 6 . MORB-normalized multi-element plots of the high-MgCr basalt (MgCr bus) and a typical olivine microphyric basalt (OMB) o r picrite from Carriacou, Grenadine Islands, Lesser Antillean Arc (Smith and Thirlwall, unpublished data).
Volcanic rocks of Early Proterozoic greenstone Belts
19
ing the formation of a continental-margin arc or mature oceanic arc, followed by crustal extension and the extrusion of a large volume of felsic pyroclastic material. The Central Svecofennian Subprovince (Norrland Geosyncline or Bothnian Basin) is characterized by the occurrence of minor mafic volcanics and leptites in a dominantly migmatised greywacke succession (Harno Formation). One comparatively complete analysis is available for a mafic rock from Ulldnger (Pharaoh and Pearce, 1984, table 1, fig. 5). The rock has relatively flat R E E patterns and a slight positive Nb anomaly (Table 2) and is generally similar to average OFT (Holm, 1985). The Southern Svecofennian Subprovince is dominated by felsic volcanics with minor mafic rocks. The Bergslagen Field represents the largest area of eruptive rocks in the Subprovince. Its stratigraphy is poorly understood, and very little trace element data is available for the eruptive rocks (“leptites”). The major element chemistry of the “leptites” shows that they form a bimodal suite dominated by rhyolites and deficient in andesites. Minor amounts of basalt erupted throughout the area as the last phase (Lagerbland and Gorbatsev, 19%). The occurrence of stratiform sulphide, and manganese and iron ores, and associated greywackes and subgreywackes, together with the nature and compositional range of the plutonic intrusions, and the chemistry of the “leptites” suggests that they represent an island-arc suite of calcalkaline rocks (Liifgren, 1979). A sequence of within-plate tholeiitic metavolcanics occurs in the Nagu-Korpo area of SW Finland, in the Southern Svecofennian Subprovince, and is slightly LREE enriched with positive Nb anomalies, and resembles OIT in composition (Ehlers et al., 1986, fig. 9; Table 2). The Early Proterozoic Kemio-Orijarv-Lohja volcanic belt occurs to the west of the Nagu-Korpo belt and contains a sequence typical of the widespread calcalkaline suites which erupted after the Nagu-Korpo rocks. It consists of a sequence of mafic and felsic metavolcanics and metagreywackes intruded by syngenetic gabbro-tonalite bodies (Colley and Westra, 1987). The subalkaline to alkaline volcanic sequence begins with a submarine sequence of mafic and intermediate lavas, overlain by metaturbidites, and then felsic pyroclastics. The geology and geochemistry of the sequence suggests that it was formed where back-arc rifting aflected both oceanic and mature arc crust (Colley and Westra, 1987). This sequence is typical of the widespread calcalkaline suites which erupted after the Nagu-Korpo rocks. The Tampere Schist Belt (1904f4 to l 8 8 9 f 5 Ma, Kahkonen et al., 1989) is about 150 km to the north of the Nagu-Korpo area. It comprises more than 50% andesites-dacites-rhyolites ranging from (almost) low-K tholeiitic, to shoshonitic and trachytic in composition, but intermediate rocks of calcalkaline affinities are most common. Subordinate units with high Ti02 contents (non-arc affinities) are also found (Kahkonen, 1987). The suite resembles those of mature island arcs or volcanic arcs formed at, or close to, active continental margins (Kahkonen (1987). The Norrbotten Porphyry Arc (1.89-1.87 Ga) is a distinctive belt of intermediate and felsic lavas which cross cuts from the northern part of the Svecofennian Domain onto the margin of the Archean craton (Pharaoh and Brewer, 1990). It includes the mafic and felsic rocks of the Kiruna Porphyries, together with the
20
TE. Smith
porphyries of the Arvidsjaur and Arjeplog districts (Fig. 2). The Kiruna porphyries (1.9 Ga, Skiold and Cliff, 1984), which include metamorphosed rhyolitic and trachytic lavas and tuffs with subsidiary basalts and andesites, overlie the Kalevian Vittangi greenstone sequence in the north and thicken to the south where the Vittangi Group is absent. The Kiruna, Arvidsjaur and Arjeplog porphyries show strong enrichment of the LREE and a relative depletion in Nb and 7h (Pharaoh and Pearce, 1984, table 1, fig. 1; Pharaoh and Brewer, 1990, fig 6; a b l e 2). They resemble the rocks of modern Andinotype high-K calcalkaline arcs and are considered to be co-magmatic with the synorogenic Haparanda Series (Pharaoh and Brewer, 1990). The earliest Proterozoic volcanic suites of the Baltic Shield (Lower Holmvatn and Kviby Groups) have been little studied but may be calcalkaline in composition. Their magmatic signatures rnay be inherited from a (Late) Archean phase of arc magmatism or may reflect the earliest Proterozoic subduction episode (Pharaoh and Brewer, 1990). The later Lapponian and Jatulian volcanic suites show a systematic distribution indicative of crustal extension and rifting. Continental tholeiitic sequences occur in the west, (Repparfjord and Altenes windows, Kautokeino, Vittangi and Kiruna greenstone belts), and are accompanied by MORB-like basalts and komatiites in the Bakkilvarri Formation of the Karasjok greenstone belt. None of the sequences are associated with significant volumes of felsic rocks. Pharaoh and Brewer (1990) have suggested that these suites erupted in environments similar to those of the Karoo. The compositional variation within and between the volcanic rocks of different areas of the Lapponian and Jatulian rocks confirms the heterogeneity of the sub-continental lithosphere on both a local and regional scale. The Kalevian volcanic rocks of Ostrobothnia and the Jormua Ophiolitic Complex have chemical compositions similar to BAT and MORB. They formed at a divergent plate margin during the opening of an ocean basin, or in a marginal basin adjacent to a Proterozoic ocean, in the break-up of the Archean craton following the Lapponian and Jatulian extensional events. The Jormua Complex was tectonically emplaced in its present location during the compressional stage of the Svecofennian Orogeny. The environment in which the volcanic rocks of the Outokumpu Allochthon erupted is uncertain but may have been a primitive oceanic arc or in a back-arc basin (Pharaoh and Brewer, 1990), which marked the onset of subduction in the area. The early Svecofennian volcanics of the Skellefte, Bergslagen, Kemio-OrijsrviLohja, and Tmmpere belts represent mature oceanic arcs and/or continental margin arcs, and back-arc basins. Subordinate amounts of high-Ti02 within-plate tholeiitic basalts occur indicating rifting of the arc crust. The presence of the within-plate basalts (WPB) and back-arc basins, together with the width of the belt (ca. 800 km), suggests these suites record the accretion of more than one arc to the continental margin (Pharaoh and Brewer, 1990). The less deformed lavas of the Norrbotten Porphyry Arc closely resemble the modern Andinotype high-K calcalkaline lavas in composition, indicating increasing maturity of the arc in
Volcanic rocks of Early Proterozoic greenstone belts
21
the later stages of the Svecofennian Orogeny. The cross-cutting trend of the Norrbotten belt reflects the re-orientation of the subduction vector to become east directed, thus setting the scene for the magmatism of the Transscandinavian Granite-Porphyry Belt (Pharaoh and Brewer, 1990). Lewisian Coniplex, northwest Scotland
The Loch Maree Group amphibolites (Sm/Nd model agc 2.0 Ga, Johnson et al., 1987) are part of the Lewisian Complex of northwest Scotland which has been correlated with the Nagssugtoqidian mobile belt of Greenland (Fig. 7). They
Fig. 7. Geographic s~ibdivisionsand gencralizcd lithotectonic elements of Early Proterozoic terranes in North America (from 1,ervry e t al., 1987). 1 = contincntal margin prism, 2 = Penibine-Wausau arc terrancs, Pciiokean Orogen.
22
TE. Smith
outcrop in a narrow belt of supracrustal rocks and are associated with narrow bands of siliceous schist, banded iron formation, graphitic schist and marble. The amphibolites are metamorphosed tholeiitic mafic volcanics which can be divided into two petrogenetic groups (Johnson et al., 1987, fig. 10). Group A, are chemically similar to MORB and are characterized by slight LREE enrichment and have no Nb anomalies in their multi-element signatures (Table 2). They are interpreted as primitive basalts which erupted rapidly, directly from the mantle, after a period of extensional rifting had caused significant crustal thinning. Group B are interpreted as mafic sills, and show LREE enrichment and negative Nb anomalies (Table 2) and are chemically similar to the (continental tholeiitic) Scourie dykes (Table 1). The Group B sills were emplaced after the extension ceased and the rift had been filled with sediment and their chemical compositions are believed to have been affected by crustal contamination (Johnson et al., 1987). The Nagssugtoqidian mobile belt of Greenland has been equated with the Trans-Hudson Orogen (Fig. 7) thus the Loch Maree amphibolites provide a link between the North Atlantic Craton and the Laurentian Shield. The Laurentirrn Shield
At least seven microcontinents (Archean cratons) aggregated at about 1.8 Ga to form the Laurentian Shield. They are separated by Early Proterozoic fold belts including the Penokean, Trans-Hudson, Wopmay, and Ketilidian orogens (Fig. 7) (Hoffman, 1989). The Early Proterozoic Penokean Orogen formed along the southern margin of the Archean Superior Craton. It is divided into a deformed continental margin prism, separated from the Wisconsin magmatic intraoceanic arc (Pembine-Wausau and Marshfield) terranes to the south by the Niagara suture zone (Fig. 7). The Huronian Supergroup can be used to illustrate the characteristics of the northern continental margin prism. It comprises a clastic wedge deposited along the northern edge of a rift zone formed at the southern margin of the Superior Craton (Jolly, 1987a). Volcanic rocks, the Elliot Lake Group at Thessalon (ca. 2450 Ma), occur only in the first of four transgressive sedimentary sequences. At the base it comprises 500 m of pillowed basalts and basaltic andesites, interlayered with quartz conglomerates, overlain by 200 m of rhyolites, and followed by 700 m of massive basalts, basaltic andesites, and andesites. The fourth subdivision recognised consists of a fault block of basalts interlayered with rhyolites. The basalts in this unit are similar geochemically to those of the third unit. The youngest volcanic unit is comprised of basalts and one hybrid andesite, unconformably overlain by a quartzite conglomerate (Jolly, 1987b). The basalts (CT) are all LILE and LREE enriched and show negative Nb anomalies in their multi-element patterns (Jolly, 1987a, fig. 3; Table 3). There are two distinct suites of rhyolites, both of which are interpreted as crustally derived. The sequence is compared to the Etendeka volcanics of the Karoo, which formed during rifting leading to the opcning of an ocean basin. Jolly (1987a) suggests that the basalts
Volcanicrocks of Early Proterozoic greenstone belts
23
TABLE 3 ljpical MORB-normalized elemental ratios and contents of selected Early Proterozoic basalts and andesites from the Laurentian Shield and southwestern U.S.A.
Elliot Lake a Quinneseca North Wausau Central Wausau a South Wausau South Wausau Watts 1a Watts za Eskimo a Povungnituka Basanite 3 Phonolite 3 Flahertya Chukotat 4 a Chukotat 5 a Chukotat 5 Chukotat 6 d Chukotat 6 Assean Lake a Fox River a Moak Lake a Opswagan Lakea La Ronge a La Ronge Lynn Lake a Lynn Lake Akaitcho a Green Mountain a Duboisa Cochetopa a Salida a Black Canyon a Pews Peas
Pewsa Peas Dos Cabezasa Gold Hill a Gold Hill a
4.50 ~0.60 x2.50 x2.50 x0.70 ~4.00 1.34 0.50 2.99 1.64 2.80 10.78 1.60 1.08 2.20 1.10 1.04 1.68 1.20 1.20 1.06 1.36 2.23 3.30 2.38 2.92 3.41 1.02 2.09 2.36 2.49 1.49 1.00 1.61 3.07 1.82 2.60 2.08
~0.60 ~6.50 ~5.00 x0.50 ~14.00 2.05 0.32 5.52 4.74 13.09 30.87 2.82 0.83 3.23 1.71 1.09 1.97 3.10 1.07 0.97 1.31 5.54 8.69 2.70 3.39 Z5.5 8.69 0.88 4.61 2.88 2.67 2.41 0.49 1.46 4.20 3.06 5.07 3.47
2.25
8
28
5
21
1.91 1.57 0.51
<1 10 20
34 28
0.60
14 16
36 28
0.70
0.34 0.64 0.95 2.92 0.64
7 2 3 3 4
0.40 0.75 1.08 3.42 0.74
0.61 1.47 1.52 1.41 1.77
4 7 7 5 7
18 13 15 28 17 28 17 17 20 21
33
0.71 2.39 1.78 1.61 2.07
4 3 8 10 14 <2 2.5 4 5 8
19 28 23 22 36 5 13 29 18 44
3.78 1.02 1.94 1.22 2.69 >1.00 0.78 1.15 2.24 1.81
3.23 0.87 1.66 1.07 2.30 >0.85 0.67 0.98 2.08 1.55
2.63
2.23 1.84 0.60
basalt; andesite; komatiite; Mg-basalts; spinifex basalts. @
1 = flat LREE; 2 = depleted LREE; 3 = Povungnituk Group; 4 = Northern Cape Smith; 5 = Southern Cape Smith; 6 = Hudson Bay islands.
24
TE. Sinith
were derived from a mantle which became enriched in LILE and LREE as a result of ancient subduction of crustal material, probably during the formation and stabilization of the Archean continental crust. The volcanic rocks of the 1860-1889 Ma Pembine-Wausau terrane comprise dominantly metamorphosed basalt, andesite, dacite and rhyolite flows and pyroclastic rocks. The Quinnesec Formation occurs in the northern part of the terrane and comprises tholeiitic basalts and basaltic andesites (and gabbro sills) characterized by limited iron enrichment, low Ti02 and HFSE contents, and LREE depletion (Table 3). These mafic rocks are associated with plagiorhyolite with low K 2 0 and R E E abundances and flat R E E patterns (Sims et al., 1989, fig. 3A) and sheeted dykcs and scrpcntinites. The chemical composition of the Quinncsec basalts ranges from that of the back-arc basalts of the Lau basin to that of the basalts of the Troodos ophiolite, and the unit is regarded as a dismembered back-arc ophiolite. The basalts are overlain by a calcalkaline andesite to rhyolite sequence with moderately enriched LREE and flat H R E E patterns (Sims e t al., 19S9, fig. 3B; B b l e 3). A bimodal suite, comprising high-Al pillow basalt and basaltic andcsitc flows interlayered with dacite to rhyolite tuffs and porphyries, occurs in the central part of the Pembine-Wausau terrane at Monico. All of the rocks show LREE enrichment, flat H R E E patterns, and Ti depletion, and are similar to the bimodal calcalkaline volcanics of the Kuroko area of Japan, and to calcalkalinc suites formed in oceanic arcs (Sims et al., 1989). 1835-1845 Ma calcalkaline andcsite, dacite and rhyolite, characterized by greater LILE and LREE enrichment than the older calcalkaline rocks (Sims et a]., 1989, fig. 6), occur in the southern part of the tcrrane. They are associated with older LREE dcplctcd tholeiitic basalts (Table 3). The older volcanic sequences are considered to have formed above a south-dipping subduction zone and the younger sequences above a north-dipping subduction zone. The Marshfield terrane, which occurs to the south of the Eau Pleine suture, consists of an apparent remnant of an 1S60 Ma mafic to felsic volcanic succession. It comprises interlayered, metamorphosed mafic volcanics, dacite porphyries, impure quartzite, chert, conglomerate and calcareous argillite which overlie Archean gneisses. No chemical data is available for the volcanic rocks. The PembineWausau and Marshfield magmatic tcrrancs were accreted to the continental margin in the interval 1840-1860 Ma, and the Quinnesec sequence is interpreted as a dismembered ophiolite caught up in the Eau Plcine suture zone. The Circum-Superior belt (Fig. S) extends north through Labrador to the Ungava Region of Quebec (Cape Smith belt) and the islands of eastern Hudson Bay to northeastern Manitoba (Fox River and Thompson Nickel belts). It compriscs Early Proterozoic supracrustals of rifted to passive margin character and separates the Trans-Hudson Orogen (Fig. 7) from the Archean Supcrior Craton (Lcwry ct al., 1987, fig. 1). The Cape Smith belt is divided into northern and southern domains, only rocks of the southern domain are found in the Hudson Bay islands and other localities. The northern domain comprises a vast dismembercd ophiolite complex (the Watts Group), and a volcano-sedimentary sequence
SEGMENTS OF THE CIRCUM-SUPERIOR
BELT
CHURCHILL PROVINCE GNEISS-UNDIVIDED
0SUPERIOR PROVINCE GNEISS-UNDIVIDED
m
2
a
za2 2
Fig. 8. Generalized map of the distribution of the Circum-Superior Belt (from Baragar and Scoates, 1987).
26
T.E. Smith
of shales, siltstones and greywackes intercalated with a basalt to rhyolite calcalkaline volcanic sequence (the Parent Group). The metavolcanics of the southern domain are, in part, interlayered with sediments of continental origin (quartzites, dolomites, iron formation, and siliceous siltstone and shales) and unconformably overlie the Archean Superior Craton in places. Tholeiitic basalts and minor felsic rocks (Povungnituk Group in Cape Smith, Eskimo and Flaherty Formations in the Hudson Bay islands) are overlain by komatiitic rocks and their derivatives (Chukotat Group) (Arndt et al., 1987; Baragar and Scoates, 1987; Picard et al., 1990; Halden, 1991). Recent U-Pb zircon dates show the belt ranges in age from approximately 2000 to 1880 Ma (Watts Group 1999 f 3 Ma; Picard et al., 1990, tholeiitic basalts of the Povungnituk Group 1960 f 5 Ma; Baragar and Scoates, 1987, alkaline volcanics of the Povungnituk Group 1960 f 3 Ma; Picard et al., 1990, basalts of the Chukotat Group 1922 f 8 Ma; Picard et al., 1990). Details of the chemical compositions of the various volcanic units are given in Arndt et a]., 1987, tables 2A, 2B and 3, figs. 5, 6 and 7; Baragar and Scoates, 1987, tables 1 A and lB, figs. 4, 5, 6, and 8 and Picard et al., 1990, tables 1, 2, 3, and 4, figs. 8 and 9). In the northern domain the Watts Group comprises a sequence of chromite-rich dunites, peridotites, pyroxenites, layered gabbros and anorthosites associated with sheeted diabase dykes and massive and pillowed lavas. Two suites of lavas and diabases similar in major element composition to oceanic tholeiites occur. The first suite has generally flat LREE patterns and 1-6 ppm Nb (Table 3), and is similar to N-MORB. The second suite has lower concentrations of Nb (<1ppm) and highly depleted LREE patterns (Table 3). The Zr/Nb (95-136), Ba/Zr (0.45-0.61), La/ Nb (1.8-2.66) and La/Ta (2.62-2.95) ratios of the second suite suggest a back-arc affinity (Picard et al., 1990). The Watts Group is interpreted either as an oceanic (Scott et al., 1988) or back-arc (Picard et al., 1990) ophiolite complex (Purtuniq Ophiolite). The Parent Group comprises approximately 40% basalts, 50% largely pyroclastic andesites, and 10% rhyolitic tuffs. They have slightly LREE enriched patterns and are identified as a calcalkaline suite which formed in a magmatic arc (Picard et al., 1990). The oldest metavolcanics of the southern domain, in the Eskimo Formation, are exposed on the islands of Hudson Bay (Fig. 8) and comprise subaerial, plagioclasephyric, tholeiitic basalts and andesites which are interbedded with dolomites and shales with minor iron formation. The basalts have relatively high concentrations of LREE, negative Nb anomalies and moderately to strongly fractionated multielement patterns ( a b l e 3) and are similar in composition to CT (Arndt et al., 1987). The overlying Flaherty Formation comprises a series of massive and pillowed flows and sills of subalkaline to transitional (Nb/Y 0.35-OSO), plagioclase-phyric basalts (Arndt et al., 1987). They have moderate LREE enrichment and positive Nb anomalies (Table 3). The Povungnituk Group, of the Cape Smith belt, has been correlated with the Flaherty Formation (Baragar and Scoates, 1987), and comprises the Lamarche sub-group of sediments which are overlain by the
Volcanic rocks of Early Proterozoic penstone belts
21
Beauparlant sub-group of uniform plagioclase-phyric, LREE-enriched tholeiitic basalts with positive Nb anomalies (Table 3). They have unfractionated HREE patterns, and relatively high concentrations of Zr, Ti, and Y and high ZriY ratios (3.2 to 9.35) and are interpreted as C T or initial rift tholeiites (Holm, 1985, B b l e l ) , erupted in an extensional environment (Picard e t al., 1990). The tholeiites a r e overlain by small volumes of within-plate alkaline basanites and phonolites having highly fractionated REE patterns (Table 3). The Chukotat Group overlies the Povungnituk Group and comprises olivinephyric komatiitic basalts, pyroxene-phyric and plagioclase-phyric tholeiitic basalts in the Cape Smith belt (Baragar and Scoates, 1987; Picard et al., 1990). Mgbasalts, interleaved with spinifex basalts at the highest levels, occur in the Ottawa and Belcher Islands of Hudson Bay (Arndt e t al., 1987). In the north of the Cape Smith belt the Chukotat lavas have flat to slightly L R E E depleted patterns, many of which are similar to modern MORB (Tdble 3). The southern sequence is more LREE enriched and is considered to be transitional between ocean ridge lavas and continental tholeiites (Table 3). The Mg-basalts are chemically similar to the rocks of the Chukotat Group and have flat to slightly enriched L R E E patterns, positive to negative Nb anomalies and low concentrations of incompatible elements (Table 3). The spinifex basalts of the Ottawa Islands have slightly higher concentrations of LREE, and the same range of concentrations of HREE as compared to the Mg-basalts, together with negative Nb anomalies and slightly fractionated multi-element patterns (Arndt et al., 1987, Tdbk 3). Picard et al. (1990) suggest that early oceanic crust (Purtuniq Ophiolite) was formed in the northern domain (1999 Ma) followed by the opening of the Povungnituk-Chukotat basin to the south by asymmetric rifting. Continentally derived sediments (Lamarche sub-group) accumulated in the developing rift, followed the eruption of the LREE-enriched tholeiites and the alkaline volcanics of the Beauparlant sub-group (1960 Ma). Further extension of the rift margin may have induced partial melting of the asthenosphere and led to the eruption of the MORB-like and L R E E enriched members of the Chukotat Group. T h e Parent Group may have formed in response to north-south compression which caused the subduction of the Chukotat volcanics in the north (Picard e t al., 1990). Arndt et al. (19S7) suggest a very similar reconstruction for the southern domain based on the volcanic sequences occurring in the Hudson Bay islands. Early contaminated continental basalts (Eskimo Formation, L R E E enriched, negative Nb and B anomalies, EN^ -7 to -8) were formed during rifting. They were succeeded by uncontaminated transitional basalts (Flaherty Formation, LREE, Nb and ?a enriched, ENd +1 to +3) and a subaqueous series of MORB-like basalts (Mg-basalts, L R E E depleted, ENd +4 to +5) erupted during the spreading phase. The occurrence of crustally contaminated komatiites (spinifex basalts, L R E E enriched, negative Nb and Ta anomalies, t~~ z 0) at the top of the sequence, suggests closure of the ocean basin. The mafic and ultramafic rocks of the Churchill-Superior boundary zone show similar variations to those described above (Halden, 1991, table 1, figs. 10, 11, 12).
TE. Smith
28
Komatiitic volcanics in the Fox River Group (ENd +5, Chauvel et al., 1987) and the Molson Dykes in Manitoba have been correlated with the Chukotat Group (1922 Ma) although they yield younger U-Pb zircon ages (Molson dykes 1883.7:;;: and 1883 f 2 Ma; Fox River Sill lSS2.9::.,5, Ma Heaman e t al., 1986). At Assean Lake the mafic rocks, arc enriched in LREE and depleted in Nb relative to MORB and are similar to the basalts of the Eskimo Formation (Table 3). The mafic rocks of both Fox River and Ospwagan Lake have high MgO contents(~14%) and relatively flat multi-element patterns which are slightly enriched in Nb like those of the Mg-basalts described by Arndt et al. (1987) (Table 3). The mafic rocks exposed at Moak Lake with much lower MgO contents ( ~ 7 % have ) similar multi-element patterns to the Fox River and Ospwagan Lake rocks at slightly highcr concentrations (Tmble 3). Halden (1991) suggests that these rocks were erupted in marginal basin formed along the northern edge of the Superior Craton, which may be a continuation of the Circum-Superior belt. The variation in their trace element chemistry rcquires local compositional variability in their mantle source regions (Halden, 1991). Gaskarth and Parslow (1987) report the close association of volcanic rocks, t N d +4 to +5 (Chauvel et al., 1987) in the Flin Flon belt (Fig. 9) having within-plate and volcanic arc signatures, and suggest that it represents a small marginal basin. The sequences are bimodal, comprising metabasalts and rhyolites, and there may also be some komatiites. The majority of the MORB-normalized patterns of these rocks show negative Nb anomalies (Gaskarth and Parslow, 1987,
\*:
I-
Fig. 9. La Ronge, Lynn Lake a n d Flin Flon domains, Trans-Hudson O r o g e n (from Lewry et al., 1987).
Volcanic rocks of Early Proterozoic peenstone belts
29
fig. l l ) , and there is also a considerable range of Ti02 content in the basalts. The associated metasediments include meta-arkoses, quartzites, and metagreywackes but stratigraphic relationships are uncertain. These are all features of rock suites erupted in marginal basins (Saunders and Thrney, 1984). This basin lies very close to, if not at, the margin of the Superior Craton (Fig. 8) and the range of volcanic rock types and compositions described above is very similar to those of the Circum-Superior belt. It is therefore suggested that the Flin Flon greenstone belt may be a marginal basin representative of the westerly extension of the Circum-Superior belt. The Trans-Hudson Orogen (Hoffman, 1981) is an Early Proterozoic orogenic belt formed by tectonic processes similar to those operating at a modern collisional margin. It is subdivided (Fig. 7) into the Cree Lake Zone (miogeosynclinal, ensialic) to the northwest and the Reindeer Lake Zone (eugeosynclinal, ensimatic) to the southeast, separated by the Wathaman Batholith (Gaskarth and Parslow, 1987). The La Ronge and Lynn Lake greenstone belts occur in the La Ronge Domain of the Reindeer Lake Zone (Fig. 9). Metarhyolites from the La Ronge domain yield U-Pb zircon ages of 1879 f 9 and 1880 f 7 Ma (Lewry et al., 1987) and felsic lavas from the Wasekwan Group in the Lynn Lake belt have a U-Pb age of 1910 (Baldwin et al., 1987). They are separated from the Flin Flon greenstone belt to the south by the Kissynew Domain (Fig. 9), comprising high grade paragneisses derived from grcywacke, and subordinate volcanic rocks and arkoses, which is at least in part contemporaneous with the greenstone belts. The volcanic sequences in the La Ronge and Lynn Lake belts comprise dominantly basalt and andesite, with local accumulations of dacite and rhyolite, which were erupted in a volcanic arc setting (Wattcrs and Pearce, 1987; Peck and Smith, 1989). d ranging from +4.0 to Igneous rocks in the Western La Ronge belt yield t ~ values f4.9 (Chauvel et a]., 19S7). The basalts and andesites of the La Ronge belt all have similar multi-element signatures (Watters and Pearce, 1987 fig. 5, table l), which lie in the range of calcalkaline suites (Tmble 3). In areas of felsic rock accumulation the basalts and andesites have higher elemental concentrations and more fractionated patterns i.e. are more calcalkaline in nature (Watters and Pearce, 1987, fig. 9). The andesites are similar in chemistry to orogenic andesites formed in continental margin arcs (CIA suite, Bailey, 1981). The La Rongc arc was previously interpreted as of entirely ensimatic character, but the incompatible trace element contents of the volcanic rocks suggest that enriched subcontinental lithosphere was involved in their genesis (Watters and Pearce, 1987). However, the ENd values have been interpreted as indicating that the volcanics are contaminated by very small (<2%) quantities of Archean crust. In the Lynn Lake belt the supracrustal sequence comprises tholeiitic basalts at the base, overlain by a bimodal sequence of basalts and felsic volcanics and an uppermost series of calcalkaline andesites (Syme, 1985; Peck and Smith, 1989). The basalts and andesites are LREE enriched with negative Nb anomalies (Xible 3), indicating that they are petrogenetically related and were erupted in a
30
T.E. Smith
volcanic arc setting (Peck and Smith, 1989, tables 2 and 3, figs. 10, 11, and 12). The chemistry of the basalts lies in the range of the low-K to calcalkaline suites, and the composition of the andesites ranges from that of orogenic andesites erupted partly on thin continental crust to those erupted on well developed continental crust or at thin continental margins (OA to CIA andesites, Bailey, 1981). Overall the geochemistry of these volcanic rocks (Peck and Smith, 1989, tables 7 and 8, figs. 11, 14, 15, 17; Table 3) suggests that they developed on relatively thin continental crust. The Wopmay Orogen is preserved in the NW Canadian Shield (Fig. 7), and contains a record of a complete Wilson Cycle, strongly suggesting that Phaneroicstyle plate tectonics operated in the Early Proterozoic. The following summary is largely taken from Hildebrand et al. (1987). The Hottah magmatic arc was established on the western margin of the Slave Craton by 1914 Ma. A marginal sea, in which sediments accumulated, formed behind the arc (1900-1890 Ma). The Akaitcho Group was deposited in discrete rift basins formed during this crustal stretching and comprises mature clastic and chemical sediments intercalated with CT at the base. These rocks are overlain by wedges of relatively deep-water feldspathic grit and coalescent submarine shield volcanoes, up to 4 km thick, made up of tholeiitic basalts and subordinate subalkaline rhyolites (felsic/mafic ratios 0.2-0.3) (Easton, 1983; Hoffman and Bowring, 1984). The basalts range from strongly LREE enriched C T at the base of the Akaitcho Group to less LREE enriched tholeiites, interpreted as oceanic, higher in the sequence (Table 3). Similar upward changes occur in the basalts of the shield volcanoes (Easton, 1983). The formation of these rocks was followed by shortening of the basin and the contemporaneous emplacement of the Hepburn intrusive suite (1986-1878 Ma), and the inception (1878 Ma) and growth of the Great Bear magmatic zone on the closed basin. Oblique folding (1860-1850 Ma), and the emplacement of the syn- and post-folding syenogranites (1858-1843 Ma), affected the Great Bear zone. Regional transcurrent faulting affected the area at some time between 18431810 Ma. The Great Bear magmatic zone shows all of the lithological, structural and volcanic features typical of a continental-margin arc. The volcanic suite is dominated by andesites, but ranges in composition from basalt to rhyolite. The major element chemistry of these rocks, which is similar to that of the volcanic rocks of the Andes and the western North American Cenozoic arcs, indicates that they are best classified as a high-K calcalkaline suite. This identification is confirmed by the incompatible trace element patterns of the basalts and andesites which are LREE enriched and have negative Nb anomalies. The majority of the volcanic rocks have high overall R E E abundances, but some of the mafic andesites have less evolved, flat, R E E patterns with abundances ~ 1 0 - 2 0times chondrite (Hildebrand et al., 1987). The dacites and rhyolites cannot be derived from the andesites by fractional crystallization. The pre-folding intrusions are calcalkaline I-type, showing smooth trends on variation diagrams. The post-folding group is dominated by I-type syenogranites with lesser amounts of intermediate, tholeiitic and calcalkaline rocks.
VoIcanic rocks of Early Proterozoic greenstone belts
31
The 1.9-1.7 Ga Ketilidian terrane formed to the south of the North Atlantic Craton (Patchett and Bridgwater, 1984). It is characterized by abundant calcalkaline volcanic and plutonic magmatism, and is considered to represent another Early Proterozoic continental margin arc. Nd, Pb and Sr isotope studies show that most of the Keltidian mobile belt is not underlain by Archean crust, and new mantle-derived crust to which Archean sedimentary that it consists of ~ 9 0 % material was added (Patchett and Bridgwater, 1984; Kalsbeek and Taylor, 1985). Central and southwestern North America
Five Early Proterozoic volcanic regimes (Condie, 1986), having different mantle extraction ages (1760-1800,1730-1740, 1720, 1680-1700, 1650 Ma), which overlap spatially, have been described in the southwestern U.S.A. (Fig. 10). A quartzitepelite sequence (1650-1750 Ma) overlies or interfingers with the five volcanic terranes. In sequences assigned to the 1760-1800, 1730-1740, and 1720 Ma terranes the mafic volcanics exceed the felsics and non-volcanogenic sediments are minor or absent. In the 1680-1700 and 1650 Ma sequences felsic volcanics exceed the mafics and non-volcanogenic sediments are abundant (Condie, 1982, table 1). The lithostratigraphic relationships, immobile incompatible element contents of the basalts, andesites and felsic volcanics (Condie, 1987, figs. 3, 4, and 5; Knoper and Condie, 1988, figs. 2, 3, 4, and 5; Robertson and Condie, 1989, figs. 8, 9, 10, 11, 12, and 13), and geochronology show that the continental crust of this area grew largely by accretion of island arc and continental-margin arc related rocks to the Archean Wyoming Craton (Condie, 1982, 1986; Bowring and Karlstrom, 1990). The volcanic sequences are assigned to the Yavapai and Mazatzal Early Proterozoic crustal provinces, which together with the Mojave crustal province, can be differentiated by their geological histories, their Pb isotope characteristics (Anderson et al., 1991), and the time of their accretion to the continent (Bowring and Karlstom, 1990; Condie, this volume). The majority of the volcanics are bimodal mafic-felsic volcanic suites. Orogenic andesites are comparatively rare but have been described from the Yavapai Supergroup of the Ash Hill terrane in the Yavapai crustal province, and in the Pedernal Hills succession in the Pecos Terrane of the Mazatzal crustal province (Condie, 1986, and this volume, fig. 1). The relative paucity of andesitic rocks may be the result of preferential erosion of subaerial portions of the arcs (Condie, 19%). The Mojave Province contains rocks ranging in age from 2300 to 1700 Ma. No geochemical data are available for the volcanic rocks and the province is regarded as cratonic or may represent a continental margin arc (Condie, this volume, table 1). It is not considered further here. Four terranes and three overlap assemblages occur in the Yavapai Province (Condie, this volume, table I). The Green Mountain terrane (1790-1780 Ma) with rocks having LREE enriched, Nb depleted, patterns, the Dubois terrane (17801750 Ma) in which rocks have flat REE patterns with no N b anomaly, and Ash Creek (mainly 1750-1740 Ma) terrane, are interpreted as representing island arcs.
32
TE. Smith
Fig. 10. Distribution of Early Proterozoic volcanic terranes in the soutliwester~~ United States (from Condie, 1987).
The Hualapai terrane (1740-1700 Ma) and Cochetopa-Salida overlap assemblage (1740-1730 Ma) with rocks which are LREE enriched with a small negative Nb anomalies, represent continental-margin arcs. The Idaho Springs-Black Canyon (1740-1720 Ma) in which rocks which arc LREE enriched with negative Nb anomalies occur, and Wet Mountain (1720-1670 Ma) overlap assemblages, and are interpreted as continental-margin back-arc basins (Table 3). The coeval Cochetopa and Black Canyon sequences comprise largely tholeiitic basalts, but include a subordinate suite of high-Mg rocks classified as komatiites and komatiitic
Volcanic rocks of Early Proterozoic greenstone belts
33
basalts (Blackburn and Vance, 1989). None of these terranes show trace element or Nd isotope evidence of forming on or near significantly older continental crust (Condie, this volume). The Mazatzal Province comprises three identifiable terranes and one overlap assemblage. The Pecos Terrane (1720 Ma) is interpreted as an oceanic arc in which MORB-like and high-Mg tholeiitic basalts volcanic rocks, with flat to slightly LREE enriched and no Nb anomaly (Dble 3), were derived from a depleted MORB-like mantle. They are associated with LREE enriched, Nb depleted, calcalkaline basalts. The field occurrence of ultramafic subvolcanic rocks reported from this terrane, which are slightly LREE enriched with no Nb anomalies, is reminiscent of Phanerozoic ophiolites (Robertson and Condie, 1989). They are interlayered with fine-grained amphibolite (metabasalt) and are cut by numerous closely spaced metagabbro and metadiabase and minor trondhjemitic dykes. The Dos Cabezas-Pinal Terrane (1700-1670 Ma), which contains a bimodal volcanic suite in which the basalts are LREE enriched and Nb depleted ( n b l e 3), is interpreted as a continental-margin arc, while the Alder Terrane (mainly ~ 1 7 0 Ma) 0 and Manzano overlap assemblage (1700-1650 are both interpreted as continental-margin back arc basins (Condie, this volume, table 1). The volcanic rocks of the Gold Hill-Wheeler Peak area comprise basal& and andesites which are LREE enriched in comparison to N-MORB and may be comparable with calcalkaline rocks erupted in a continental-margin arc (Condie and McCrink, 1982) Ga Payson ophiolite (Dann, 1991) comprises In the Alder terrane the ~ 1 . 7 3 gabbro-diorite intrusions, dyke swarms, sheeted dykes and submarine basalts, which are overlain by dacitic breccias and a thick sequence of turbidites. The sheeted dykes comprise mainly tholeiitic basalt and subordinate andesite with island-arc affinities. The dykes show Ti02 enrichment versus MgO, have low (<1.0%) Ti02 contents and selective enrichment of LILE over HFSE relative to MORB. They plot in the island-arc tholeiite field on Ti-V plot (Dann, 1991, figs. 3A, 3B, 3C). It is suggested that the ophiolite formed in situ as the floor of an intra-arc basin (Dann, 1991). The complex chronological and spatial relationships existing between the volcanic terranes show that most of southwestern North America was assembled between 1.8 and 1.65 Ga. The growth of the crust by accretion is marked by four principal Early Proterozoic collision events (Condie, this volume, table 2). The Grcen Mountain-Wyoming Craton (Archean) collision (1780-1765 Ma) occurred when the Green Mountain and Dubois terranes accreted to the Wyoming Craton during the Cheyenne Orogeny. The collision of the Mojave and Yavapai crustal provinces took place about 1700 Ma in the Ivanpah Orogney; the Pecos and Dubois terranes also collided at about 1700 Ma during the Pecos Orogeny. The collision of the Oklahoma and Mazatzal provinces occurred at about 1650 Ma during the Mazatzal Orogeny.
TE. Smith
34
The South American Shield In South America, Early Proterozoic greenstone belts are poorly known but show many of the features already described from the Baltic Shield and North America. In the northern Guyana Shield, bimodal, tholeiitic and calcalkaline volcanic sequences, with few rocks in the 63-68% Si02 range, are typical. Basalts dominate the lower sections and are overlain by a more varied sequence of porphyritic mafic, intermediate, and felsic volcanics (Thble 4). High-Mg basalts TABLE 4 Typical MORB-normalized elemental ratios and contents of selected Early Proterozoic basalts and andesites from South America, India, China and Australia
Mazaruni a Mazaruni 1a Mazuruni Piumhi PiumIii PiumIii a Piumlii
__ 1.29 2.50 3.21
1.51 2.05 3.26 3.16
1.69 3.51 8.12 1.68 2.31 9.02 8.34
Birimian 2 a Birimian 3a Birimian 4 a Birimian 5 a Birimian 6 a Birimian 7 a Birimian 8 a Birimian 8 Orange River Grp a Orange River Grp
1.18 1.51 0.93 1.91 0.66 1.06 1.51 2.74 2.49 4.63
1.16 1.76 0.66 6.47 0.39 1.11 1.55 5.14 8.62 13.45
Aravalli a Dahaa
2.08 0.71
5.54 0.68
Tongkuangyu Fm a Songjiashan G p a Zhongtiao G p a Xiyanghe Gp a Xionger
2.69 1.32 2.30 2.44 4.30
1.85
Willyama a Willyama Oonagalabi a L. Marts Range a U. Harts Range a
3.09
27 18 30
0.50 0.59 0.12
5 5 6
23 19 21
0.59 0.69 0.14
2.08 3.32
7 9
18 22
2.43 3.89
6.42 1.89 3.74 9.46 10.75
2.56 1.82 2.05 3.50 5.05
5 2 10 10 8
20 22 49 39 35
3.00 2.13 2.40 4.10 5.91
2.55 5.63 4.59 1.07 4.31
0.99 1.60 2.33 2.19 2.97
4 16 9 2 4
22 49 57 45 34
1.15 1.88 2.72 2.5 3.47
31
Basalt; andesite; peridotitic komatiite; basaltic komatiite. 1 = high NzO3; 2 = Tsalabya Formation, Mauretania; 3 = Mako Formation, Senegal; 4 = YaourC Formation, Ivory Coast; 5 = Haute-Comoe Formation, Ivory Coast; 6 = Bouroum I1 Formation, Burkina Faso; 7 = Bouroum I Formation, Burkina Faso; 8 = Liptako Formation, Niger. a
Volcanic rocks of Early Proterozoic greenstone belts
2120
- 1252
35
2310
0.1, " " " " " " ' Rb Ba Th K Nb La Ce Sr P Zr Srn Ti Y Yb
Fig. 11. MORB-normalized multi-element plots showing typical compositions of the tholeiites (1252), high-AI203 basalts (2120), and pyroxene andesites (2310) of the Issineru Formation, Mazuruni greenstone belt, Guyana Shield.
and ultramafic schists, possibly of komatiitic affinities, are also present (Gibbs, 1987). Komatiites, with enriched L R E E and flat H R E E distribution patterns associated with tholeiites, and minor andesites (Table 4) (Sm-Nd age 1.84 Ga), occur in the Piumhi Massif of southeastern Brazil (Jahn and Schrank, 1983). The base of the dominantly volcanic Issineru Formation, which makes up the lower 5-7 km of the Mazaruni greenstone belt in the Guyana Shield, is obscured by intrusive contacts (Renner and Gibbs, 1987). The formation is overlain by the Haimaraka Formation which is composed of hematitic quartzites, tuffaceous conglomerates, and greywackes. The lower part ( ~ km) 3 of the section is made up of massive and vesicular pillow basalts, and gabbros. Most of these mafic rocks are tholeiites with relatively flat LREE patterns without Nb anomalies (Fig. 11, Table 4) and are similar to N-MORB in composition. They are overlain by a sequence dominated by intermediate (basaltic-andesites and andesites) with felsic (tuffs and rhyolites) rocks, and minor amounts A1203-rich plagioclase-phyric basalts. These basalts have LREE enriched patterns and moderately fractionated multi-element signatures (Fig. 11, n b l e 4), like those of CAB or CT, but are not consistently Th enriched (Renner and Gibbs, 1987, fig. 5). The majority of the andesites which occur in the upper part of the succession have slightly (Fe-rich tholeiitic) to moderately enriched (hornblende- and pyroxene-phyric, calcalkaline) LREE patterns, most of which have negative Nb anomalies (Renner and Gibbs, 1987, fig. 6, table 2; Fig. 11, Table 4) and are similar in composition to CIA of continental margin arcs (Table 1). Tholeiitic felsic rocks also occur and are characterized by moderately enriched L R E E patterns, with negative Eu anomalies with high overall abundances, some of which have negative Nb anomalies. There are also calcalkaline felsics that have highly enriched LREE patterns with lower overall abundances. The majority of these felsics have negative Nb anomalies in their multi-element signatures (Renner and Gibbs, 1987, fig. 5). These data show that the lower part of the Issineru Formation is characterized by N-MORB basalts, and there is a predominance of intermediate and felsic rocks above. Definitive chemical signatures are absent but Renner and Gibbs (1987) suggest that this association may have been formed in rifted continental crust opening into a marginal basin.
36
TE. Smith
The African Shield Immense stretches of juvenile continental crust, known as the Birimian terranes (Abouchami e t al., 1990, fig. 1) were formed in a very short period of time, around 2.1 Ga, in West Africa (Mauritania, Algeria, Mali, Senegal, Guinea, Ivory Coast, Burkina Faso, Niger, and Ghana) (Vachette, 1964; Ledent et al., 1969; Bassot and Caen-Vachette, 1984). Two major Birimian units have been identified, the first composed dominantly of volcaniclastic rocks of felsic to intermediate affinity, and the second is comprised mostly of a bimodal, volcanic sequence (Abouchami et al., 1990). Tholeiitic mafic rocks are dominant in the bimodal sequence and occur with calcalkaline andesites, dacites and rhyolites. Some of the tholeiites may have formed in a mid-ocean ridge environment (N’Gom, 1985; Fabre, 1987; Zonou, 1987), but others, which are associated with subordinate andesites, and are depleted in Ti and Nb, may be representative of an immature island arc or back-arc basin environment (Zonou, 1987; Dia, 1988). Most Birimian lavas have relatively high A I 2 0 3 , Rb, Ba, and K, and low TiO2, R E E and HFSE contents as compared to MORB (Abouchami et al., 1990, tables l a and lb, fig. 3 ) . The majority of the lavas have flat R E E patterns although some are LREE depleted and some are LREE enriched (Abouchami et al., 1990, ) from +1.2 to +4.9 which, together with their fig. 4, Table 4). E N ~ ( Tvaries Sm/Nd ratios, suggests that they were derived from at least two depleted mantle sources both of which were different from the source of present day MORB (Abouchami et al., 1990). Chemically the lavas are similar to back-arc or low-Ti continental flood basalts (Bble 4) but their geological setting is inappropriate for either environment. The Nb concentrations are generally higher than those of arc magmas and the trace element signatures lack negative Nb anomalies ( B b le 4). Abouchami et al. (1990) suggest that the common occurrence of pillow lavas, and the absence of rocks of continental crustal affinities, associated with the bimodal volcanics is strongly suggestive of an oceanic environment of eruption. The Ce/Nb ratios are lower than those of IAB and the basalts plot in the fields of MORB and oceanic flood basalts of the Nauru Plateau on the Ce/Nb vs. Ce diagram (Abouchami et al., 1990, fig. 12.). It is concluded that some of the Birimian mafic rocks may have erupted as oceanic flood basalts (Abouchami et al., 1990) but others may have formed in mid-ocean ridge and immature arc or back-arc basin environments (N’Gom, 1985; Fabre, 1987; Zonou, 1987; Dia, 1988). Evidence for subduction beneath southern Africa since the Archean is limited to the Early Proterozoic Orange River Group, and the mid-Proterozoic Sinclair Formation in Namibia (Fig. 12). The Southern Cape Conductive Belt (Fig. 12) beneath the Namaqualand Metamorphic Complex may also be the result of Proterozoic subduction (Duncan, 1987). The Orange River Group (2.0 Ga) comprises a volcano-sedimentary sequence 8 km thick and extending approximately 350 km E-W (Reid et al., 1987). It was produced largely by subaerial volcanic activity, and is characterized by lateral and vertical discontinuity. Lavas, pyroclastic tuffs, volcanogenic sediments, breccias, agglomerates and conglomerates all occur. The
Volcanic rocks of Early Proterozoic greenstone belts
37
Fig. 12. Major tectonic provinces and subprovinces of southern Africa (from Reid et al., 1987).
volcanic rocks range in composition from mafic to felsic, but are dominated by andesites containing phenocrysts of plagioclase, augite, orthopyroxene, and magnetite. The volcanic rocks are penetrated by granitic rocks of the Vioolsdrif Batholith, and together these units form the Richtersveld Subprovince of the Namaqua Province of South Africa (Fig. 12). The boundaries of the subprovince are structural or intrusive, and it is affected by greenschist facies to amphibolite-facies metamorphism. All of the volcanic rocks show high overall concentrations of R E E with strong LREE fractionation (Reid et al., 1987, fig. 7). On N-MORB normalized diagrams (Reid et al., 1987, fig. 8) the mafic rocks, basaltic-andesites and andesites all have the same geochemical signatures. These show enrichments in Th (K, Rb, Ba) and the LREE, and depletions in Nb, Ta, Zr, Ti, Y, and the H R E E suggesting
38
TE. Smith
that the volcanics represent a (high-K) calcalkaline suite (Table 4). The initial Sr, Pb, and Nd isotopic ratios of least metamorphosed rocks of the entire volcanic suite (basalt to rhyolite) are similar. The E S ~ ( Tof) + lo , the high p value (high U/Pb ratio), and 6Nd cz 0 (compared to bulk earth values) suggest that the parental magmas were derived from a mantle source which had been affected by subduction zone metasomatism, a conclusion which is strongly supported by the presence of abundant andesites in the suite. The trace element and isotopic data suggest that the suite formed as a result of fractional crystallization, and that old Archean crust was not involved in the formation of the abundant andesites. Similar volcanic suites develop at continental margins such as the Andes, but ancient representatives of such arcs would be expected to be represented by high grade metamorphic and batholithic rocks as a result of very deep erosion. This observation, together with the absence of direct or indirect evidence of underlying Archean crust, led Reid et al. (1987) to suggest that the Orange River Group represents a mature oceanic island arc of Proterozoic age. The Indian Shield In India, the Early Proterozoic Aravalli Supergroup of Rajasthan is unconformable on an Archean banded gneiss complex and includes a thick basal unit of tholeiitic basalts, with minor picrites (komatiites), intercalated with thin bands of quartzite and conglomerate (Ahmad and Rajamani, 1991). The original mineralogy and textures of these volcanic rocks have been destroyed by deformation and metamorphism in the greenschist to amphibolite facies. The picrites are considered to have been formed by adiabatic melting of LREE-enriched sources at pressures up to 50 Kb. The tholeiites are considered to be derived from shallow lithosphere (low pressure) LREE-enriched sources, and cannot be related to the picrites in terms of partial melting or fractional crystallization. No felsic magmatism is reported associated with these mafic and ultramafic rocks. There is insufficient trace element chemistry to compare this CT sequence with the standard modern rock suites. However, the source characteristics and physical conditions of magma generation indicate that the basal Aravalli volcanism could be related to deep mantle plumes (Ahmad and Rajamani, 1991). The Dalma “Epidiorites” of the Chaibasa basin, Singhbhum Craton, eastern India, represent an Early Proterozoic metavolcanic sequence (1.6 Ga). It comprises ocean floor tholeiitic, alkaline, and komatiitic volcanics overlain by high-magnesian vitric tuffs and co-magmatic intrusive komatiites (Gupta e t al., 1980). They are underlain by phyllites, ferruginous shales, quartzites and iron formation. These authors consider that the Dalma belt was formed in a rift developed in Archean protocontinental crust, and is representative of a series of greenstone belts which surround the Archean Singhbhum Craton. They note that it has many features in common with the Circum-Superior belt of northern Quebec, and propose a tectonic model based on upwelling of the mantle. A small amount of trace element data has been reported from the tholeiites (Chakraborti and Bose, 1985) which
Volcanicroch of Early h.oterozoic greenstone belts
39
shows that they are more L R E E depleted than average MORB, and have lower average Y and Zr contents (Table 4). These authors conclude that the chemistry of the tholeiites and the overall geological setting strongly suggest eruption in a marginal basin.
China Four Early Proterozoic supracrustal successions have been described from the Zhongtiao Mountains of east-central China (Dazhong e t al., 1990, fig. 1). They are defined on the basis of unconformities and tectonic and intrusive contacts. The Jiangxian Group (2400-2200 Ma) which is >2000 m thick, and the Zhongtiao Group ( ~ 2 1 0 0Ma) which is ~ 7 0 0 0m thick, comprise chiefly metapelites, quartzites and marbles with small amounts of bimodal felsic to mafic volcanics dominated by felsic ash flow tuffs. The Danshansi Group is made up of approximately 700 m of quartzites and conglomerates. The Xiyanghe Group ( ~ 1 8 3 0Ma) comprises more than 5000 m of bimodal volcanics and red beds. The basalts of the Tbnguangyu Formation of the Jiangxian Group have moderately enriched LREE patterns, and those of Songjianshan Group, correlated with the Upper Jiangxian Group, have slightly LREE enriched patterns (Dazhong et al., 1990). Both suites have low concentrations of Nb and Y and La/Nb ratios >2.0 (Table 4). In the Zhongtiao Group REE, Nb and Y concentrations are higher, and the rocks show LREE enrichment and have negative Nb anomalies in their multi-element plots (Table 4). Basalts of the Xiyanghe Group have the highest concentrations of R E E and show L RE E enrichment and large negative Nb anomalies in their multi-element plots (Table 4). The high Ce/Nb and Th/Nb ratios of these rocks are similar to, or even higher than, those of continental within plate basalts (CWPB) and subduction-related basalts (Dazhong e t al., 1990, fig. 8). The ratios do not fall on mixing lines between MORB and average continental crust and are interpreted as indicating that the basalts contain a subduction component. This conclusion is not necessarily valid because it assumes that CWPB acquire their high Ce/Nb and Th/Nb ratios as a result of crustal contamination whereas many authors consider that these signatures are acquired as a result of the recycling of sediments through the mantle (see Marsh, 1987). The felsic rocks of all of the groups are enriched in LREE and have negative E u anomalies. When plotted on the Ta-Yb and Nb-Y diagrams of Pearce et al. (1984) they scatter between the arc and within plate fields. The volcanic rocks of the Jiangxian and Zhongtiao Groups are associated with quartzite-pelite-carbonate successions suggestive of deposition in a relatively stable sedimentary basin. The Danshansi Group is made up chiefly of quartzites and conglomerates, and the younger Xiyangye Group comprises bimodal volcanics and red beds, which are interpreted as being deposited in near-shore marine and fluvial, tectonically active basins. Dazhong et al. (1990) conclude that most of the data are consistent with the deposition of the Early Proterozoic supracrustal sequences, like those of the SW United States, in a continental-margin back-arc
40
TE. Smith
basin in which no oceanic crust was formed. Dazhong et al. (1990) imply that these sequences, plus the Xionger Group described below, were all formed during the accretion of successive arc-related sequences to the southern margin of the Asian continent. However, it is possible that some of the mafic volcanics of these sequences are CWPB derived from mantle through which sediments have been recycled. These suites may therefore represent all of the elements of a Wilson cycle. The Xionger Group is an Early to mid-Proterozoic (1710 Ma) volcanic suite 8 km thick, which is exposed for a length of 600 km in the Qinling Mountains of Central China (Chengzao, 1987). It is among the best preserved and most complete of the complexes described in this review. The Group, which is practically unmetamorphosed, overlies an Archean metamorphosed complex, the Taihua Group, and is overlain by terrigenous clastic rocks of the Ruyang Group and the Guandaokou Group. Four subdivisions can be recognised in the Xionger Group; (1) terrigenous clastic sediments, (2) andesites with a few basaltic-andesites, ( 3 ) rhyolites and dacites, (4) andesites with thin beds of terrigenous clastic rocks. The suite is dominated by andesites, the majority of which contain plagioclase phenocrysts with subordinate augite and hornblende. The distribution of Na 2 0 and K 2 0 in this volcanic suite indicates an evolutionary trend from basaltic andesite through alkali basalt, andesite, latite, and dacite, to rhyolite which is typical of active continental margins (Chengzao, 1987, table 1, fig. 8). The MORBnormalized multi-element plots (Chengzao, 1987, fig. 13), of the andesites, dacites, and rhyolites are all strongly LREE enriched ( a b l e 4) and have similar shapes. Overall abundances are lowest in the andesites and highest in the rhyolites. The multi-element patterns are all characterized by negative Nb anomalies, and are very similar to high-K andesites like those erupted at the active Andean margin in Central Chile (Chengzao, 1987, fig. 11). Regional synthesis (Chengzao, 19S7) suggests that the volcanic rocks of the Xionger Group resulted from plate subduction on the southern margin of the North China Microplate. The Kuanping Group, comprising a melange of metamorphosed oceanic tholeiites, greywacke and marble, is identified as a fossil subduction complex (accretionary wedge) located along the southern edge of the Xionger Group. The Guandaokou and Launchuan Groups contain carbonate flysch-like formations interpreted as sediments deposited in a forearc basin. The Ruyang and Luoya Groups are lithologically and stratigraphically similar to these forearc deposits but occur to the north of the Xionger Group. They are interpreted as sediments laid down in a back-arc basin.
The Australian Shield Proterozoic igneous rocks are widespread in Australia (Fig. 13), and become less abundant in the younger sequences. The following summary is based largely on Wyborn et al. (1987) who emphasize the differences between these suites and Phanerozoic subduction-related suites. They state that there are no ophiolites,
Volcanic rocks of Early Proterozoic greenstone belts
41
Fig. 13. Location of Australian Proterozoic domains (from W b o r n et al., 1987).
ocean floor crust, compositional equivalents of modern subduction suites, or any other evidence of a Wilson cycle. However, Wilson (1987) and Sivell (1988) use plate tectonic explanations for the volcanic sequences which they describe. The suites are mafic, felsic or bimodal, and there are no major intermediate sequences. Continent-wide volcanic events can be recognised at 2500-2000 Ma (including the Widgiemooltha Dyke Suite which is similar to the Early Proterozoic dyke swarms described below), 2000-1870 Ma, 1870-1820 Ma, 1810-1620 Ma, 1600-1200 Ma, and 1200-600 Ma. Depleted tholeiites are common in the early suites and enriched tholeiites and alkali rocks are dominant after 1800 Ma. The igneous rocks are most commonly associated with periods of extension and the formation of large sedimentary basins (2000-1870 Ma, 1810-1620 Ma, 1600-1200 Ma, and 1200-600 Ma), and less commonly with orogenic or deformational events (1870-1840 Ma, ca. 1600 Ma). The basins are characterized by an early rift phase, in which sediments and most of the igneous rocks are deposited in relatively narrow fault-bounded grabens. These sequences are overlain by
42
TE. Smith
widespread, dominantly sedimentary rocks which were deposited in a thermal subsidence phase. Anorogenic granites are intruded contemporaneously with, but spatially separated from, the sedimentary sequences. The volcanic suites of the rift sequences are dominantly bimodal, mafic, or alkaline associations, and those associated with deformation are dominated by felsic compositions. Some examples are described below. Basalts depleted in incompatible elements, having affinities to modern oceanic basalts, which erupted during the rift phase of the 2000-1870 Ma period include the Stag Creek Volcanics (Pine Creek Inlier) the Ding Dong Downs Volcanics and Biscay Formation (Halls Creek Inlier) Narracoota Volcanics (Glengarry subbasin) and the Woodward Dolerite (Halls Creek Inlier). The contemporaneous amphibolites of the Cook Gap Schist (Gawlor Craton) and the Zamu Dolerite (Pine Creek Inlier) have been interpreted as CT A number of layered igneous complexes, comprising gabbros, norite, troctolites, anorthosites, peridotites and pyroxenites were also emplaced at this time. Examples include the McIntosh and Panton Sills, the Alice Downs Ultrabasics (Halls Creek Inlier) and the Harry Anorthosite Complex. Felsic volcanism was limited in extent during this phase but occurs in the bimodal Ding Dong Downs suite, the Gerowie Tuff, and tuffs from the Mount Bonnie Formation (Pine Creek Inlier). The overlying flych sequence may have been derived from these felsic volcanics, both have low initial 87Sr/87Sr ratios and high Rb/Sr ratios, which indicate that they have had short crustal histories. The chemistry of these rocks is similar to that of the felsic rocks erupted in the 1870-1820 Ma phase described below. The majority of the igneous rocks of the 1870-1820 Ma phase are subaerial felsic volcanics which are often intruded by granitic intrusives, Most of the igneous rocks were emplaced immediately after the orogenic event, which was characterised by low-pressure metamorphism, and an unusually high geothermal gradient. Chemically the igneous rocks bear little resemblance to modern calcalkaline rocks. They range from 60 to 78% Si02 (predominantly 68-75%). They are high in K20, Rb, La, Ce, Th, and U, and very low in MgO, CaO, Ni, and Cr, as compared to Phanerozoic analogues or Archean tonalites and trondjhemites. They are interpreted as being derived from slightly older Proterozoic infracrustal sources. Cover rocks of the Harts Range meta-igneous complex and Irindina supracrustal assemblage structurally overlie Early Proterozoic basement sequences of the Strangways Metamorphic Complex, including both the Oonagalabi and Entia gneiss complexes, in the eastern Arunta Block (Fig. 13). The early formed supracrustal tholeiites are interlayered with pelitic and calc-silicate gneisses, ultramafic rocks, quartzites and iron formation of the Oonagalabi and Entia basement (deformed and metamorphosed ca. 1800 Ma). They display slight to moderate LREE enrichment and flat HREE distributions (Tible 4) and chondritic Zr/Nb and Zr/Y ratios (Sivell, 1988, fig. 5), and were formed by 15% to 30% partial melting of a relatively undepleted mantle source in an extensional ensialic-
Volcanic rocks of Early Proterozoic greensfone belts
43
rift environment. Later formed metabasites, intrude the basement Entia Gneiss Complex (together with granitoid intrusives of Cordilleran affinities). Their LREE and LILE enriched patterns, pronounced depletions in HFSE, La/Nb and Cr/Y ratios resemble those of basalts formed at convergent plate margins (Sivell, 1988, fig. 5). Meta-tholeiites (amphibolites) occur within laterally extensive interlayered pelitic, calcareous and quartzoze schist and gneisses that form the Harts Range cover sequence and resemble N-MORB and CT in their chemical compositions (Sivell, 1988, fig. 5). Tholeiites from the upper part of the cover sequence show LILE enrichment and pronounced decoupling between the LREE and HFSE resulting in high La/Nb ratios (Tmble 4). It is suggested that the source from which these rocks were derived has been re-enriched with a subduction-related component. The wide range of trace element compositions shown by these Early to mid-Proterozoic metamorphosed mafic volcanic rocks is indicative of chemical heterogeneity in their mantle sources. The sequence of volcanic rocks may result from repeated transition from rift-related to subduction-related tholeiitic magmatism representing a distinctive style of Proterozoic tectonics that involved alternating extensional and compressional ensialic orogenic phases (Sivell, 1988). Felsic rocks which are predominant in the bimodal Gawler Range Volcanics (1600-1500 Ma) were formed by partial melting of a refractory mafic-intermediate lower crustal source (Giles, 1988). This was the result of heat supplied, during crustal rifting and underplating, by mantle diapirism which also produced the mafic volcanics. It is suggested that this process of crustal building is fundamentally different to that which produced modern calcalkaline arcs. The volcanic rocks of the Lower Proterozoic Willyama Complex (1820 f 60 Ma) of the Broken Hill Block (Fig. 11) comprise a complete sequence from mafic through intermediate to felsic (James et al., 1987). The mafic rocks show Th and LREE enrichment and Nb and Ti depletion (James et al., 1987 fig. 4) on N-MORB-normalized plots ( a b l e 4). James et al. (1987) conclude that this tholeiitic suite was formed by partial melting of a slightly enriched N-MORB source followed by fractional crystallization and crustal contamination. They suggest that some intermediate rocks were formed by fractional crystallization, but major crustal involvement may be necessary to produce the large quantities of felsic rocks present in the lower part of the complex. The Mount Isa Inlier (Fig. 6) contains Proterozoic volcanic rocks from four phases of igneous activity (Wilson, 1987), examples of which are briefly described below. The Leichhardt Metamorphics (1865 Ma), comprise metamorphosed rhyolites and dacites. Chemically they have high KzO, Ba, Rb, Sr, and LREE and low TiO2, MnO, Nb, Y , Z n and Zr, compared to most modern felsic volcanics and are similar to orogenic Cainozoic rhyolites (Ewart, 1979). Wilson (1987) has suggested that they formed by partial melting of a pelitic source. Sm-Nd model ages suggest a maximum crustal residence time of 200 Ma (Page et al., 1984). The Magna Lynn Metabasalt and Argylla Formation (1777 Ma) contain metabasalt, andesite, dacite, rhyolite, and clastic sediments. The Magna Lynn Metabasalt is dominated by
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intermediate compositions near the base (53-56% Si02) and mafic compositions (<49% SiOz) near the top. The suite is subalkaline and shows moderate iron enrichment (possibly tholeiitic), has Ti02 between 1.5% and 1.8%, and K20, and P205 that are generally lower than those of within-plate basalts. Wilson (1987) suggests that these mafic rocks, containing three times chondritic concentrations of incompatible elements, were formed by 30% partial melting of a mantle source followed by fractional crystallization. The vast majority of samples from the Argylla Formation are quartz-normative alkali rhyolites which most closely resemble the felsic rocks of Cainozoic bimodal suites. They have high K 2 0 (>5%) and total iron (2%-7%) and H R E E contents (15-20 x chondrite), (Wilson, 1987, figs. 4a and 5). Petrogenetic modelling suggests that these felsic rocks were formed by partial melting of a shale source at relatively shallow depth (Wilson, 1987). The Eastern Creek Volcanics (1700 Ma) are dominantly made up of metabasalt and metamorphosed clastic sediments. There are two mafic volcanic sequences in the formation, the lower member is characterized by Ti02 values around 3%, and the upper member by values between 1 % and 1.7%. In each member the more primitive rocks resemble OFT and the more evolved rocks are classified as C T According to Wilson et al. (1984) the K/Rb, Zr/Nb, and Zr/Hf ratios of the mafic rocks are consistent with derivation from a mantle source. The within-suite variation can be accounted for by fractional crystallization of olivine, plagioclase, clinopyroxene and minor magnetite, in magma reservoirs in the base of the crust. The final cycle was formed between 1680 and 1610 Ma and is characterized by highly potassic volcanics, ranging from mafic lavas to felsic tuffs, including the Fiery Creek Volcanics, which are interbedded with apparently unmetamorphosed clastic sediments and dolomites. The Zr/TiO2 :Nb/Y ratios indicate that the mafic rocks of the Fiery Creek Volcanics are alkaline basalts, a conclusion consistent with the presence of relict olivine in some samples, and with their REE patterns. The felsic rocks of this unit are trachytes and alkali trachytes with enriched LREE patterns. Their composition is unlike that of any modern volcanic suite and is considered to be the result of alteration. Partial melting of pelitic material to yield the felsic volcanics of the Leichhardt Metamorphics and the Argylla Formation implies either the deposition of a thick pile of sediments or major tectonic thickening of the crust. Either process could take place at a continental margin, in a back-arc basin or in association with a subduction zone. In addition, a mantle convection driven model involving collision is considered to be the most likely mechanism to have produced the 80% shortening of parts of the Mount Isa Inlier (Wilson, 1987).
DISCUSSION AND CONCLUSIONS
Tholeiitic and boninitic (picritic)-noritic dyke swarms were emplaced in the continental crust worldwide, during the period 2.5-2.0 Ga, indicating the existence of large scale chemical heterogeneities in the mantle at that time (Weaver and
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Tarney, 1981, 1983; Hall et a]., 1987; Sheraton et al., 1987; Sial et al., 1987; Zirney and Weaver, 1987). The heterogeneities probably formed initially as a result of the voluminous extraction of tholeiitic and komatiitic melts from the mantle and the subsequent formation of the Archean cratons and associated sub-continental lithosphere (Sun e t al., 1979; Condie, 1985; Hall e t al., 1987; Tarney and Weaver, 1987). Similar associations of picritic lavas, showing LREE enrichment and Nb depletion, associated with enriched tholeiites, also occur in the Archean Ameralik dyke suite and in younger continental flood basalts in the. Karoo and Deccan (Xirney and Weaver, 1987). In addition, thirteen Kaapvaal volcanic sequences ranging in age from 3000 to 2100 Ma, and basalts from three of the Karoo volcanic provinces (lS0 Ma) all have identical trace element signatures, showing LREE enrichment and Nb and Ti depletion (Myers e t al., 1987). The similarities in the range of trace element compositions of the Early Proterozoic dyke suites, the Ameralik dykes, Kaapvaal volcanics and the Karoo and Deccan basalts, produced over the past 3000 Ma, suggest that the processes of development of sub-continental lithosphere have followed the same general pattern throughout the period. The processes must involve re-supply of LILE and LREE to the mantle through subduction, or some related process capable of recycling crustal materials into the mantle (Hall et al., 1987; Tarney and Weaver, 1987). The operation of plate tectonic processes during the Phanerozoic is well established, and the prolonged re-supply of LILE and LREE to the mantle supports the view (Kerr, 1991) that similar tectonic processes have taken place continuously since the Archean. It has been suggested, that intracontincntal rifting and underplating of Archean cratons, and not plate tectonics, best explain the Proterozoic orogenesis and tectonic processes occurring in Australia, and possibly world wide (for summary see Etheridge et al., 1987; Wyborn et al., 1987). This view has been criticised (McCulloch, 1957) and some of the studies quoted above (Wilson, 1987; Sivell, 1988) imply that a plate tectonic interpretation may better fit the Australian data. In addition, contrary to the view expressed by Etheridge et al. (1987) Early Proterozoic island arc sequences containing abundant andesites, and ophiolite complexes, do exist. The othcr data used to support of the hypothesis (Etheridge et al., 1987) is not compelling and it seems unlikely that the thermotectonic processes taking place in Australia were different to those of the rest of the world during the Early Proterozoic. In the Early Proterozoic large areas of stabilized crust began to behave as internally rigid units similar to, but smaller than, modern continental plates. Early Proterozoic volcanic suites, that have striking similar characteristics to each of the different modern volcanic associations, are preserved on these stable cratons, and have been described above (see also Pharaoh et al., 19S7). Many of the suites contain komatiites which are virtually absent from Phanerozic sequences, which implies that there has been some secular change in volcanic activity since the Proterozoic. Pharaoh and Pearce (1984) conclude that despite secular changes
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in the geochemistry and thermal structure of the crust and mantle since the Proterozoic, the chemical compositions of volcanic rocks, up to 2.0 Ga old, may be compared directly with those of modern suites. On balance geological and structural relationships, and geochemistry (Condie, 1989), strongly suggest that the Early Proterozoic suites erupted in thermotectonic environments similar to those of Phanerozoic sequences. The significance of the occurrence of the komatiites is assessed below. The komatiites which occur in Early Proterozoic volcanic sequences are generally associated with tholeiitic sequences, and may be LREE-enriched or depleted, even within the same formations (see Bakkilvarri Formation of the Baltic Shield Fig. 4). However, their trace element contents may not be representative of the original magma because they erupt at very high temperatures and are capable of assimilating any type of wall rock (Kerr, 1991). The Early Proterozoic komatiites are less magnesian (MgO 16%-20%) and less abundant than their Archean equivalents (MgO > 20%). The maximum MgO content of lavas is both temperatureand pressure-dependent, and Ahmad and Rajamani (1991) have suggested that the decrease in the volume and MgO content of komatiites, from the Archean to the Early Proterozoic, is the result of the decreasing heat budget in the earth’s mantle with age (see also GaAl and Gorbatschev, 1987). In addition, they believe that Aravalli komatiites (<20 wt.% MgO) may have been formed by partial melting at pressures as high as 50 kb, when deep mantle plumes rose into the lithosphere. They suggest that the high-MgO magma caused melting of the lithosphere to produce the closely associated tholeiitic sequence. Thus the presence of komatiites may reflect the greater importance of plume tectonics in the Early Proterozoic as compared to the Phanerozoic. The secular distribution of komatiites and the striking similarities between Early Proterozoic and Phanerozoic volcanic suites suggest that a modified form of plate tectonics, involving more plume activity, occurred during the Early Proterozoic.
ACKNOWLEDGEMENTS
Thanks are due to Drs. K.C. Condie and Y. Kahkonen, and an unknown reviewer, for constructive comments which greatly improved the final manuscript. Figure 1was reproduced with permission of the American Geophysical Union and the Geological Society of London gave permission for the use of Figures 2, 7, 8, 9, 10,12 and 13.
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Lees, G.J., Roach, R.A., Shufflebotham, M.M. and Griffiths, N.H., 1987. Upper Proterozoic basaltic volcanism in the northern Massif Armoricain, France. In: TC. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC. London, Spec. Publ., 33: 503-524. Le Roex, A.P., Dick, H.J.B., Erlank, A.J., Reid, A.M., Frey, FA., and Hart, S.R., 1983. Geochemistry, mineralogy and petrogenesis of lavas erupted along the southwest Indian Ridge between the Bouvet triple junction and 11 degrees east. J. Petrol., 24: 267-318. Lewry, J.E, Macdonald, R., Livesey, C., Meyer, M., Van Schmus, W.R. and Bickford, M.E., 1987. U-Pb geochronology of accreted terranes in the Trans-Hudson Orogen, Northern Saskatchewan, Canada. In: T.C. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC.London, Spec. Publ., 3 3 147-166. Lofgren, C., 1979. Do leptites represent Precambrian island arc rocks? Lithos, 1 2 159-165. Ludden, J., Gelinas, L. and ’Rudel, P., 1982. Archean metavolcanics from the Rouyn-Noranda district, Abitibi Greenstone Belt, Quebec, 2. Mobility of trace elements and petrogenetic constraints. Can. J. Earth Sci., 19: 2276-2287. Marsh, J.S., 1987. Basalt geochemistry and tectonic discrimination within continental flood basalt provinces. J. Volcanol. Geotherm. Res., 32: 35-49. McClennan, S.M. and Taylor, S.R., 1991. Sedimentary rocks and crustal evolution: tectonic setting and secular trends. J. Geol., 99: 1-21. McCulloch, M.T., 1987. Sm-Nd isotopic constraints on the evolution of Precambrian crust in the Australian continent. In: A. Kroner (Editor), Proterozoic Lithosphere Evolution. Am. Geophys. Union, Geodyn. Ser., 17: 115-130. McKenzie, D.P. and Weiss, N., 1975. Speculations on the thermal and tectonic history of the earth. Geol. J. R. Astron. SOC.,42: 131-174. Morgan, G.E. and Briden, J.C., 1981. Aspects of Precambrian palaeomagnetism with new data from the Limpopo mobile belt and Kaapvaal craton in southern Africa. Phys. Earth Planet. Inter., 24: 142-168. Myers, R.E., Cawthorn, R.G., McCarthy, T.S. and Anhaeusser, C.R., 1987. Fundamental uniformity in the trace element patterns of the volcanics of the Kaapvaal Craton from 3000 to 2100 Ma: evidence for the lithospheric origin of these continental tholeiites. In: T.C. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC. London, Spec. Publ., 33: 315-325. N’Gom, P.M., 1985. Contribution B I’ttude de la sdrie birrimienne de Mako dans le secteur Sabodala (Sinkgal Oriental). Thkse 3e cycle, Universitt de Nancy I, 134 pp Nisbet, E.G. and Fowler, C.M.R., 1983. Model for Archaean plate tectonics. Geology, 11,376-379. Page, R.W., McCulloch, M.T. and Black, L.P., 1984. Isotopic record of major Precambrian events in Australia. Proc. 27th Int. Geol. Congr., 5: 25-72. Park, A X , 1985. Accretion tectonism in the Proterozoic Svecokarelides of the Baltic Shield. Geology, 1 3 725-729. Park, R.G., 1988. Geological Structures and Moving Plates. Blackie, Glasgow, 337 pp. Patchett, P.J. and Bridgwater, D., 1984. Origin of continental crust of 1.9-1.7 Ga age defined by Nd isotopes in the Keltidian terrain of South Greenland. Contrib. Mineral. Petrol., 87: 311-318. Patchett, J. and Arndt, N.T., 1986. Nd isotopes and tectonics of 1.9-1.7 Ga crustal genesis. Earth Planet. Sci. Lett., 78: 329-338. Pearce, J.A., 1982. Trace element characteristics of lavas from destructive plate boundaries. In: R.S. Thorpe (Editor), Andesites. Wiley, New York, N.Y., pp. 525-548 Pearce, J.A., 1983. Role of the sub-continental lithosphere in magma genesis at active continental margins. In: C.J. Hawkesworth and M.J. Norry (Editors), Continental Basalts and Mantle Xenoliths.
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Shiva, Nantwich, pp. 230-249 Pearce, J.A., 1987. An expert system for the tectonic characterization of ancient volcanic rocks. J. Volcanol. Geotherm. Res., 3 2 51-65. Pearce, J.A. and Norry, M.J., 1979. Petrogenetic implications of Ti, Zr, Y, and Nb variations in volcanic rocks. Contrib. Mineral. Petrol., 6 9 33-47. Pearce, J.A., Harris, N.B.W. and Tindle, A.G., 1984. Trace element discrimination diagrams for the tectonic interpretation of granitic rocks. J. Petrol., 25: 956-983. Peck, D.C. and Smith, TE., 1989. The geology and geochemistry of an early Proterozoic volcanic-arc association at Cartwright lake: Lynn Lake Greenstone Belt, northwestern Manitoba. Can. J. Earth Sci., 26: 716-736. Perfit, M.R., Gust, D.A., Bence, A.E., Arculus, R.J. and Taylor, S.R., 1981. Chemical characteristics of island-arc basalts: implications for mantle sources. Chem. Geol., 3 0 227-256. Pharaoh, TC., 1985. Volcanic and geochemical stratigraphy of the Nussir Group of Arctic Noway - an early Proterozoic greenstone suite. J. Geol. SOC.London, 1 4 2 259-278. Pharaoh, TC. and Pearce, J.A., 1984. Geochemical evidence for the geotectonic setting of the early Proterozoic metavolcanic sequences in Lapland. Precambrian Res., 1 0 283-309. Pharaoh, TC., Warren, A. and Walsh, N.J., 1987. Early Proterozoic metavolcanic suites of the northernmost part of the Baltic Shield. In: TC. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC.London, Spec. Publ., 33: 41-58. Pharaoh, TC. and Brewer, T.S. 1990. Spatial and temporal diversity of early Proterozoic volcanic sequences - comparisons between the Baltic and Laurentian Shields. Precambrian Res., 47: 169189. Picard, C., Lamothe, D., Piboule, M. and Oliver, R., 1990. Magmatic and geotectonic evolution of a Proterozic oceanic basin system: the Cape Smith thrust-fold belt (New Quebec). Precambrian Res., 47: 223-249. Piper, J.D.A., 1982. The Precambrian palaeomagnetic record: the case for the Proterozoic supercontinent. Earth Planet. Sci. Lett., 5 9 61-89. Reid, D.I., Welke, H.J., Erlank, A.J. and Moyes, A,, 1987. The Orange River Group: a major Proterozoic calcalkaline volcanic belt in the western Namaqua Province, southern Africa. In: TC. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC.London, Spec. Publ., 33 327-346. Renner, R. and Gibbs, A.K., 1987. Geochemistry and petrology of the metavolcanic rocks of the early Proterozoic Mazaruni greenstone belt, northern Guyana. In: TC. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC. London, Spec. Publ., 33: 289-309. Robertson, J.M. and Condie, K.C., 1989. Geology and geochemistry of early Proterozoic volcanic and subvolcanic rocks of the Pecos greenstone belt, Sangre de Cristo Mountains, New Mexico. In: J.A. Grambling and B.J. Tewksbury (Editors), Proterozoic Geology of the Southern Rocky Mountains: Boulder, Colorado, Geol. SOC.Am., Spec. Pap., 235: 119-145. Saunders, A.D. and Tarney, J., 1984. Geochemical characteristics of basaltic volcanism within back-arc basins. In: B.P. Kokelaar and M.E Howells (Editors), Marginal Basin Geology. Geol. SOC.London, Spec. Publ., 1 6 59-76. Scott, D.J., St-Onge, M.R., Lucas, S.B. and Helmstaedt, H., 1988. Purtuniq ophiolite: oceanic crust preserved in the ca. 1.9 Ga Cape Smith thrust-fold belt, northern Quebec. GAC-MAC-CSPG Joint Annual Meeting, Program with Abstracts, 13. Sheraton, J.W., Thonison, J.W. and Collerson, K.D. 1987. Mafic dyke swarms of Antarctica. In: H.C. Halls and W.E Fahrig (Editors), Mafic Dyke Swarms. Geol. Assoc. Can., Spec. Pap., 34: 419-432.
Volcanic rocks of Early Proterozoic greenstone belts
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Sial, AN., Oliviera, E.P. and Choudhari, A., 1987. Mafic dyke swarms of Brazil. In: H.C. Halls and W.F. Fahrig (Editors), Mafic Dyke Swarms. Geol. Assoc. Can., Spec. Pap., 3 4 467-481. Sims, P.K., Van Schmus, W.R., Schulz, K.J. and Peterman, Z.E., 1989. Tectonostratigrphic evolution of the early Proterozoic Wisconsin magmatic terranes of the Penokean orogen. Can. 3. Earth Sci., 26: 2145-2158. Sivell, W.J., 1988. Geochemistry of the metatholeiites from the Harts Range, Central Australia: implications for mantle source heterogeneity in a Proterozoic mobile belt. Precambrian Res., 40/41: 261- 275. Skiold, T and Cliff, R., 1984. Sm-Nd and U-Pb dating of Early Proterozoic mafic-felsic volcanism in northernmost Sweden. Precambrian Res., 26: 1-13. Smith, TE. and Holm, P.E., 1987. The trace element geochemistry of metavolcanics and dykes from the Central Metasedimentary Belt of the Grenville province, southeastern Ontario, Canada. In: TC. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. Soc. London, Spec. Publ., 33: 453-470. Sun, S.S., 1980. Lead isotope study of young volcanic rocks from mid-ocean ridges, ocean islands and island arcs. Philos. Trans. R. Soc. London, 1297: 409-445. Sun, S.S., Nesbitt, R.W. and Sharaskin, A.Ya., 1979. Geochemical characteristics of mid-ocean ridge basalts. Earth Planet. Sci. Lett., 44: 119-138. Syme, E.C., 1985. Geochemistry of metavolcanic rocks in the Lynn Lake Belt. Manitoba Department of Energy and Mines, Mineral Resources Division, Geological Report 84-1, 84 pp. Tarney, J. and Windley, B.E, 1977. Chemistry, thermal gradients and evolution of the lower continental crust. J. Geol. SOC.London, 134: 153-172. Tarney, J. and Weaver, B.L., 1987. Geochemistry and petrogenesis of early Proterozoic dyke swarms. In: H.C. Halls and W.E Fahrig (Editors), Mafic Dyke Swarms. Geol. Assoc. Can., Spec. Pap., 34: 81-94. Taylor, S.R., 1987. Geochemical and petrological significance of the Archaean-Proterozoic boundary. In: T.C. Pharaoh, R.D., Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC.London, Spec. Publ., 33: 3-8. Thompson, R.N., Morrison, M.A., Dickin, A.P., Hendry, G.L., 1983. Continental flood basalts ... arachnids rule OK? In: C.J. Hawkesworth and M.J. Norry (Editors), Continental Basalts and Mantle Xenoliths. Shiva, Nantwich, pp. 158-185. Vachette, M., 1964. Essai de synth&sedes dtterminations d’dges radiomttriques de formations cristallines de l’Ouest Africain (CBte d’Ivoire, Mauritanie, Niger). Ann. Fac. Sci. Clermont Ferrand, 25, 7 PP. Vivallo, W and Claesson, L-A., 1987. Intra-arc rifting and massive sulphide mineralization in an early Proterozoic volcanic arc, Skellefte district, northern Sweden. In: T.C. Pharaoh, R.D., Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites, Geol. Soc. London, Spec. Publ., 33: 69-80. Watters, B.R. and Pearce, J.A., 1987. Metavolcanic rocks of the La Ronge Domain in the Churchill province, Saskatchewan: geochemical evidence for a volcanic arc origin. In: TC. Pharaoh, R.D., Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. Soc. London, Spec. Publ., 33: 167-182. Weaver, B.L. and Tarney, J., 1981. The Scourie dyke suite: Petrogenesis and geochemical nature of the Proterozoic sub-continental mantle. Contrib. Mineral. Petrol., 7 8 175-188. Weaver, B.L. and Tarney, J., 1982. Andesitic magmatism and continental growth. In: R.S. Thorpe (Editor), Andesites. Wiley, London, pp. 639-661. Weaver, B.L. and Tarney, J., 1983. Chemistry of the sub-continental mantle: inferences from Archaean and Proterozoic dykes and continental flood basalts. In: C.J. Hawkesworth and M.J. Norry (Editors), Continental Basalts and Mantle Xenoliths. Shiva, Nantwich, pp. 209-229.
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Welin, E., 1987. The depositional evolution of the Svemfennian supracrustal sequences in Finland and Sweden. Precambrian Res., 35: 95-114. Wilson, I.H., 1987. Geochemistry of Proterozoic volcanics, Mount Isa Inlier, Australia. In: TC. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. Soc. London, Spec. Publ., 33 409-423. Wilson, I.H., Derrick, G.M. and Perkin, D.J., 1984. Eastern Creek Volcanics: their geochemistry and possible role in copper mineralisation at Mount Isa, Queensland. BMR J. Aust. Geol. Geophys., 9: 317-328. Wilson, M.R., Hamilton, P.J., Fralick, A.E., Aftalion, M. and Michard, A., 1985. Srn-Nd, U-Pb and 0 isotope systematics of granites and Proterozoic crustal evolution in Sweden. Earth Planet. Sci. Lett., 7 2 376-388. Wilson, M.R., Ohlander, B., Cuney, M. and Hamilton, P.J., 1987. Geodynamic significance of contrasting granitoid types in northern Sweden. In: A. Kroner (Editor), Proterozoic Lithosphere Evolution. Am. Geophys. Union, Geodyn. Ser., 17: 161-173. Winchester, J.A., Max, M.D. and Long, C.B., 1987. Trace element geochemical correlation in the reworked Proterozoic Dalradian metavolcanic suites of the western Ox Mountain and NW Mayo Inliers, Ireland. In: TC. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. Soc. London, Spec. Publ., 33: 489-502. Wood, D.A., 1980. The application of a Th-Hf-Ta diagram to problems of tectonomagmatic classification and to establishing the nature of crustal contamination of basaltic lavas of the British TertiaIy volcanic province. Earth Planet. Sci. Lett., 50: 11-30. Wood, D.A., Joron, J.L. and Treuil, M., 1979. A re-appraisal of the use of trace elements to classify and discriminate between magma series erupted in different tectonic settings. Earth Planet. Sci. Lett., 45: 326-336. Wyborn, L.A.I., Page, R.W. and Parker, A.J., 1987. Geochemical and geochronological signatures in Australian Proterozoic igneous rocks. In: TC. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC.London, Spec. Publ., 33 377-394. Zonou, S., 1987. Les formations leptyno-amphibolitiques et le complexe volcanique et volcano-sCdimentaire du ProtCrozoique infCrieur de Bouroum Nord (Burkina Faso, Afrique de I’Ouest). Etude pttrographique, gCochimique: approche pCtrogCnCtique et Cvolution gkodynamique. Thkse, UniversitC de Nancy I, 294 pp.
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Chapter 2
THE PROTEROZOIC OPHIOLITE PROBLEM H.H. HELMSTAEDT and D.J. SCOTT
INTRODUCTION
Interpreted as fragments of ancient oceanic lithosphere, ophiolite complexes play an important role in extrapolating processes of plate tectonics into preMesozoic times (e.g., Coleman, 1977, 1984; Moores, 1982). As their sheeted dikes appear to represent physical evidence for sea-floor spreading, the presence of ophiolites in orogenic belts is considered as a key indicator for the operation of the Wilson Cycle, suggesting that geosynclines and orogenic belts evolve as natural consequence of plate motions by the opening and closing of oceans (Wilson, 1968). In order to qualify as a section through the floor of a true ocean basin, an ophiolite was thought to have to conform to the Penrose conference definition (Anonymous, 1972), meaning that it should be composed, from bottom to top, of tectonized harzburgite, a layered mafic-ultramafic plutonic complex grading upward into high-level gabbros and sheeted dikes, overlain by massive and pillowed mafic volcanic rocks and pelagic sedimentary rocks. However, ophiolites are highly diverse, not only in chemical characteristics and internal structure, but also with respect to the tectonic environment of formation, many having originated in island arc or marginal basin settings rather than in full-fledged ocean basins (e.g., Moores, 1982; Coleman, 1984; Searle and Stevens, 1984). An expanded version of the original Penrose ophiolite definition, proposed by Moores (1982), allows greater precision in inferring origin and emplacement by also considering the pre-emplacement rocks (crystalline basement, platformal sedimentary rocks and continental rise deposits) and the types of post-emplacement deposits associated with the ophiolites. In spite of the increased attention given to these associated rock types, however, the tectonic setting of many well-studied ophiolites remains controversial (e.g., Searle, 1991). While their sheeted mafic dike complexes suggest spreading in an oceanic environment, only in exceptional circumstances (e.g., Oman) provide these ophiolites some minimum limits for the width of former ocean basins (Glennie et al., 1974). Although most pre-Mesozoic ophiolites are incomplete and structurally dismembered, field evidence for the operation of the Wilson Cycle during the Paleozoic (Dewey and Spall, 1975) and latest Proterozoic (e.g., LeBlanc, 1981) is compelling. However, the apparent absence of ophiolites older than about 800-900 Ma, conforming to the Penrose conference definition, left geologists divided over
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H.H. Helmstaedt and D.J.Scott
the importance of plate tectonics during the earlier Proterozoic. In spite of Wilson’s hypothesis that, in intracontinental orogenic belts, most of the evidence of earlier oceans is destroyed by subduction (Wilson, 1968) and sutures arc difficult to recognize (see also Dewey, 1977), some considered the lack of ophiolites in older Proterozoic belts as inconsistent with the operation of Wilson Cycles, reasoning that these belts resulted entirely from ensialic orogenic processes (e.g., Kroner, 1977, 1981, 1983; Baer, 1978, 1981; Glickson, 1981; McCall, 1981). However, the similarities in sedimentary and igneous rock records and in tectonic styles between Phanerozoic and Proterozoic orogens led others to the conclusion that the Wilson Cycle operated since at least 2000 Ma (e.g., Burke et al., 1976; Hoffman, 1980, 1988; Lewry, 1981; Windley, 1981, 1984; Condie, 198%; Park et al., 1984). In a most perceptive discussion of the Proterozoic ophiolite problem, Moores (1986) suggested three possible reasons for the lack of ophiolites in older Proterozoic orogenic belts: (1) Plate tectonics as now conceived was not operating in mid-lower Proterozoic time (1-2.5 Ga). (2) Ophiolites as now known are present in Proterozoic suture belts but have gone unrecognized. (3) Ophiolites arc not present in Proterozoic suture belts older than 1 Ga. Instead their place is taken by rock sequences that resemble ophiolites but differ in significant respects from Phanerozoic examples. He dismissed the first possibility, because most of the “petrotectonic suites” characteristic of plate tectonics arc present in Proterozoic belts. He also considered the second possibility unlikely, as he thought that most Proterozoic suture belts had been mapped in sufficient detail to recognize ophiolites. In support of the third possibility, he cited examples from the circum-Superior belt in Canada, the Goias mafic-ultramafic complexes in Brazil, and the Giles complex of Western Australia. He recognized that in all three examples mafic-ultramafic complexes arc present in the position occupied by ophiolites in Phanerozoic orogenic belts and that each resembles the upper, magmatic part of an ophiolite, except that this part is much thicker. He concluded that the Proterozoic magmatic oceanic crust must have been thicker and that during the tectonic emplacement of the ocean floor over the continental margin, not enough section could be lifted to expose the suboceanic mantle. Using a theoretical approach, Hynes (1987) came to the opposite conclusion that a hotter Early Proterozoic mantle should lead to the formation of thinner oceanic plates. This, in turn, should inhibit back-arc spreading, one of the principal mechanisms of ophiolite formation (Moores, 1982; Coleman, 1984), providing a possible explanation for the apparent absence of Proterozoic ophiolites. However, this conclusion is at variance with findings of other authors (e.g., Burke et al., 1976; Windley, 1981; Bickle, 1986) that greater heat loss in the Archean should have led to the formation of a thicker oceanic crust and that extensive rifting in marginal basins gave rise to greenstone belts, considered as proto-ophiolites by Windley (1981).
The Proterozoic ophioliteproblem
57
The recent discovery of three nearly 2-Ga-old ophiolites throws a new light on the Proterozoic ophiolite debate. The Jormua complex in northeastern Finland (Kontinen, 1987), is structurally dismembered but contains all the essential ingredients of the Penrose conference definition, including serpentinized peridotite. The combined thickness of the reconstructed crustal portion (see Fig. 1) is in line with that of Phanerozoic ophiolites. The Purtuniq ophiolite in the Cape Smith belt, Northern Quebec, Canada (St-Onge et al., 1988; Scott et al., 1989; Scott, 1990; Scott and Bickle, 1991) is allochthonous and structurally dismembered, but the reconstructed section is also incomplete, representing only the crustal portion of the idealized Penrose ophiolite. However, this crustal portion is significantly thicker than that of Phanerozoic ophiolites (Fig. 1). The 1.73 Ga Payson ophiolite in central Arizona, on the other hand, is structurally more or less intact, being preserved in situ in an intra-arc setting (Dann, 1991). The fact that these older ophiolites had gone unrecognized for so long (possibility 2 of Moores, 1986) underscores the continued vital role of geological field work in formulating Precambrian tectonic models. The present contribution thus highlights some of the field aspects of Proterozoic ophiolite occurrences. Our review will concentrate on a comparison of the youngest Proterozoic ophiolites from the Late Proterozoic Pan-African belts with the much older ophiolites from the Early Proterozoic belts of the Canadian and Fennoscandian shields. We will examine the geotectonic significance of these ophiolites and discuss whether they represent possible links between Archean greenstone belts (proto-ophiolites of Windley, 1981) and Phanerozoic ophiolites.
PROBLEM OF PRESERVATION OF PROTEROZOIC OPHIOLITES
By definition, ophiolites occur in or near suture zones where the inferred remnants of ocean floor have escaped subduction. As shown by Dewey (1977), sutures resulting from continental collision are highly complex structures, because terminal collision is generally preceded by suturing on various smaller scales, involving the collision of arcs and microcontinents as well as the closure of marginal basins. As studies of Phanerozoic orogenic belts have shown, the preservation of relatively young and buoyant marginal-basin floor in smaller-scale mini-sutures appears to be more common than remnants of much older and cooler main-ocean floor in a mega-suture (e.g., Burke et al., 1976, 1977). The observation that ophiolites, obducted mainly during mini-suturing, have the tendency to become isolated at high structural levels above a downward-tapering mega-suture (Dewey, 1977), explains the small chances for their survival in the cryptic terminal sutures of deeply eroded Proterozoic orogens. It is therefore not surprising that from approximately 25 Proterozoic orogenic belts, identified as possible suture zones (Burke et al., 1977), only one fully developed and two dismembered ophiolites had been described, all of these occurring in the late Proterozoic
YELLOWKNIFE 2700 Ma
12 11 10 9
8
SAMAIL 100 Ma
MODERN OCEAN FLOOR
PAYSON 175 Ma
. . . . JABAL ESS
700 M a
Oceanic Layer
780 Ma
Dacite breccia Mafic volcanic rocks
5
I c n w
a
BOU AZZER
1970 Ma
L]Sedimentary rocks
E
Y
Y
4
JORMUA
-
6 VOURINOS
1730 Ma
?
7
t
3
2 1
0 MOHO -1 Upper
Mantle
-2 -3 -4
-5 Sheeted dikes
-6 Gabbro, isotropic or layered Ultramafic cumuhte rocks Upper mantle
- peridotite, serpentmite
Pre-ophiolitic granodiorite
-7 -8 -9
-10
Fig. 1. Simplified composite sections to compare internal structure and thickness of Proterozoic ophiolites to those of Phanerozoic ophiolites, modern oceanic crust, and mafic section of Archean Yellowknife greenstone belt. Datum level is contact between tectonized peridotites and layered cumulates. Vounnos, Samail, and modem Ocean floor after Coleman (1977);Bou Azzer after LeBlanc (1981);Jabal Ess after Shanti and Roobol(l979); Payson after Dann (1991); Jormua after Kontinen (1987); Purtuniq after Scott (1990) (see also Fig. 7); Yellowknife after Helmstaedt et al. (1986). Vertical black lines on Bou Azzer section indicate dikes in upper part of peridotite; dikes also occur in the volcanic rocks. The upper part of the section shows intercalated mafic flows and sedimentary rocks. Upper contact of Purtuniq section is a thrust (see Figs. 6 and 7), and the pre-thrusting relationship of the sedimentary Spartan Group to the mafic volcanic section is unknown.
Xke Proterozoic ophioliteproblem
59
Pan-African sutures of northern Africa (LeBlanc, 1975, 1976) and Arabia (Bakor et al., 1976). Considering the evidence put forward by earlier authors, that the Wilson Cycle operated since the Early Proterozoic (for a discussion, see also Windley, 1984), it is also no surprise that a more focused search following the identification of possible sutures led to the discovery of numerous other ophiolites and ophiolite fragments. No doubt, continued search will ’unearth’ pieces of even more ancient ocean floor, rendering conclusions based on the presently known Proterozoic ophiolites obsolete.
PAN-AFRICAN OPHIOLITES
Apart from a few additions along its northeastern and southern margins (Fig. 2), the African continent was assembled during the Pan-African orogeny at the end of the Proterozoic. Although the “Pan-African” event was first thought of as a period of major basement reactivation (Kennedy, 1964), it was proposed later that a number of Pan-African belts are comparable to Phanerozoic orogens and may represent major sutures (e.g., Burke et al., 1977). However, this interpretation found general acceptance only after ophiolites and remnants of ophiolites were discovered in the trans-Saharan belt of northwestern Africa (e.g., LeBlanc, 1976), and a number of geologists remained unconvinced that similar belts without ophiolites could be interpreted in terms of a Wilson Cycle (e.g., Kroner, 1983). Kroner (1979, 1980) referred to Pan-African belts marking possible suture zones as “orogenic belts” in order to distinguish them from ensialic “mobile belts”, involving major basement reactivation, and other basement regions showing a PanAfrican thermal overprint (Fig. 2). He also distinguished the magmatic arc terrain of the Arabian-Nubian Shield (Figs. 2 and 4) as a separate Pan-African province. We examine ophiolites and possible ophiolites from three different Pan-African terrains using examples from the transSaharan and Damara belts, and from the Arabian Shield. Trans-Saharan suture
Bou Azzel; Morocco Generally considered to be the first Proterozoic ophiolite identified, the Bou Azzer ophiolite (LeBlanc, 1975, 1976, 1981) is located in the Pan-African segment of the Anti-Atlas in Morocco (Fig. 2). This segment is identified by Burke et al. (1977) as the northwesterly extension of the trans-Saharan suture zone. According to the descriptions of LeBlanc (1981), the ophiolite appears to be complete, both in the sense of the Penrose definition and the expanded ophiolite association of Moores (1982), but it is tectonically dismembered into several thrust sheets obducted from northeast to southwest onto the northern margin of the Eburnean domain of the West African Craton. At the craton margin, approximately 2000 Ma Eburnean gneisses are unconformably overlain by 300 to 400 m thick quartzites
H.H. Helmstaedt and D.J: Scott
60
0>
2500 Ma
> 1000 < 2500
Ma
Pan-African mobile zones Zones with Pan-African Elthermal overprint Pan-African orogenic belts
L l l ophiolite remnantsPhanerozoic belts
Fig. 2. Distribution of areas affected by Late Proterozoic Pan-African event with location of ophiolites. Modified after Kroner (1979, 1980). K = Kalahari Craton, C = Congo Craton, WA = West Africa Craton, S = East-Sahara Craton, N = Nubian Shield, A = Arabian Shield. I = Pharusian belt, 2 = Mauritanide-Rokelide belt, 3 = Dahomey belt, 4 = West Congo belt, 5 = Katanga belt, 6 = Darnara belt, 7 = Gariep belt, 8 = Saldania belt, 9 = Zambezi belt, 10 = Mozambique belt. B = Bridgetown Formation, BA = Bou Azzer, ED = Eastern Desert of Egypt, G = Grootderm Formation, M = Matchless arnphibolite, N = Natal (part of Middle Proterozoic Narnaqua belt), T = Timetrine complex.
and stromatolitic limestones which are in thrust contact with the approximately 2 km thick Bleida Formation, a tectonically imbricated, approximately 790-Maold volcano-sedimentary continental-rift sequence (LeBlanc, 1981). The ophiolite is separated from the Bleida zone by a large thrust sheet of gneisses, either derived from the Eburnean craton, or representing a microcontinental fragment of unknown affinity (see Church, 1991; Hefferan et al., 1991). A major mylonite zone is developed along the contact between the gneisses and the ophiolite. A number of exposures over a strike length of more than 30 km were used by LeBlanc (1975, 1976, 1981) to reconstruct a 4-5 km thick ophiolite section
The Proterozoic ophioliteprobletn
61
interpreted to represent a coherent sample of Proterozoic oceanic lithosphere. This section comprises, from bottom to top (thicknesses are approximate, see also Fig. 1): (1) 2000 m of serpentinites derived from harzburgites and dunites with tectonite fabric. (2) 500 m of layered gabbros, grading from ultramafic cumulates upwards into hornblende microgabbros and quartz-microdiorites. (3) 500 m of mafic flows, parts of which are pillowed. (4) 1500 m of interlayered pillowed flows, greywackes, and tuffs with associated jaspillites and calcareous tuffs. Large sill-like bodies and stocks of quartz diorite are concentrated at the boundary between layered gabbros and mafic flows and interfinger with the gabbros. A well-developed sheeted dike complex has not been identified, but swarms of basic and keratophyric dikes, up to l m thick, cut the layered gabbros and appear to be feeders to the mafic flows. The dikes have chilled margins and are perpendicular to the igneous layering. Within the mafic flows, dikes are locally sheeted, occupying up to 70% of the total volume, and rodingitized mafic and microdioritic dikes are common within the serpentinites. Mafic volcanism along the Eburnean continental margin has been dated as approximately 790 Ma (based on a Rb-Sr whole-rock isochron of rocks in the metamorphic aureole around mafic dikes (LeBlanc, 1981). The major PanAfrican deformation of the Bou Azzer ophiolite complex was accompanied by greenschist-facies metamorphism dated as approximately 685 Ma (LeBlanc, 1981), and post-tectonic granodiorites have yielded a U-Pb zircon date of 615 Ma (Ducrot and Lancelot, 1978). The deformed ophiolite is unconformably overlain by conglomerates and feldspathic sandstones of the Tiddiline Formation, preserved in fault-bounded grabens, that in turn are overlain by Infracambrian and Paleozoic cover rocks. Small pods of chromitite within the serpentinized peridotites resemble podiform chromite deposits typical of ophiolites (LeBlanc, 1981). Copper mineralization, though not of economic concentrations, occurs near the base of the mafic flows, in breccias and mafic dike swarms below the pillowed flows, and in spilitic rocks at the top of the ophiolite. The well-known Bou Azzer cobalt ore bodies, located along the margins of the serpentinite massifs, postdate the emplacement of the ophiolite (LeBlanc, 1975; LeBlanc and Billaud, 1982; Buisson and LeBlanc, 1985). According to LeBlanc (1981), the Bou Azzer ophiolite probably evolved within a marginal sea that opened and closed in less than 100 Ma. The mafic flows plot within the ocean-floor tholeiite field on the Ti/Cr vs. Ni diagram of Beccaluva et al. (1979), but abundant diorite, trondhjemite and tonalite as well as low-Ti metabasalts (Bodinier, 1984; Naidoo et al., 1991) (Fig. 3a) also suggest characteristics of ophiolites formed within or near immature island arcs. On the basis of remapping, structural analysis, and geochemical studies in the Pan-African rocks of the Bou Azzer inlier, Saquaque et al. (1989) and Naidoo el al. (1991) suggest that the internal structure of the ophiolitic rocks is more complex
62
H.H. Helrnstaedt and D.J. Scott
100
Bou Azzer ophiolite
10
10
m
m LT 0
[r
0
i -p r
I -
1
a l 0.
i!
l
E
d
v)
.1
.01
.1
data from Bodinier et al., 1984 I
I
,
,
I
,
n-7 I
9
7
data from Al-Shanti and Gass, 1983
,018
7
I
I
I
I
I
I
I
I
I
I
n=21 I
I
I
I
I
3
1
Purtuniq ophiolite, MORB suite
Jormua ophiolite
10
m
m
LT
[r
0
0
50 f m
Fa
-
1
Q
l
l
Q
5
v)
data from Kontinen, 1987
.01
n=9
data from Scott, 1990
" " " ' " " ' " " ' Sr K Rb BaTh Ta NbCe P Zr HfSmTi Y Yb ScCr Ni
n=14
3
Sr K RbBaTh Ta NbCe P Zr HfSmTi Y YbScCr Ni
100
100
Purtuniq ophiolite, OIB suite
10
10
m
m [r 0
-. 0 a
LT
0
I .
H
-a l
1
l
Q
5
f
m
v)
.1
.1
data from Scott, 1990
n=5
data from Dann, 1991
n:20
71ze Proterozoic ophioliteproblem
63
than previously thought. These authors contend that the tectonized peridotites, gabbros, and sheeted dikes are part of a distinct terrane wedged between a fore-arc-basin terrane in the north, and an accretionary-melange terrane in the south. Pillowed metabasalts within the melange, viewed by LeBlanc (1981) as the extrusive parts of the ophiolite dislodged from the intrusive parts by thrust faults (see also Bodinier et a]., 1984), are reinterpreted as tectonic slivers unrelated to the ophiolite and separated from it by major transcurrent faults. As the tectonic significance of the melange zone south of the ophiolite is still a matter of debate (Church, 1991; Hefferan et al., 1991), it is obvious that further structural studies are necessary to verify whether the individual ophiolite components originated apart and were juxtaposed by transcurrent faulting, or whether they originated together and were dismembered by thrust faulting.
Pharusian belt The Pan-African suture marked by the Bou Azzer ophiolite continues southeastand southwards along the eastern margin of the West African Craton (Burke and Dewey, 1970; Caby et al., 1981, 1989 (with references therein)). It is outlined by elongated positive gravity anomalies which correspond to mafic and ultramafic rocks with petrological and geochemical similarities to ophiolites, but lacking a typical ophiolite pseudostratigraphy. The Timetrine complex in Mali (Fig. 2), is a typical example, consisting of four elongate massifs of serpentinized ultramafic rocks with lenses of metagabbro and diabase that are part of a nappe complex above thick epicontinental deposits, west of the suture. The ultramafic massifs are surrounded by chlorite-albite and sericite schists, which are structurally overlain by pillowed mafic flows with feeder dikes, but it is not clear whether the contacts of the ultramafic rocks are tectonic (LeBlanc, 1976) or intrusive (Caby et al., 1981). Nevertheless, general agreement exists that they are part of an oceanic domain telescoped between the highly deformed, passive- continental-margin sequence on the West African Craton, in the west, and the magmatic assemblage of the Tilemsi arc, in the east (Caby et al., 1989). Evidence for the operation of a Wilson Cycle during the evolution of this segment of the trans-Saharan Pan-African belt is compelling even without preservation of typical ophiolites. Rifting along the eastern margin of the West African Craton, leading to the opening of the “Pharusian Ocean” at about 800 Ma, is well established by the development of a passive-continental-margin sequence and alkaline magmatism associated with continental fragmentation (Caby et al., 1981). The Gourma trough, located at the embayment in the West African Craton south of the Timetrine complex (Fig. 2), is interpreted as an aulacogen that evolved in a failed arm at a high angle to the major rift. Ocean closure commenced with eastward subduction under the ensimatic Tilemsi arc a t about 730 Ma. This arc was then accreted against the eastern continent (western Hoggar region of the Saharan Craton) prior to collision between the West African and Saharan Cratons at about 635 Ma leading to extensive reworking of the pre-Pan-African basement of the latter (Caby et al., 1989). The suture between the West African Craton and
64
H.H. Helrnstaedt and D.J. Scott
the Tilemsi arc is marked by a zone of high-pressure metamorphism, resulting in the development blue amphiboles in the oceanic metabasalts at Timetrine, and of white schists and eclogites near Gourma, to the south (see also Caby, 1987).
Dahomey belt Although complete ophiolites have not been recognized, greenstones interbedded with marine sedimentary rocks in the Dahomey belt (Fig. 2) suggest that the “Pharusian Ocean” extended southwards into this region (Burke and Dewey, 1973). However, according to Kroner (1979), the field evidence is consistent only with the existence of a “miniocean”, because the continental rises of the opposing sides remained close enough to supply turbidites across the entire ocean floor.
Damara belt, soulhem Aj?ica Geotectonic interpretations of the Damara belt in southwestern Africa (Fig. 2) provide an interesting contrast to those of the trans-Saharan belt, because the lack of well-preserved ophiolites led to a continuing lively controversy over the importance of plate-tectonic processes in Precambrian orogeny as well as over criteria for the recognition of ocean-floor spreading and subduction processes in ancient rocks (see Kroner, 1983). Contending that convincing evidence for plate separation and formation of oceanic crust cannot be found, Martin and Porada (1977) and Kroner (1977, 1983) have considered the NE-SW trending branch of the Damara orogen as a type example of ensialic orogeny, involving the closure of wide intracratonic basins without subduction of oceanic crust, but through interstacking of continental crust and possible delamination of subcrustal lithosphere (see also Kroner, 1981). Others have argued for platetectonic models, beginning with rifting and seafloor spreading and followed by northward subduction of the Damaran oceanic crust and continental collision during which the northwestern edge of the Kalahari Craton was overridden by the Congo Craton (e.g., Hartnady, 1979; Barnes and Sawyer, 1980; Kasch, 1983, 1986). A clearly defined suture has not been identified in the southern Damara orogen, but several occurrences of serpentinite, talc schist, metagabbro and metabasalt, closely associated with S- to SE-vergent thrust faults within the Southern Zone and the Southern Margin Zone, suggest the existence of a complex collision zone (Kasch, 1983, 1986; Hartnady et al., 1985). The ultramafic rocks have a depletedmantle composition (Barnes, 1982) and are therefore considered to represent fragments of oceanic lithosphere (Kasch, 1983). The chemical composition of the metamorphosed tholeiitic volcanic and intrusive sequence of the Matchless Amphibolite (up to 3000 m thick), forming a remarkably straight zone along the southern margin of the Khomas trough (Fig. 2), changes from that of continental flood basalt in the west to ocean-ridge basalt in the east (Breitkopf and Maiden, 1987). As this change is comparable with the north to south compositional variations of the tholeiitic basalts along the central graben of the Red Sea, the formation of the Matchless belt is thought to be consistent with deposition in
The Proterozoic ophioliteproblem
65
an environment changing from advanced continental rifting to the formation of small ocean basins (Breitkopf and Maiden, 1987). More work is needed before the existence of a major Damaran ocean basin (the “Adamastor Ocean” of Hartnady et a]., 1985) can be verified, thereby establishing whether models involving the operation of a Wilson Cycle provide a more plausible explanation for the evolution of this orogen than “ensialic” orogenic models such as that proposed by Kroner (1983). Hartnady et al. (1985) also proposed that the ‘Adamastor Ocean” extended southwards through the Gariep and Saldania belts of Namibia and South Africa (Fig. 2). In the Gariep belt, mafic metavolcanic rocks of the Grootderm Formation and an associated melange with blocks of serpentinite may represent remnants of an oceanic domain, but typical ophiolites have not been discovered (see also Kroner, 1979; Hartnady et al., 1990). The metavolcanic rocks have the geochemical signature typical of hotspot-related oceanic islands (Smith and Hartnady, 1984; Hartnady et al., 1985). Metabasalts of the Bridgetown Formation in the Saldania belt further south (Fig. 2) may have originated in a similar tectonic setting.
Arabian Shield The Arabian-Nubian Shield (Figs. 2 and 4) consists of a series of Late Proterozoic intra-oceanic island arcs accreted with various microplates and slivers of ophiolites between about 950 and 600 Ma and cratonized by post-orogenic granitoid intrusions, uplift and cooling about 570-500 Ma ago (e.g., Greenwood e t al., 1976; Al-Shanti and Gass, 1983; Pallister et al., 1988; Dixon and Golombek, 1988; Berhe, 1990). Although generally considered to be a northward extension of the Pan-African Mozambique belt (e.g., Kennedy, 1964; Coleman, 1984), Kroner (1980) distinguished this mainly juvenile island-arc terrain as a separate province (Fig. 2), noting that its rock types and tectonic relationships are unlike those in the other Pan-African belts, especially those in the generally ensialic “mobile belts” such as the Mozambique belt. Some of the microplates within the arc terrain, however, may be rifted fragments of the Early Proterozoic Mozambique belt (Stoeser and Camp, 1985). The contrast between the Arabian-Nubian Shield and the Pan-African “orogenic belts” is also evident in the distribution of ophiolitic rocks. Ophiolites do not occur in megasutures between major cratons but, as first pointed out by Brown and Coleman (1972), occupy cryptic sutures that mark the sites of collisions between microplates or arc systems (Fig. 4). This interpretation has been generally confirmed by geochronological studies on the Arabian Shield (summarized in Pallister et al., 19SS), but caution is warranted, as many ophiolites are allochthonous and may not accurately mark the sutures between individual arc systems (Gass, 1981). Shackleton et al. (1980), for instance, believe that the numerous maficultramafic masses in the Eastern Desert of Egypt (Fig. 2) have moved too far to justify the identification of linear ophiolite belts as sutures (see also Church, 1988).
H.H. Helmstaedt and D.J. Scott
66
Yanbu Sutur
0
100
200
300
0
Phanerozoic cover rocks
Arc terranes
c'.y Sub-terrane
with continental affinity
Fig. 4. Distribution of ophiolites in Arabian Shield. Modified after Pallister e t al. (1988) and Stoeser and Camp (1985). AA = A1 Amar, AD = Ad Dafinah, AR = Ar Ridaniyah, AU = A1 Uwayja, BT = Bir Tuluhah, BU = Bir Umq, DZ = Darb Zubaydah, H = Hamdah, Ha = Halaban, JE = Jabal Ess, JT = Jabal Tays, JW = Jabal a1 Wask, T = Tathlith, Th = T h u w a h , Uus = Umm ash SharaL.
Whereas some of the mafic-ultramafic complexes on the Arabian-Nubian Shield conform to the Penrose definition, they lack certain units of the expanded ophiolite association (Moores, 1982); in particular, crystalline basement rocks with shelf and continental rise deposits typically found at the structural base of Tethyan ophiolites are missing. Instead, the Arabian ophiolites are juxtaposed against
The Proterozoic ophioliteproblem
67
island arc rocks or volcanogenic sedimentary rocks reminiscent of Cordillerantype ophiolites. Their outcrop patterns and associated rock types show surprising similarities not only to Archean and Early Proterozoic granite-greenstone terrains (as noted by Engel et al., 1980), but also to Phanerozoic accreted terranes, such as those of the western North American Cordillera and Alaska (as proposed by Pallister et al., 1988). As age relationships have been studied most thoroughly in the Arabian Shield (see Pallister et al., 1988), our discussion is confined to the ophiolites of this region. This shield has been subdivided into five terranes, all of which are bounded by complexly deformed belts with slivers of ophiolites interpreted as suture zones (Stoeser and Camp, 1985) (Fig. 4). Three juvenile arc terranes in the western part (Asir, Hijaz, and Midyan), divided by the NE-striking Yanbu and Bir Umq sutures, are bounded in the east by the northerly-trending Nabitah suture zone against the Mif terrane, a composite arc terrane possibly including an embedded pre-Pan-African microplate (Pallister et al., 1990). The Mf terrane is separated from the Ar Rayn terrane in the easternmost part of the shield by another NNW-striking, composite suture zone (Urd and Al Amar sutures). Brief descriptions of these suture zones will give an impression of the variety of ophiolitic rocks and their geological settings. Ynnbu suture This northeasterly-trending belt, offset along a sinistral, northwesterly-trending strike-slip system, contains the Jabal Ess and Jabal a1 Wask ophiolites, respectively among the best preserved and largest ophiolite complexes described from the Arabian Shield (Pallister et al., 1988) (Fig. 4). Both are in tectonic contact with deformed and metamorphosed arc volcanic and sedimentary rocks (lower and middle formations of the Farri Group) and are unconformably overlain by similar volcanic and sedimentary rocks (Al Ays Group) with a local basal conglomerate containing ophiolite debris (Stoeser and Camp, 1985; see also Pallister et al., 1987).
Jubul Ess ophiolite. The Jabal Ess ophiolite in the northwestward displaced segment of the Yanbu suture, contains all the units of the Penrose definition (Fig. l ) , but as many contacts are faulted or intruded, a coherent internal sequence cannot be demonstrated, and the 3000 m thickness estimated for the exposed section by Shanti and Roobol (1979) must be regarded as a minimum. The base of the complex is an up to 250 m thick melange that contains blocks of all units of the ophiolite in a serpentinite matrix and is interpreted as an original low-angle thrust along which the ophiolite was emplaced over slightly metamorphosed shales. The complex was folded twice after emplacement, and the uppermost units, consisting of pelagic sediments, pillowed mafic flows (upper Farri Group), and sheeted dikes, crop out in a n approximately 3 km wide, northwesterly-trending, steeply dipping belt representing the core of a faulted and refolded synform (Shanti and Roobol, 1979). The folded basalt flows (300 m thick)
68
H.H. Helmstaedt and D.J. Scott
are flanked on both sides by sheeted dikes, gabbro, and serpentinized harzburgite and dunite, which locally contains podiform chromitite. The sheeted-dike complex is 200-600 m thick and has gradational boundaries with the flows above and gabbro below. It consists almost entirely of meta-dolerite dikes (30 cm-2 m wide) which, immediately below the steeply dipping basalt flows, have chilled margins that are perpendicular to the flow boundaries. Layered and isotropic gabbro occurs as a thin septum (
7he Proterozoic oplziolite problem
69
magmatic arcs were built (Pallister et al., 1988). Camp (1984) suggested that the Yanbu suture and its southwestward continuation into northeastern Sudan mark the former site of a southeast-dipping subduction zone.
Bir Umq suture This northeast-trending belt (Fig. 4) of deformed arc-related igneous and sedimentary rocks (Samran Group), containing imbricate thrusts faults, overturned folds, and allochthonous sheets of marble, basalt and chert (Birak Group), is interpreted as the former site of a southeast-dipping subduction zone above which the Bi f arc system in the northern part of the Asir terrane developed (Stoeser and Camp, 1985). The Thunvah ophiolite (Nassief et al., 1984) (Fig. 4) consists mainly of serpentinized peridotite tectonites and cumulates with tectonically interleaved zones of layered and massive gabbro. Although most of the finer-grained hypabyssal rocks are deformed and recrystallized to the extent that dikes and flows cannot be distinguished, local transitions from massive gabbro and pillowed flows into sheeted dikes were noted by Nassief (as quoted by Pallister et al., 1988). The Bir Umq complex, near the intersection of the Bir Umq and Nabitah sutures (Fig. 4), is an east- to northeast-trending thrust sheet of serpentinized peridotites with associated hornblende gabbro, diorite, and pillow basalt interlayered with cherts and tuffs. According to Pallister et al. (1987, 1988), sheeted dikes were not well-preserved, and a coherent ophiolite pseudostratigraphy has not been demonstrated. From zircon ages of 820-870 Ma and 838 Ma for the Thunvah and Bir Umq complexes, respectively, Pallister et al. (1988) concludes that the ophiolitic rocks represent oceanic crust that was formed less than 20 Ma prior to early arc magmatism in the region. An older zircon component in the Thunvah gabbro (> 1250 Ma), however, suggests incorporation of older continental crust material. Nabitah suture zone The Nabitah orogenic belt is a complex transpressive suture zone (Quick, 1991) separating the western composite arc terrane from the eastern Afif terrane, also a composite arc terrane with possible pre-Pan African basement in its southeastern part (Pallister et al., 1990) (Fig. 4). It is named after the serpentinite-lined Nabitah fault zone in the southern part of the belt, east of which are two mafic-ultramafic bodies near Bthlith and Hamdah. In the central part of the Nabitah belt, the zone of serpentinites is displaced northwestward along the left-lateral Najd strike-slip fault system. Larger mafic-ultramafic masses occur near Bir Tuluhah, at Bir Nifazi (Darb Zubaydah ophiolite) and within the Ad Dafinah belt (Fig. 4). The Bir Tuluhah complex is a fault-bounded, steeply dipping slab of ultramafic rocks showing an east-west progression from metabasites through peridotite tectonite, cumulus peridotite, cumulus gabbro to non-cumulus gabbro (Pallister et al., 1988). The contact relationships to pillow basalts west and east of this slab are not clear, but similar dikes of gabbro, diorite, and plagiogranite have been found
70
H.H. Helmstaedt and D.J. Scott
within the mafic-ultramafic rocks and the basalts, suggesting that the rocks belong to an ophiolitic association. The Darb Zubaydah ophiolite (Quick, 1990), is approximately 830 Ma old and preserves a largely intact section consisting of ultramafic rocks, gabbro, granodiorite, and interbedded volcanic and sedimentary rocks that is exposed in an east-dipping blow. The sequence is extensively intruded by diabase dikes, but a classic sheeted-dike complex is not developed. Quick (1990) interprets the 9-13 km thick ophiolitic section in terms of continuous crustal evolution within or near an island arc. Compositions of the oldest extrusive unit range from tholeiites with MORB characteristics to calc-alkaline andesites and rhyolite, suggesting that the rocks formed in a relatively unevolved island arc or in a back-arc basin sufficiently close to an arc to receive calc-alkaline lava flows and coarse-grained, arc-derived detritus. No complete ophiolitic sequence is identified in the discontinuous maficultramafic bodies of the Ad Dafinah belt and the Nabitah fault zone (Fig. 4). A relatively large lens east of the fault zone, near Tmthlith, containing serpentinite, metamorphosed microdiorite, diabase, and basalt without a coherent ophiolite structure, is intruded by a gabbro which postdates the possible ophiolitic rocks (Pallister et al., 1988). At Hamdah, flat, sheet-like serpentinite bodies, overlying hornblende and biotite schists with a possible thrust contact, extend as much as 40 km east of the Nabitah fault zone (Pallister et al., 1988). A structurally complex unit above the serpentinite contains foliated gabbro and diorite, metabasalt, diabase, metaandesite, and metagreywacke and may represent an ophiolitic melange. Whereas the age of the Tuluhah complex, in the northern part of the Nabitah belt, is about 847-823 Ma, that of the Hamdah complex appears to be about 100 Ma younger (Pallister et al., 1988). As the oldest arc rocks also range from about 840 Ma in the north to about 730 Ma, the Nabitah belt is probably not a suture resulting from a single subduction event. Although considering the possibility that the Bir Tuluhah suture could be correlative with an older suture identified west of the Nabitah fault zone in the south, Pallister et al. (1988) suggested that age relationships of ophiolitic and arc rocks in the northern part of the Nabitah belt are more consistent with the interpretation that the Bir Tuluhah belt is a continuation of the Bir Umq suture that was transposed to the northwest along the Nabitah fault zone.
Urd and A1 Amar sutures The Al Amar suture of Stoeser and Camp (1985), separating the Afif terrane from the Ar Rayn terrane at the eastern edge of the Arabian Shield (Fig. 4), is marked by a deformed accretionary complex (Abt Formation) that is bounded on each side by a belt of ophiolitic rocks. The western belt, defined as the Urd suture by Pallister et al. (1988), is a 1-10 km wide zone, best exposed at Halaban, containing lenses of metagabbro associated with serpentinized peridotite, talc-anthophyllite schist and listwaenite. The eastern belt, occurring along the Al Amar-Idsas fault and referred to as Al Amar-Idsas suture by Pallister et al.
TIze Proterozoic oplziolite problem
71
(1988), appears to merge in the south with the western belt, possibly as the result of a broad synclinorium structure. The eastern belt is described by Al-Shanti and Gass (1983) as a melange with blocks of oceanic crust in a matrix of sheared serpentinite. The melange occurs as a sheet-like layer underlying the schists of the Abt Formation and is considered to be part of an accretionary complex. Blocks within the melange are generally 1-2 m in diameter and are rarely larger than 100 m. Most blocks are basalts and dolerites, but blocks of chromitite and plagiogranite have also been identified. Two large blocks with well-defined sheeted-dike structure have been identified south of Jabal Idsas. At Jabal Thys, west of the Al Amar-Idsas fault, much larger blocks are found in the melange, and a number of prominent hills are centered on single gabbro blocks. Many gabbro blocks show igneous lamination and cyclic layering from melanocratic, olivine and pyroxene-rich gabbros to anorthosites. The spindle-shaped chromite grains in the serpentinized peridotite are indicative of the high-temperature deformation textures seen in tectonized peridotites of many ophiolites. Judging from the slightly elevated content in large-ion-lithophile elements and somewhat depleted high-field-strength elements with respect to MORB (Al-Shanti and Gass, 1983), the mafic rocks of the Al Amar suture resemble island-arc tholeiites (Fig. 3b). However, the tectonic significance of the melange, including the type of ocean floor subducted and the polarity of subduction, are still a matter of debate (see Pallister e t al., 1988). EARLY PROTEROZOIC OPHIOLITES
Baltic Shield Based on facies relationships and structures in the Karelides (Fig. 5), the western boundary zone of the Archean Presvecokarelian (Karelian) craton of eastern Finland may represent an Early Proterozoic suture, along which the Proterozoic Svecofennian domain of western Finland and central and northern Sweden was accreted to the Archean craton as a result of the Svecokarelian orogeny (e.g., Burke et al., 1977; Simonen, 1980; Koistinen, 1981; Park et al., 1984; Park, 1988, 1991). Park et al. (1984) summarized the tectonic evolution of this Early Proterozoic margin as follows: (1) Stabilization of Archean Presvecokarelian craton by 2500 Ma. (2) Early Proterozoic craton break-up beginning with the intrusion of layered mafic-ultramafic masses in northern Finland at 2450 Ma. Deposition of the shallow-marine Jatulian sequence and intrusion of basic sills and dikes with associated volcanism between 2300 and 2050 Ma. (3) Deposition of Outokumpu assemblage (consisting of carbonaceous metapelites, dolostones and silica-rich exhalites with tholeiites and komatiites) in a N-S shallow marine basin over stretched continental crust. Rifting is suggested by intrusions of gabbro and sheets of ultramafic rocks at about 1970 Ma.
72
H.H. Helmstaedt and D.J. Scott
Fig. 5. Proterozoic orogenic belts and terranes of Laurentia and Baltic Shield (in pre-Iapetus position). Modified after Hoffman (1988,1989). 1 = Jormua complex in Svecokarelian suture, 2 = Purtuniq ophiolite in Cape Smith belt (CS).3= Payson ophiolite in Mazatzal block of central Arizona. BL = Belcher belt, CH = Cheyenne belt, CS = Cape Smith belt, GS = Great Slave Lake shear zone, FR = Fox River belt, K R = Keweenawan rift zone, MK = Makkovik orogen, TH = Thompson belt.
(4) Deposition of turbidites (Kalevian) about 1950-1900 Ma. (5) Emplacement of thrust nappes (Outokumpu nappe) onto the basal facies of the Kalevian during convergence of Archean Iisalmi and Kuhmo terranes (Park, 1991) at about 1900 Ma. The Outokumpu assemblage, containing serpentinite, pillow basalt, and cherts was earlier interpreted as ophiolitic in the sense of the Steinmann Trinity (Wegmann, 1928), and this interpretation has played an important role in plate tectonic models for the Karelides. Although the absence of sheeted dikes and the internal stratigraphy of the Outokumpu assemblage in the Outokumpu region are generally not consistent with an ophiolite section in the sense of the Penrose
71ze Proterozoic ophioliteproblem
73
definition (Park et al., 1988), a complete ophiolite section was described from the Jormua complex, the northernmost of a chain of mafic-ultramafic bodies extending along the Karelides and thought to be related to the Outokumpu association (Kontinen, 1987). It is thus possible that the Outokumpu basin was a t least locally floored by oceanic crust. The Jormua conzplex As described by Kontinen (1987), the Jormua complex comprises an intensely faulted body of serpentinites and associated mafic intrusive and extrusive rocks about 2-5 km in width and 20 km in length. Five distinct lithological units have been recognized that, in spite of tectonic disruption and lower-amphibolite-facies metamorphism, can be reconstructed into their pre-deformation configuration (Fig. 1): (1) A basal metaserpentinite with cross-cutting metabasite dikes and metagabbro bodies that invariably exhibits tectonic contacts with the country rocks consisting of Presvecokarelian gneisses and granitoid rocks (> 1000 m). (2) Bodies of metagabbro locally dissected by metabasite dikes (>lo0 m). ( 3 ) A dike complex consisting of narrow subparallel metabasalt and metadolerite dikes that are locally sheeted but otherwise have numerous interdike screens of metagabbro and serpentinite (300 m). ( 5 ) A thin sequence of metacarbonate rocks, metachert, black schists and calcareous metatufites conformably overlying the metabasalts (<200 m) and overlain by metaturbidites of the Upper Kalevian (>>500m), presumably also with conformable contact. Metaserpentinite, the predominant rock type of the complex, consists almost entirely of antigorite with accessory chromian magnetite and rare relics of meshtextured lizardite showing probable olivine pseudomorphs. Relics of primary layering or tectonite fabrics are not recognized. From the observation that mafic dikes cutting the serpentinite are not rodingitized, Kontinen (1987) inferred that they did not intrude fresh peridotite. The formation of the Jormua complex may have begun with protrusions of serpentinized mantle into early rifts that localized the spreading complex indicated by the sheeted dikes. With a total outcrop area of less than 1 km2, gabbro is a relatively minor component of the Jormua complex. Although its contacts are generally faulted, magnesian metagabbro cuts the serpentinites as small stocks and dikes, and ilmenite-rich meta-ferrogabbro occurs as septae between dikes of the dike unit. Both gabbro types are characterized by irregular grain-size variations and the general absence of igneous layering. Metatrondhjemites occur as irregular segregations within the gabbros or as cross-cutting dikes within the mafic dike unit. Outcrops of the thick and locally spectacular dike complex occupy several square kilometres and, where not faulted, the contacts with gabbros are gradational. As the contacts of the metabasalts are faulted, a transition of the dikes into this unit cannot be documented, but at least one outcrop of dikes with
74
H.H. Helmstaedt and D.J. Scott
pillow basalt screens is described by Kontinen (1987). The width of individual dikes ranges from a few millimetres to 6 m, but is generally between 20 cm and 120 cm. An earlier generation of generally thicker metadolerites and plagioclase-phyric dikes can be distinguished from a younger generation of thinner and finer-grained, aphyric metabasalt dikes. The younger dikes generally split the older dikes resulting in numerous half dikes or marginless septa of the older dike genera tion. Extrusive rocks are volumetrically minor with respect to the dike complex and consist of an approximately 300 m thick, faulted unit of mainly metabasaltic, pillowed flows with minor massive flows and breccia. The scarcity of dikes in these flows suggests that they were never in close proximity with the dike complex, or that most of the dikes of the intrusive complex are older than this flow unit. Within error limits of about 11 Ma, zircon dates of 1960 and 1954 Ma for a coarse-grained metagabbro and a metatrondhjemite, respectively, are virtually identical and suggest that the Jormua complex is coeval with other maficultramafic bodies of the Outokumpu assemblage (Kontinen, 1987; Park, 1988). The Jormua complex lies also in the same tectonostratigraphic position as the Outokumpu assemblage, above a major east-directed thrust fault on the parautochthonous or basement rocks of the eastern Savo province (Park, 1988), and overlain by Kalevian flysch. However, the Jormua gabbro and basalts have MORB affinity (Huhma, 1986; Kontinen, 1987) (Fig. 3c), whereas the Outokumpu assemblage was derived from a more depleted source, suggesting intra-arc or backarc affinities (Park, 1988; 1991). This suggests that the Jormua nappe includes a fragment of the ocean floor of the Outokumpu basin. Whether this basin was a narrow rift, as proposed by Kontinen (1987), or whether it was a much wider ocean, depends on the interpretation of the Iisalmi terrane (Park, 1991), west of the Jormua nappe. Although this was earlier viewed as a piece of the Archean Karelian craton, the rifting and rejoining of which, respectively, created and closed the Outokumpu basin (Park et al., 1984; Kontinen, 1987; Park, 1988), its Archean and Jatulian evolution differ sufficiently from that of the craton to the east (Kuhmo terrane) suggesting that it may represent an exotic terrane (Vayrynen, 1939; Park, 1991). Typical of the preservation of ophiolites elsewhere, the Jormua complex is situated along the relatively minor Savo suture between the Iisalmi and Kuhmo terranes (Kontinen, 1987; Park, 1991), and no ophiolites are known from the major Svecokarelian suture between the Karelian collage, including the Iisalmi terrane, and the Svecofennides to the west (for schematic cross-sections illustrating this relationship, see fig. 4 of Park, 1991). Canadian Shield The Canadian Shield exposes a major part of Laurentia, the Precambrian nucleus of North America, the Early Proterozoic assembly and growth of which was summarized in two papers by Hoffman (1988, 1989). As pointed out by Hoffman (1988), Laurentia owes its existence to a network of Early Proterozoic
R e Proterozoic oplzioliteproblem
75
orogenic belts, and its assembly in the Early Proterozoic is comparable to the Late Proterozoic (Pan-African) and Phanerozoic accretion, respectively, of Africa and Eurasia. Arguments over the nature of the Early Proterozoic belts in North America were similar to those in the Pan-African belts, because many are cryptic sutures preserving only the deformed margins of adjacent Archean microcontinents. However, accreted Early Proterozoic island arcs and associated oceanic deposits occur in some belts (Schulz, 1987) (Fig. 5 ) and, though complete ophiolites have not been described, a variety of plate tectonic models have been proposed (see Hoffman, 1988). Of special interest is the Trans-Hudson orogen in Manitoba and Saskatchewan (Fig. 5 ) , with an internal zone of intra-oceanic rocks (Stauffer, 1984) between the ensialic Thompson belt, along the margin of the Archean Superior Province in the east, and the Cordilleran-type Wathaman-Chipewyan batholith (Lewry et a]., 1981) in the north and west. Several domains of this internal zone were considered to be typical Archean granite-greenstone terrains, until isotopic dating and geochemical studies (see Sangster, 1972, 1978; Stauffer, 1984; Van Schmus et al., 1987; Chauvel et al., 1987; Gaskarth and Parslow, 1987; Watters and Pearce, 1987) established that the 1910-1880 Ma volcanic rocks and associated 1890-1840 Ma plutonic rocks are island arc rocks that formed in an Early Proterozoic oceanic basin, the “Manikewan Ocean” of Stauffer (1984). Submarine mafic volcanic rocks of the Amisk Group in the Flin Flon domain show a continuous gradation from tholeiitic to calc-alkaline compositions and have been interpreted as representing the transition from the lower to upper parts of mature island arcs (see Stauffer, 1984). Other mafic volcanic rocks within the boundary zone between the 3ansHudson orogen and the Superior Province, including those in the Fox River belt (Fig. 5), have chemical compositions that are consistent with a marginal-basin origin (Halden, 1991). In what is interpreted as the pinched extension of the Trans-Hudson orogen, in northern Quebec, rocks of the internal zone, including the crustal portion of an ophiolite, were obducted together with a continental margin sequence onto the Superior Craton and are preserved as the infolded Mippe of the Cape Smith fold and thrust belt (Hoffman, 1985, 1988; St-Onge et al., 1987, 1988; St-Onge and Lucas, 1990a). The Cape Smith belt The Purtuniq ophiolite occurs in a large thrust sheet along the northern margin of the Cape Smith belt, an Early Proterozoic fold and thrust belt located on the northern margin of the Archean Superior structural province, astride the tip of the Ungava Peninsula of northern Quebec. The debate over the geotectonic significance of this belt is typical of that for other orogens, as it was interpreted successively as marking the site of: (1) a major collisional suture zone (Gibb and Walcott, 1971; Dewey and Burke, 1973; Burke and Dewey, 1977); (2) an ensialic fold belt (Baer, 1978); (3) a narrow rift zone (Baragar and Scoates, 1981); and (4) a suture after closure of a small, short-lived oceanic basin (Hynes and Francis,
16
H.H. Helmstaedt and D.J. Scott
1982). As basement of the Superior Province is continuous from south to north around the eastern margin of the belt (see also Doig, 1983), Hoffman (1985) suggested that the entire belt is a tectonic klippe that is isolated from its root zone to the north (Sugluk suture zone) by a basement antiform. Based on the work of Baragar and Scoates (1981), Moores (1986) cited the Circum-Superior belt as an example of a Proterozoic belt in which thick maficultramafic sequences resembling the upper parts of ophiolites are present in the position occupied by ophiolites in Phanerozoic suture belts. Noting that the volcanic sequences are described as grading from platformal and rift-related at the bottom to oceanic at the top, he also predicted that major thrust faults should exist along which the oceanic rocks are juxtaposed over the initial rift sequences. As a result of a three-year mapping program in the eastern part of the Cape Smith belt, St-Onge et al. (1988) distinguished the following five tectonostratigraphic units (from south to north): (1) A fluvio-deltaic sedimentary sequence that unconformably overlies gneisses of the Superior Province. (2) Sedimentary rocks and tholeiitic volcanic rocks of the Povungnituk Group (Bergeron, 1959) interpreted by Hynes and Francis (1982) to have been deposited in pre-1960 Ma, north-facing continental rift margin. ( 3 ) Mainly pillowed mafic volcanic rocks (ca. 1920 Ma) of the Chukotat Group (Bergeron, 1959) ranging in composition from Mg-rich komatiitic basalts to mid-ocean ridge basalts thought to indicate a change from continental rift-type volcanism to the formation of transitional oceanic crust (Hynes and Francis, 1982; Francis et al., 1981, 1983). (4) Deep-water sedimentary rocks of the Spartan Group (Lamothe e t al., 1984). ( 5 ) Older mafic and ultramafic volcanic and intrusive rocks (ca. 1998 Ma) of the Watts Group (Lamothe et al., 1984), now interpreted as oceanic crust and referred to as Purtuniq ophiolite by St-Onge et al., 1988). Structural work by St-Onge et al. (1987, 1988) and St-Onge and Lucas (1989, 1990a) confirmed Hoffman’s (1985) prediction that the Cape Smith belt represents an infolded tectonic klippe, because units 2 to 5 were tectonically imbricated during transport from north to south. Whereas the Povungnituk and Chukotat groups (units 2 and 3) may be considered par-autochthonous with respect to the basement of the Superior Province, the Spartan and Watts groups (units 4 and 5 ) are clearly exotic. Interestingly, although the amphibolite-facies, mafic volcanic and plutonic assemblage of the Watts Group (Begin, 1989), preserved in the northernmost and structurally highest thrust stack, was previously interpreted as the metamorphic equivalent of the Povungnituk Group, the thrust fault at its base has been recognized as one of the more profound structural features in the area, for it can be traced for most of the length of the Cape Smith belt (Bergeron, 1959; Baragar and Scoates, 1987). North of the Cape Smith belt, the Archean basement gneisses of the Superior Province are folded into an antiform which plunges to the west beneath allochthonous magmatic arc rocks, yielding U-Pb zircon ages of 1863-1830 (Parrish,
l h e Proterozoic ophioliteproblem
77
1989). This arc terrane is wedged against the northern margin of the Cape Smith belt (St-Onge and Lucas, 1990b), and its accretion is thought to have caused the intensive reimbrication observed in the thrust stacks of the earlier emplaced Cape Smith klippe (Lucas and St-Onge, 1991). The Purtuniq ophiolite At its type locality, near Lac Watts (Fig. 6), the Watts Group consists mainly of metamorphosed basalts with gabbro dikes and sills as well as layered gabbros and ultramafic cumulates that are cut by numerous internal thrust faults and are bounded below by a major thrust fault against graphitic pelites and semipelites of the Spartan Group (see also Fig. 7). Detailed mapping by St-Onge et al. (1987, 1988) and St-Onge and Lucas (1989) show that a 3 km thick sequence of pillow basalts with gabbro sills and mafic dikes, forming the lowest thrust sheet, is overlain by successive thrust slices of layered gabbro and layered peridotite. After preliminary structural sections suggested that the Watts Group may represent the structurally dismembered, crustal portion of an ophiolite, a search near the base of the mafic volcanic sequence, in the thickest part of the lowest thrust sheet, east of Lac Watts (Fig. 6), led to the discovery of a sheeted-dike complex (St-Onge e t al., 1988; see also St-Onge and Lucas, 1990a). Careful restoration of the faulted and folded sequence, based on cross-sections constructed by utilizing the structural relief provided by a major cross-fold east of Lac Watts (Figs. 6 and 7) (Scott, 1990, using plunge data of St-Onge and Lucas, 1989, and Lucas, 1989), yielded an approximately 9 km thick, composite section showing all the rock types, primary structures and pseudostratigraphy of the crustal portion of an ophiolite (Fig. 1). Structural correlation of units from separate thrust sheets has been confirmed by whole-rock geochemical (Scott, 1990; Scott et al., 1991; see also below) and isotopic studies (Hegner and Bevier, 1991). Pb-Pb whole-rock dating of sheeted dikes and cumulate rocks from different thrust sheets yielded identical isochron ages of 1970 f 4 and 1970 f 11 Ma, respectively, confirming an essentially contemporaneous origin, and combined data for layered cumulates, sheeted dikes, and pillowed basalts constrain the age of the ophiolite to 1980 f 3 Ma (Hegner and Bevier, 1991). The latter age is in good agreement with a U-Pb zircon age of 1998 f 2 Ma determined for a metagabbro from the layered complex (Parrish, 1989). The transition zone between the mafic dike complex and pillowed basalt, as mapped by Scott and Bickle (1991) (Fig. S), shows sheeted dikes with zones of massive amphibolite grading through alternating fine-grained basalt with local pillow outlines, amphibolites, and up to 50 m wide bands of sheeted dikes, into pillowed flows with some mafic dikes and local plagiogranite. Individual dikes range in width from less than 20 cm up to 50-60 cm (Fig. 9) and are oriented roughly perpendicular to the overlying flows. Typically sheeted dikes, many of them preserving only one chilled margin, are locally intruded by finer-grained, chloritic dikes cutting the sheeted dikes at low angles. Although not easily distinguishable in outcrop, two chemically distinct groups of sheeted dikes are present (Scott, 1990). The majority of dikes are similar to modern MORB, whereas the less
78
............. .............
......... ......... ....... ....... ......... .......... .......... .......... ............. ............. ............. ............. ............. ............... ................ ........... ........... ........... .............. ............. ............ ............ ......... .......
H.H. Helmstaedt and D.1 Scott
Fig. 6 . Simplified geological map of Lac Watts area, eastern portion of Watts Group, Cape Smith belt, northern Quebec (location of area is outlined on map in inset showing northern part of Ungava Peninsula, northern Quebec) (after St-Onge and Lucas, 1989). Line segments&’, BB’, CC’,and DD’ are lines of cross-sections in Fig. 7. Circled letters identify individual thrust sheets and are keyed to cross-sections in Fig. 7. Location of sheeted dikes, east of Lac Watts, is indicated by arrow in central part of the map.
Natts Group Composite Crustal Section
B' 10
-D1 Uwst fault L
9
P P
8
C'
7 6
5 4
3 2 1 1 krn
2 1 0 krn
mafic cumulate rocks ukramafic cumulate rocks
L
mafic cumulate rocks
iP
L
Fig. 7. Vertical cross-sections along lines indicated on Fig. 6 (after Scott, 1990). Legend same as Fig. 6. Inset on right of figure shows reconstructed composite section at the same scale as cross-sections. Locations of individual segments (numbered) used to construct th composite section are shown in cross-sections.
H.H. Helmstaedt and D.J. Scott
80
Watts Group, Purtuniq Ophiolite
******** *
-
piagiogranite 0 9 9 0
*
0 0 0 0 0
75 *
0 0 0 9
*** ** ** . 4 , * * epidote banded fine-grained basalt, with pillows ’ : :1: fi 2 :; : pillowed basalts, epidote selvages
.)
***
9 9 0 9 9 9
9 9
0
0 0 0 0 0 9 ~ 0 9 9 9 0 0 0
*09 Q%* < *0 0 %>* 0 * 0 ** * * * * * 9 * * * * 0 * * * * massive medium- to coarse-grained arnphibolite
sheeted mafic dikes
0 0 0
0 9
0 9
0 0 0 9 0
44 J strike and dip of dike 5 6 j strike and dip of schistosity Y
0
retrograde shear zone
200 m
100 .....
Fig. 8. Detailed geological map of sheeted-dike locality (for location see Fig. 6) showing transition zone between sheeted mafic dikes and dominantly pillowed mafic volcanic rocks at base of thrust sheet E, east of Lac Watts (see also section AA’ in Fig. 7) (modified after Scott and Bickle, 1991).
The Proterozoic ophioliteproblem
81
Fig. 9. Photograph showing a cleaned outcrop of sheeted mafic dikes (at Locality 1,Fig. 8). Length of hammer is 35 cm.
numerous dikes, which are always chilled against the MORB-like dikes, resemble modern ocean-island tholeiities (Figs. 3d, 3e). Coarse-grained mafic rocks with cumulus textures are volumetrically the largest component of the Watts Group in the eastern part of the Cape Smith belt. Compositional layering is defined by modal variations of the metamorphic products of primary plagioclase and clinopyroxene, and individual layers range in thickness from the centimetre- to metre-scale. Numerous stratiform lenses of layered ultramafic rocks with gradational contacts are found throughout the gabbros. However, ultramafic lenses with tectonic contacts are also found. The thickness of segment 2 of the layered gabbro sequence (Fig. 7) is about 2100 m, whereas that of segment 4 is about 1800 m. An approximately 2200 m thick unit of layered ultramafic rocks has intrusive contacts with layered gabbros above and below. Compositional layering, defined by modal variations in relict primary clinopyroxene and olivine, and their respective metamorphic-recrystallization products, ranges from dunite, through wehrlite and clinopyroxenite on the scale of centimetres to tens of metres. Serpentinization of olivine is extensive, and the freshest dunitic rocks rarely contain more than 5-10% relict olivine. Numerous irregularly shaped bodies of coarse-grained, massive clinopyroxenite are found at all levels within the mafic and ultramafic cumulate sequence.
H.H. Helmstaedt and D.J. Scott
82
As shown by Scott (1990), the rocks of the Watts Group belong to two magmatic suites. The older, comprising pillowed and massive basalts, sheeted dikes, layered cumulate rocks which are dominantly gabbroic, and rare plagiogranite, is tholeiitic I 61.
61'
r-------
'2'
I
116'
,
114'
-I I
\z
112'
km
D
Fig. 10. Comparison of map patterns and areal dimensions of various ophiolites (black) and greenstone terrains. A. Cape Smith belt; B. Samail ophiolite, Oman; C. Greenstone terrain of western Slave Province, NWT, Canada; D. Newfoundland. Note the similarity in surface area of the Cape Smith klippe (Purtuniq ophiolite in black, Chukotut and Povungnituk groups dotted) t o that of the Humber Zone ophiolites of Newfoundland and that of the Samail ophiolite. The pattern of the Slave Province greenstone belts is comparable with that of the ophiolites of the Arabian Shield (Fig. 4) and the highly deformed Dunnage Zone of Newfoundland.
f i e Proterozoic ophioliteproblem
83
and compositionally and petrographically similar to rocks formed at modern ocean ridges (Fig. 3d). It is characterized by cNd(t)-values of +3.3 to +4.7 (Hegner and Bevier, 1991). The younger suite comprises some of the sheeted dikes, layered cumulate rocks ranging from mafic to ultramafic, and the massive clinopyroxenite intrusions. It resembles tholeiitic suites from modern oceanic islands (Fig. 3e) and has cNd(t)-values ranging from +3.0 to +3.4 (Hegner and Bevier, 1991). The older suite is interpreted to have originated near a constructive plate margin, whereas the younger suite suggests hot-spot related magmatism in an oceanic-island setting (Scott, 1990). The igneous rocks of the Watts Group include all of the igneous crustal members of an ophiolite assemblage as defined by the Penrose Conference participants (Anonymous, 1972). However, tectonized harzburgite, thought to represent obducted-oceanic mantle, has not yet been identified in the Cape Smith belt and, in addition, the Watts Group is dismembered by thrust faults. Nevertheless, as Anonymous (1972, p. 25) stated explicitly that “faulted contacts between mappable units are common” and that “whole sections may be missing”, the Watts Group may be viewed as an incomplete and structurally dismembered ophiolite, and the name “Purtuniq ophiolite”, proposed by St-Onge et al. (1988) is in keeping with the original ophiolite terminology. The Purtuniq ophiolite represents a n unusually large fragment of exceptionally well preserved oceanic crust that can be reconstructed with relative certainty in spite of the tectonic dismemberment and moderate metamorphic recrystallization. An important reason for the preservation of such a large and coherent thrust stack (see also Fig. 10) is the absence of voluminous post-emplacement granitoid batholiths suggesting that, apart from the intrusion of small tonalitic plutons between 1880 and 1840 Ma (Parrish, 1989), the Superior Province basement beneath the Cape Smith klippe was not remobilized significantly after obduction. Although one can only speculate as to the reasons for the absence of an uppermantle portion from the Purtuniq ophiolite, this might be explained by the unusual thickness of the crustal section, which with 9 km (considered to be a minimum estimate by Scott, 1990) is significantly higher than that of most Phanerozoic ophiolites (Fig. 1) and may have been a limiting factor during the obduction process (see also Moores, 1986). The older age of the ophiolite with respect to the rift-related volcanic sequence of the Povungnituk Group to the south (1998 Ma vs. ca. 1960 Ma, Parrish, 1989) shows that the Early Proterozoic ocean basin of the Purtuniq ophiolite did not form by a south to north progression from rift to ocean, ruling out a simple Wilson Cycle for the evolution of the suture in the vicinity of the Cape Smith belt. Proterozoic accreted terranes of southwestern United States
As pointed out by Hoffman (1988), most of the southwestern United States is underlain by juvenile Early Proterozoic crust that was accreted to the Archean Wyoming Province of Laurentia between about 1800 and 1600 Ma (see also
84
H.H. Helmstaedt and D.1Scott
Karlstrom and Houston, 1984; Condie, 1986; Bennett and DePaolo, 1987) (Fig. 5). Proterozoic rocks in Arizona, exposed in the transition zone between the Colorado Plateau and the Basin and Range Province, are the result of crust formation during the Yavapai (1790-1690 Ma) and Mazatzal (1710-1620 Ma) cycles. Volcanicplutonic suites and associated greywacke-pelite-facies sedimentary rocks of the older cycle, interpreted as relics of island arcs and associated sedimentary basins (Condie, 1982b, 1986), dominate in northwestern Arizona, whereas felsic volcanic rocks and shelf-type sedimentary rocks of the younger cycle occur mainly in central and southeastern Arizona. Rocks of the younger cycle unconformably overlie the older arc rocks in central Arizona and were deformed and thrust northwestwards over the older rocks during the Mazatzal orogeny at ca. 1650-1600 Ma (Karlstrom et al., 1987; Karlstrom and Bowring, 1988). The geological setting of an 1730 Ma ophiolite within Yavapai-cycle rocks of the Mazatzal block, near Payson in central Arizona (Fig. 5 ) resembles that of the ophiolites in the Arabian Shield, with the exception that the Payson ophiolite is younger than adjacent arc rocks and appears to be preserved in situ in an intra-arc basin (Dann, 1991). The Payson ophiolite As described by Dann (1991), the Payson ophiolite consists of layered gabbros, gabbros and quartz diorites, tonalitic dikes and plutons, mafic dike swarms, sheeted dikes, and submarine basalts, emplaced within a magmatic arc complex and overlain by a thick sequence of turbidites. Several domains of sheeted dikes lie between coarse gabbro in the northeast and mafic flows in the west and southwest, suggesting pseudostratigraphic relationships typical of those in crustal portions of ophiolites (Fig. 1). The dikes strike northwest and dip about 70" northeast, implying that the entire ophiolite section may have been tilted about 20" to the southwest. As the dikes and gabbros show intrusive contacts with the arc rocks, the ophiolite is thought to be in its in situ position, and the absence of exposed upper mantle rocks is explained by incomplete rifting of the arc basement. Chemical compositions of the mafic dikes are characteristic of island-arc tholeiite and show the selective enrichment of large-ion-lithophile over high-fieldstrength elements, relative to MORB, typical of basalts from some back-arc basins (Saunders and Tmrney, 1984) and supra-subduction-zone ophiolites (Pearce et al., 1954) (Fig. 30. A U-Pb zircon age of 1730 Ma on diorite from the ophiolite is bracketed by U-Pb zircon ages of 1750 Ma from granodiorite of the arc basement and 1720 Ma from ash beds in the overlying turbidites (unpublished data by Bowring, as quoted in Dann, 1991). The tectonic environment, chemical characteristics, and age constraints of the Payson ophiolite are thus consistent with formation in a short-lived intra-arc basin (Dann, 1991). DIVERSITY OF PROTEROZOIC OPIlIOLITES
As shown in the preceding paragraphs, remnants of tectonically dismembered fragments of ocean floor have now been identified in Early as well as Late
The Proterozoic ophioliteproblem
85
Proterozoic orogenic belts. Although most are incomplete, and all are highly deformed, sequences of the nearly 2000 Ma Jormua and Purtuniq complexes display sheeted-dike complexes and are sufficiently well preserved to qualify as ophiolites in the sense of the Penrose conference definition (Anonymous, 1972). One of the most important former arguments against the operation of the Wilson Cycle in the Early Proterozoic, based on the apparent lack of physical evidence for spreading in an oceanic environment, is thus obviated. In spite of the relatively small number of preserved examples, it is evident that there is no principal difference between Early and Late Proterozoic ophiolites, and that Proterozoic ophiolites on the whole are as diverse as their Phanerozoic counterparts. This diversity is roughly comparable with the division by Moores (1982) of Phanerozoic ophiolites into a Tethyan type, emplaced onto a welldeveloped passive margin, and a Cordilleran type, the emplacement of which is less clear-cut, but probably involves preservation of oceanic lithosphere or back-arc basin floor between accreted island arc and microcontinental terranes. The Proterozoic rock record shows a spectrum from small fragments of intra- or back-arc-basin floor and pieces of oceanic lithosphere trapped between juvenile arc terranes (e.g., Arabian Shield, Payson ophiolite, Flin-Flon domain) to areally more extensive ophiolites emplaced as part of major thrust sheets in fold and thrust belts directly on the deformed margins of older continental cratons (e.g., Bou Azzer (?), Jormua, Purtuniq). As both end members of this spectrum existed already in the Early Proterozoic, it is unnecessary to invoke a transition in tectonic style from predominantly “ensialic orogeny” to “collisional orogeny”, caused by Phanerozoic-style plate tectonics, between the earlier and later parts of the Proterozoic (e.g., Kroner, 1983). If both orogenic styles are viable, they must have coexisted throughout the entire Proterozoic. The expanded ophiolite association of Moores (1982) serves as a useful working model for the recognition of remnants of Proterozoic ophiolites. As discussed by Moores (1982), the presence or absence of the various ophiolite units, their nature and thickness, and the tectonic style are a function of the tectonic setting of the ophiolite that must be inferred as much from the associated rock types as from the ophiolite itself. Whereas the diversity of the ophiolites depends on the type of oceanic environment in which they form, the diverse emplacement styles and thus the chances of survival of Proterozoic ophiolites appear to be controlled largely by the size of the continental fragments involved in plate accretion, that in turn determines the intensity and type of suturing (see also Dewey, 1977). Arguments concerning the ultimate origin of Proterozoic ophiolites, including the question whether any of them represent “true” Proterozoic ocean floor, are thus as fraught with uncertainties as those about their Phanerozoic counterparts. Although questions regarding the nature and size of Early Proterozoic oceans (i.e., the extent of possible Wilson Cycles) will also remain debatable, it is reassuring that the locations of remnants of ocean floor can he predicted using the basic tenets of Phanerozoic plate tectonics. The “Proterozoic ophiolite problem” is thus not one of lack of preservation (Moores, 1986), but one of recognizing
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these remnants and restoring them to their pre-deformation configuration. The Jormua complex was recognized as an ophiolite (Kontinen, 1987) as the result of detailed structural work in an area that was previously interpreted as a possible suture zone (for a review, see Park, 1991). The Purtuniq ophiolite was identified as the result of the structural mapping by St-Onge and his co-workers (St-Onge e t al., 1988) near a major suture recognized by earlier workers (see Hoffman, 1985). As predicted by Moores (1986), it is located in the structurally highest slice of a succession grading from continental to oceanic mafic volcanic rocks (Baragar and Scoates, 1981; Hynes and Francis, 1982). As the crustal sections of these two ophiolites differ vastly in complexity and thickness (Fig. l ) , it is too early to derive general conclusions about the emplacement of Proterozoic ophiolites. However, the fact that only the thinner Jormua ophiolite section includes a piece of possible upper mantle supports the suggestion of Moores (1986) that the maximum total thickness of ophiolites (10-12 km) may be related to the ability of the buoyant continental margin to lift ocean floor above sea level. During obduction, an exceptionally thick slice of oceanic crust would be expected to detach from its mantle portion which is subsequently subducted. Such mechanism may account for the lack of tectonized harzburgite in the thick Purtuniq section. Late Proterozoic ophiolites have now been identified also in previously suspected suture zones of India (Phulad ophiolite in Rajasthan suture; see Volpe and MacDougall, 1990) and the Carpathian region of Europe (Stara Planina ophiolite in the Thracian suture, Bulgaria; see Haydoutov, 1989), and it is likely that the recent re-interpretations of several Proterozoic orogens in Australia as collisional belts (e.g., Myers, 1990) will lead to further discoveries of old ophiolites.
OPHIOLITE ANALOGUES IN ARCHEAN GREENSTONE BELTS?
Although typical ophiolite sections have not been described from Archean greenstone belts, it has been proposed that some of these belts represent ancient oceanic crust and thus are possible ophiolite analogues (e.g., Harper, 1986; Helmstaedt et al., 1986; D e Wit et al., 1987). Windley (1981) suggested that Archean greenstone belts are “proto-ophiolites” formed by extensive rifting in marginal basins prior to the development of stable cratons in the Early Proterozoic. However, evidence for the existence of sizeable continentaI plates between 3200 and 2500 Ma, including some with diamondiferous roots, has been accumulating (e.g., Helmstaedt and Schulze, 1989; Gurney, 1990), and it now appears likely that the Cordilleran and Tethyan types of ophiolites recognized in the Early Proterozoic developed earlier, as the size of continental plates gradually increased during Late Archean time. This is confirmed by the observation that extent and thickness of ocean-floor material of the Early Proterozoic Purtuniq ophiolite are already comparable with those of the larger Phanerozoic ophiolites (Figs. 10a,b,d). Although many Archean volcanic belts may be products of ensialic rifting or arc volcanism, it is proposed here that some greenstone belts may be
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analogous to ophiolites emplaced by (1) wedging of ocean-floor material between accreted arc terranes and (2) thrusting of oceanic crust onto continental crust. The remarkable similarity of the juvenile arc terranes of the eastern Arabian Shield with Archean and Early Proterozoic greenstone belts was noticed earlier by Engel et al. (1980) (compare Figs. 4 and lOc). As pointed out initially by Brown and Coleman (1972), the Arabian ophiolites are aligned in sutures between various accreted arc terranes and microcontinental fragments, and properties and ages of many of these ophiolites are consistent with their being remnants of ocean floor on which the island arcs were built (Pallister et al., 1988). A similar model was proposed by Hoffman (1986) for the formation of the greenstone belts of the Slave Province of the Canadian Shield, and the N-S alignment of greenstone belts in that province (Fig. 1Oc) was interpreted as a suture between an older microcontinent (Anton terrane) in the west and an accretionary terrane (Contwoyto terrane) in the east (Kusky, 1989). The distribution pattern of greenstone belts in the western Slave Province is comparable also with that of ophiolites in younger orogenic belts, such as the Dunnage zone of the Newfoundland Appalachians (Fyson and Helmstaedt, 1988) (Figs. lOc, 10d), that appear to represent remnants of thrust sheets dismembered by later deformation and magmatic activity (Dunning and Chorlton, 1985). The Cape Smith belt (Fig. lOa), comprising a large thrust stack of mainly mafic volcanic rocks similar to rock types found in Archean greenstone belts (e.g., Condie, 1981), provides an Early Proterozoic example of allochthonous greenstone belt formation and may be used to illustrate the difficulties of recognizing oceanic crust in Archean greenstone belts. Unlike most other greenstone belts, the Cape Smith belt has not been modified extensively by post-emplacement granitoid intrusions and late faulting, and primary lithologic and structural relationships are well preserved. All three volcanic suites within the belt, the Povungnituk, Chukotat, and Watts Groups, derived from a continental-rift-margin, a rift-ocean-transitional, and a truly oceanic setting, respectively, are tectonically juxtaposed, and the continental-rift sequence is structurally imbricated with the sedimentary rocks of the continental margin (for a summary, see St-Onge and Lucas, 1990a). Rock types of all three suites could be interpreted as “greenstone belts”, but only one of these, the Watts Group, comprises truly oceanic rocks. However, as the Watts Group is also imbricated by numerous thrusts, and the ophiolite structure did not become apparent until detailed cross sections were constructed, it is unlikely that the individual ophiolite units could have been pieced together if the thrust sheets of the Cape Smith belt were dismembered by pervasive granitoid intrusions and other post-emplacement deformation. The existence of thrust stacks with ophiolites of Purtuniq dimensions (Fig. 10a) in the Early Proterozoic Cape Smith belt clearly suggests that obduction of oceanfloor rocks onto continental basement did not ’suddenly’ begin in the Proterozoic. Remnants of allochthonous slices of even more ancient sea-floor rocks are almost certainly present in Archean greenstone belts (e.g., D e Wit et al., 1987; Fyson and Helmstaedt, 1988; Kusky, 1989; Armstrong et al., 1990), but their ophiolitic
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nature is even more disguised than that of their Proterozoic counterparts as a consequence of dismemberment during obduction and multiple post-emplacement intrusive and deformation events. Thus, strict adherence to the Penrose conference model of an ophiolite may not be helpful in recognizing these rocks.
ACKNOWLEDGEMENTS
We are grateful to Marc St-Onge and Steve Lucas of the Geological Survey of Canada for providing the opportunity to partake in the discovery and study of the Purtuniq ophiolite during their mapping project in the Cape Smith belt of northern Quebec. The first author thanks Karl Kasch for the hospitality on his farm (his Ph.D. thesis area) near Omitara, Namibia, during an unforgettable field trip through the Damara belt in 1981. Whereas the Purtuniq ophiolite appears to represent one of the best preserved pieces of Proterozoic ocean floor, the ’ophiolite’ on Karl’s farm is definitely one of the most dispersed. Karl nevertheless presented a convincing argument for a suture and managed to assemble all the important rock types of an ophiolite within easy walking distance. Chris Hartnady provided stimulating discussions on Damara belt geology during a recent visit a t the University of Cape Town. Martin van Kranendonk, at Queen’s University, was particularly helpful in finding references. Nancy Thomas drafted most of the figures, and Karin Helmstaedt proofread the manuscript. Careful and constructive reviews by K. Condie, Y. Dilek, and J.S. Pallister led to significant improvements of the paper. H.H. Helmstaedt’s research is financed by NSERC Canada operating grant A 8375.
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Lucas, S.B., 1989. Geometrical, rheological and mechanical evolution of continental trust belts: examples from the Cape Smith belt, Northern Quebec, Canada. Ph.D. thesis, Brown University, Providence, R.I., 253 pp. Lucas, S.B. and St-Onge, M.R., 1991. Evolution of Archean and early Proterozoic magmatic arcs in northeastern Ungava Peninsula, Quebec. Geol. Surv. Can., Pap., 91-1C 109-119. Martin, H. and Porada, H., 1977. The intracratonic branch of the Damara Orogen in South West Africa, I. Discussion of geodynamic models. Precambrian Res., 5: 311-338. McCall, G.J.H., 1981. Progress in research into the early history of the earth: a review, 1970-1980. Spec. Publ. Geol. SOC.Aust., 7: 3-18. Moores, E.M., 1982. Origin and emplacement of ophiolites. Rev. Geophys. Space Phys., 20: 735-760. Moores, E.M., 1986. The Proterozoic ophiolite problem, continental emergence and the Venus connection. Science, 234: 65-68. Myers, J.S., 1990. Precambrian tectonic evolution of part of Gondwana, southwestern Australia. Geology, 18: 537-540. Naidoo, D.D., Bloomer, S.H., Saquaque, A., and Hefferan, K., 1991. Geochemistry and significance of metavolcanic rocks from the Bou h e r - E l Graara ophiolite (Morocco). Precambrian Res., 53: 79-97. Nassief, M.O., Macdonald, R., and Gass, I.G., 1984. The Jebel Thunvah Upper Proterozoic ophiolite complex, western Saudi Arabia. J. Geol. SOC.London, 141: 537-546. Pallister, J.S., Cole, J.S., Stoeser, D.B., and Quick, J.E., 1990. Use and abuse of crustal accretion calculations. Geology, 18: 35-39. Pallister, J.S., Stacey, J.S., Fischer, L.B., and Premo, W.R., 1987. Precambrian ophiolites of Arabia: A summary of geologic settings, U-Pb geochronology, lead isotope characteristics, and implications for microplate accretion. Saudi Arabian Deputy Ministry for Mineral Resources Open-File Report USGS-OF-07-10, pp. 81. Pallister, J.S., Stacey, J.S., Fischer, L.B., and Premo, W.R., 1988. Precambrian ophiolites of Arabia: geologic settings, U-Pb geochronology, Pb-isotope characteristics, and implications €or continental accretion. Precambrian Res., 38: 1-54. Park, A.F., 1988. Nature of the Early Proterozoic Outokumpu assemblage, eastern Finland. Precambrian Res., 38 131-146. Park, A X , 1991. Continental growth by accretion: A tectonostratigraphic terrane analysis of the evolution of the western and central Baltic Shield, 2.50 to 1.75 Ga. Geol. SOC.Am. Bull., 1 0 3 522537. Park, A X , Bowes, D.R., Halden, N.M. and Koistinen, TJ., 1984. Tectonic evolution at an early Proterozoic continental margin: the Svecokarelides of eastern Finland. J. Geodyn. 1: 359-386. Parrish, R.R., 1989. U-Pb geochronology of the Cape Smith Belt and Sugluk block, northern Quebec. Geosci. Can., 16: 126-130. Pearce, J.A., 1982. Trace element characteristics of lavas from destructive plate boundaries. In: R.S. Thorpe (Editor), Andesites. John Wiley and Sons, London, pp. 52.5-548. Pearce, J.A., Lippard, S.J. and Roberts, S., 1984. Characteristics and tectonic significance of suprasubduction zone ophiolites. In: B.P. Kokelaar and M.E Howells (Editors), Marginal Basin Geology. Geol. SOC.London, Spec. Publ., 16: 77-94. Quick, J.E., 1990. Geology and origin of the Late Proterozoic Darb Zubaydah ophiolite, Kingdom of Saudi Arabia. Geol. SOC.Am. Bull., 102 1007-1020. Quick, J.E., 1991. Late Proterozoic transpression on the Nabitah fault system - implications for the assembly of the Arabian Shield. Precambrian Res., 53: 119-147. Sangster, D.F., 1972. Isotopic studies of ore-leads in the Hanson Lake-Flin Flon-Snow Lake mineral belt, Saskatchewan and Manitoba. Can. J. Earth Sci., 9: 500-513.
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Sangster, D.E, 1978. Isotopic studies of ore-leads of the circum-Kisseynew volcanic belt of Manitoba and Saskatchewan. Can. J. Earth Sci., 15: 1112-1121. Saquaque, A., Admou, H., Karson, J., Hefferan, K. and Reuber, I., 1989. Precambrian accretionary tectonics in the Bou Azzer-El Graara region, Anti-Atlas, Morocco. Geology, 17: 1107-1110. Saunders, A.O. and Tarney, J., 1984. Geochemical characteristics of basaltic volcanism within back-arc basins. In: B.P. Kokelaar and M.E Howells (Editors), Marginal Basin Geology, Geol. SOC., Spec. Pub]., 16, pp. 77-94. Schulz, K.J.,1987. An early Proterozoic ophiolite in the Penokean orogen. Geol. Assoc. Can., Program with Abstracts, 12 87. Scott, D.J., 1990. Geology and geochemistry of the early Proterozoic Purtuniq ophiolite, Cape Smith Belt, Quebec. Ph.D. Thesis, Queen’s University, Kingston, Ont., pp. 289. Scott, D.J. and Bickle, M.J., 1991. Field relationships in the early Proterozoic Purtuniq ophiolite, Lac Watts and Purtuniq map areas, Quebec. Geol. Sum. Can., Pap., 91-1C 179-188. Scott, D.J., St-Onge, M.R., Lucas, S.B. and Helmstaedt, H., 1989. The 1998 Ma Purtuniq ophiolite: imbricated and metamorphosed oceanic crust in the Cape Smith Thrust Belt, northern Quebec. Geosci. Can., 16: 144-148. Scott, D.J., St-Onge, M.R., Lucas, S.B., and Helmstaedt, H., 1991. Geology and chemistry of the early Proterozoic Purtuniq ophiolite, Cape Smith, northern Quebec, Canada. In: Tj. Peters (Editor), Ophiolite Genesis and Evolution of the Oceanic Lithosphere. Kluwer, Dordrecht, pp. 825-857. Searle, M.P., 1991. Conference report: Symposium on ophiolite genesis and evolution of oceanic lithosphere. J. Geol. SOC.London, 148 203-204. Searle, M.P. and Stevens, R.K., 1984. Obduction processes in ancient, modern and future ophiolites. In: I.G. Gass et al. (Editors), Ophiolites and Oceanic Lithosphere. Geol. SOC.London, Spec. Publ., 13 303-320. Shackleton, R.M., Ries, A.C., Graham, R.H. and Fitches, W.R., 1980. Late Precambrian ophiolite melange in the eastern desert of Egypt. Nature, 285: 472-474. Shanti, M. and Roobol, M.J., 1979. A late Proterozoic ophiolite complex at Jabal Ess in northern Saudi Arabia. Nature, 279: 488-491. Simonen, A., 1980. The Precambrian in Finland. Geol. Sum. Finl. Bull., 304 1-58. Smith, H.S. and Hartnady, C.J.H., 1984. Geochemistry of Grootderm Formation lavas: indication of iectonic environment of extrusion. Abstr. Conf. Middle to Late Proterozoic Lithosphere Evolution, University of Cape Town, pp. 20-21. Stauffer, M.R., 1984. Manikewan: an early Proterozoic ocean in central Canada, its igneous histoly and orogenic closure. Precambrian Res., 25: 257-281. Stoeser, D.B. and Camp, V.E., 1985. Pan-African microplate accretion of the Arabian Shield. Geol. SOC. Am. Bull., 9 6 817-826. St-Onge, M.R. and Lucas, S.B., 1989. Geology, eastern portion of the Cape Smith Thrust-Fold Belt, parts of the Wakeham Baye, Cratere du Noveau-Quebec and Nuvilik Lakes map areas, northern Quebec. Geol. Sum. Can., Maps 1721A to 1735A, scale 1: 50 000. St-Onge, M.R. and Lucas, S.B., 1990a. Evolution of the Cape Smith Belt: Early Proterozoic continental underthrusting, ophiolite obduction, and thick-skinned folding. In: J.E Lewry and M.R. Stauffer (Editors), The Early Proterozoic Trans-Hudson Orogen of North America. Geol. Assoc. Can., Spec. Pap., 37: 313-351. St-Onge, M.R. and Lucas, S.B., 1990b. Early Proterozoic collisional tectonics in the internal zone of the Ungava (Trans-Hudson) orogen: Lacs Nuvilik and Sugluk map areas, Quebec. Geol. Surv. Can., Pap., 90-1C 119-132. St-Onge, M.R., Lucas, S.B., Scott, D.J. and BCgin, N.J., 1987. Tectonostratigraphy and structure of the Lac Watts-Lac Cross -Riviere Deception area, central Cape Smith Belt, northern Quebec. Geol.
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Surv. Can., Pap., 87-1A 619-632. St-Onge, M.R., Lucas, S.B., Scott, D.J., Begin, N.J., Helmstaedt, H. and Carmichael, D.M., 1988. Thinskinned imbrication and subsequent thick-skinned folding of rift-fill, transitional-crust and ophiolite suites in the 1.9 Ga Cape Smith Belt, northern Quebec. Geol. Surv., Can. Pap., 88-1C 1-18. Van Schmus, NR., Bickford, M.E., Lewty, J.E and Macdonald, R., 1987. U-Pb geochronology in the Trans-Hudson orogen, northern Saskatchewan, Canada. Can. J. Earth Sci., 24: 407-424. Vayrynen, H., 1939. On the geology and tectonics of the Outokumpu ore field and region. Camrn. Geol. Finl. Bull., 124. Watters, B.R. and Pearce, J.A., 1987. Metavolcanic rocks of the La Ronge domain in the Churchill province, Saskatchewan: Geochemical evidence for a volcanic arc origin. In: TC. Pharaoh et al. (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC.London, Spec. Publ. 33, pp. 167-182. Volpe, A.M. and MacDougall, J.D., 1990. Geochemistry and isotopic characteristics of mafic (Phulad ophiolite) and related rocks in the Delhi Supergroup, Rajasthan, India: implications for rifting in the Proterozoic. Precambrian Res., 48: 167-191. Wegrnann, C.E., 1928. Uber die Tektonik der jungeren Faltung in Ostfinnland. Fennia, 50 (16). Wilson, J.T., 1968. Static or mobile earth: The current scientific revolution. Am. Philos. SOC.Proc., 112: 309-320. Windley, B.E, 1981. Precambrian Rocks in the light of the plate tectonic concept. In: A. Kroner (Editor), Precambrian Plate Tectonics. Elsevier, Amsterdam, pp. 1-20. Windley, B.E, 1984. The Evolving Continents. Wiley, New York, N.Y., 399 pp.
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Chapter 3
PROTEROZOIC RIFTS J.C. GREEN
INTRODUCTION
By the end of the Archean, a period of intense volcanism, plutonism, accretion and cratonization had led to the establishment of more or less stable continental platforms. These have been estimated (Condie, 1989b) to have covered as much as 60% of the area of the present continents. These cratons have subsequently been affected by extensional forces, presumably due to mantle diapirism and asthenospheric flow, that led to the production of a variety of Proterozoic crustal features. Some of these extensional regimes appear to have produced only isolated ensialic basins, whereas others were more linearly organized and led to major continental rifts, both abortive and successful. The change to the “modern” stable cratonhifting behavior was not abrupt (see also Burke and Dewey, 1973; Kroner, 1977); several large, rifted intracratonic basins formed during the Late Archean (e.g. in Australia: Blake and Groves, 1987; in South Africa: Thnkard et al., 1982). One of the more controversial aspects of Phanerozoic continental rifting has been the question of the thermotectonic origin of individual rifts. In essence, the contrasting models describe a “passive” rift as resulting from exogenic forces, transmitted through a rigid lithosphere, that result in fracturing and attenuation of the lithosphere (e.g., Baikal rift); this in turn induces magmatism from decompression melting of upper mantle (as described by McKenzie and Bickle, 1988). An “active” rift is one generated by thinning over an endogenic mantle plume or hot spot (Burke and Dewey, 1973; Baker and Morgan, 1981; Turcotte and Emerman, 1983). These plumes may in turn be the result of energy exchanges at the core/mantle boundary. Such arguments about the active/passive character and origin apply as well to Proterozoic rifts but in general have not been extensively explored. This chapter reviews Proterozoic cratonic rifts. Principal attention is given to the major continental rifts, with lesser concentration on isolated ensialic basins and on back-arc rifts associated with convergent plate boundaries. It is not intended to be an exhaustive survey. Other extensional regimes are treated in the chapter on Proterozoic dike swarms (Thrney, this volume). Layered intrusions (von Gruenewaldt and Harmer, this volume), the anorthosite suite (Wiebe, this volume), and anorogenic granites (Anderson and Morrison, this volume) may also
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be related to extension but are less clearly associated with tectonic structures and supracrustal rock assemblages than the rifts reviewed in this chapter. Evidence for major extensional tectonism comes from several aspects of Proterozoic crustal elements, including the lithology, stratigraphic sequence, and thickness of strata and their structural and stratigraphic relations to older crust. The accumulation of a significant thickness of mafic volcanic rocks in a cratonic setting is considered a prime criterion for rifting. The geochemistry of the igneous rocks can also be useful. In many cases the rocks have been so recrystallized during metamorphism that they have lost most or all vestiges of primary textures and structures, leaving their gross stratigraphic relations and geochemistry as bases for interpretation. Later tectonic dislocations have further obscured origins in some situations, making definitive inferences even less certain. Sedimentary assemblages commonly interpreted as products of continental rifting (Condie, 1989b) include immature, commonly arkosic, fluvial sandstones and conglomerates, indicative of rapidly subsiding, subaerial basins (these are typically red-beds in Middle to Upper Proterozoic sequences); lacustrine deposits; evaporites; stable-shelf deposits including some carbonates and iron-formations; and possibly turbiditic graywackes and pelites deposited on trailing continental margins or rapidly subsiding ensialic basins. Volcanism is a very important element of rift evolution, generally coming at an early stage and either continuing throughout rifting or fading as sedimentation increases. The volcanic rocks are typically bimodal in composition, with mantlederived tholeiitic mafic rocks dominant and subordinate rhyolites produced largely by partial melting of the continental crust. Alkalic rocks may be present. Numerous geochemical diagrams and correlations have been published in an attempt to infer the tectonic environment of volcanic rocks from their major and trace elements (e.g. Pearce and Cann, 1973; Pearce et al., 1977; Wood, 1980; Mullen, 1983; Meschede, 1986). Relying principally on those elements less likely to have been mobilized during rock alteration or metamorphism, these discriminants may be the only remaining evidence for a rift origin for many metamorphosed volcanic units. However, these discriminants are calibrated from a limited number of Phanerozoic and especially Cenozoic suites of known tectonic situation, and many anomalies (that is, rocks of well-known tectonic setting with geochemistry that does not correspond well with the diagrams) have been reported. In particular, the Ti/Y/Zr diagram has been shown to imply incorrect tectonic setting in several provinces (Holm, 1982; Duncan, 1987). Therefore, these discriminant diagrams cannot be considered completely definitive. Even assuming their general utility, there are many Proterozoic igneous suites for which a tectonic setting remains ambiguous on the basis of these geochemical characteristics. Whether this is the result of rock alteration (element mobility), substandard analyses, or the inherent limitations of the correlations is not always clear. Subcontinental mantle source heterogeneity is probably significant, and detailed geochemical and petrological study of each suite may be necessary for a definitive conclusion (e.g. Marsh, 1987; Myers and Breitkopf, 1989). In general, a rift origin has been inferred by most workers where
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the major-element geochemistry shows the characteristics of tholeiitic (as opposed to calc-alkaline) magmas (e.g., generally high Ti; Fe-enrichment) and within-plate trace-element ratios. Rifts that develop to the extent of drastically thinning or separating the continental crust can produce basalts with the characteristics of back-arc or marginal basins or even mid-ocean ridge basalts (MORBs). The geochemical character of Proterozoic mafic rocks is useful as a probe of the ancient mantle and its evolution, using growth models for the radiogenic isotopes (Patchett, this volume). Condie (1989a) has found that Proterozoic basalts in general appear to have been derived from mantle sources enriched relative to those of typical Archean greenstones, perhaps as a result of extensive postArchean erosion and recycling of Late Archean crust. Evidence from U/Pb, Nd/ Sm, and Rb/Sr systems also can be useful in determining the degree of interaction of magmas with continental lithospheric mantle and crust. “Successful” continental rifts - those that have produced a new ocean - have freed the cratonic fragments to drift in some remarkable trajectories. It is thus not a straightforward matter to search out evidence for rifting on once-contiguous but now widely separated cratons. A combination of paleomagnetism, geochronology, geochemistry, and tectonic analysis is necessary to reconstruct Proterozoic plate configurations. Several recent studies, for instance, have proposed more or less similar arrangements for a supercontinent in the latest Proterozoic (e.g. Bond et al., 1984; Moores, 1991; Dalziel, 1991; Hoffman, 1991). Continental rifting reached major culminations at three times during the Proterozoic: at 2.0-1.8 Ga, 1.2-0.9 Ga, and at 0.8-0.6 Ga. This “pulse of the Earth” is discussed by Hoffman (1989) who attributes the onset of the activity to the accumulation of radiogenic heat beneath old, stable supercontinents far from subducting margins. In particular, the end of the Proterozoic was a period of widespread continental breakup, before Paleozoic reassembly of the fragments in Gondwanaland and Pangaea (Hoffman, 1991). In the remaining portion of this chapter, Proterozoic cratonic rifts will be reviewed in general order of decreasing age.
EARLY PROTEROZOIC RIFTS: 2.5-1.9 Ga
The cratonic framework established near or at the end of the Archean through widespread volcanic arc formation, accretion, and plutonism, remained remarkably stable during the following few hundred million years. Although some major ensialic sedimentary basins developed (e.g. Witwatersrand, Bansvaal, Huronian), and significant mafic dike swarms and other mantle-derived intrusions were emplaced during this period, little evidence of major continental rifting is found until about 2.1 f 0.1 Ga. The large mafic intrusive events of this period probably reflect incipient cratonic rifting, and include the Great Dyke in the Zimbabwe craton (2461 & 16 Ma;
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Podmore and Wilson, 1987; Wilson and Prendergast, 1989); the Jimberlana intrusion of Western Australia (2420 f 30 Ma; McClay and Campbell, 1976); the roughly linear belt of mafic and ultramafic Koillismaa intrusions cutting the Kola Block of northern Finland and adjacent Russia (about 2440 Ma; Alapieti et al., 1990); the 2.45 Ga Matachewan dikes of Ontario (Nelson et al., 1990); the Scourie dikes of northwest Scotland (2390 Ma; Chapman, 1979; Weaver and Tirney, 1981); the 2220 Ma Nipissing diabase complex of the southern Canadian Shield (Corfu and Andrews, 1986); and the 2120 f 67 Ma Kenora-Kabetogama dike swarm of northern Minnesota and adjacent Canada (Southwick and Day, 1983; Southwick and Halls, 1987). South Afiican basins
Several large depositional basins developed during Early and Middle Proterozoic time on the Archean cratons of southern Africa (Tankard et al., 1982) (Fig. 1). Although their tectonothermal origins are not clear, those that show evidence of cratonic rifting as opposed to more even subsidence will be briefly reviewed here. The Ventersdorp Supergroup of north-central South Africa, previously commonly regarded as Proterozoic, is a largely volcanic sequence up to 8 km thick that was emplaced on rifted basement and Witwatersrand Supergroup supracrustal rocks. Its age has been problematic, but recent zircon U/Pb ages establish it as Late Archean (2.75-2.70 Ga; Myers et al., 1987; Crow and Condie, 1988). The Transvaal and Griqualand Supergroups constitute a major rift-related accumulation in the Kaapvaal craton, originally covering at least 500 000 km2 .'
7
ZIMBABWE
i
i
Fig. 1. Northeastern part of South Africa and environs showing Proterozoic basins. J = Johannesburg; = Swaziland. Solid line: approximate limit of Waterburg basin (9;dashed line: Soutpansberg (5') and correlative (?) Palapye (P)basins; dotted line: approximate limit of Transvaal basin. Shaded: approximate limit of Ventersdorp Supergroup (VS).After several sources.
SW
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(Fig. 1). This sequence is up to 12 km thick in the northeastern Pansvaal, and it has been dated at about 2250 Ma (Rb/Sr; see Bnkard et al., 1982) to about 2150 Ma (Rooiberg Felsite; Twist and French, 1983). Primary textures are preserved; the rocks are essentially undeformed, but have undergone some burial metamorphism. Initial subsidence produced fluvial conglomerates, immature sandstones, siltstones and shales unconformably on Ventersdorp volcanic rocks and Archean basement. These are overlain by 800 m of bimodal volcanic rocks (low-Al continental tholeiites and high-K rhyolites), interbedded with lacustrine and fluvial sedimentary rocks. Transgression of a sea brought deposition of deltaic and shallow-marine clastic and carbonate rocks. Cyclic tidal and subtidal deposition produced limestone, dolostone, chert, and iron-formation, some of which contains pseudomorphs after glass shards. A major prograding event followed, depositing the Pretoria and Postmasburg Groups, 2 to 7 km thick. Although this sequence is mainly shallow-marine, several unconformities record periodic uplift and minor erosion. Sedimentary rocks in this sequence include shale, quartz arenite, and carbonates. In addition, mafic volcanic rocks were erupted, both subaerial (Hekpoort and Dullstroom in the Transvaal sequence) and subaqueous (Ongeluk in Griqualand West, Machadodorp in Transvaal). At the top of the Pansvaal Supergroup are the Rooiberg felsites, which constitute one of the largest accumulations of felsic volcanic rocks in the geologic record (Xvist and French, 1983). Approximately 3 to 5 km thick, they comprise a sequence of very large, flood-like quartz latite rheoignimbrites and lavas (?). They may be the result of lower crustal melting by the early phases of the mafic Bushveld intrusion which subsequently spread out beneath this basin (Twist and French, 1983; Xvist and Harmer, 1987). Geochemistry of the mafic to intermediate volcanic formations of the Pansvaal Supergroup (Abel Erasmus, Hekpoort, Machadodorp, and Dullstroom) has been recently studied by Crow and Condie (1990). These rocks show both tholeiitic and calc-alkaline character; Machadodorp basalts most resemble MORBs. Crow and Condie (1990) conclude that two or three mantle sources were involved, and some large ion-lithophile element (LILE) enrichment occurred through subduction processes rather than crustal contamination. Circum-Superior belt
The Circum-Superior belt is a large, Early Proterozoic, rift-related feature that forms the northern boundary of the Archean Superior Province of the Canadian Shield (between the Superior and Hearne and Rae provinces, Fig. 2). It includes several segments, from east to west: the Labrador Trough; the Cape Smith Fold Belt; the Ottawa and Belcher Islands in Hudson Bay; and the Fox River and Thompson belts in Manitoba. The Circum-Superior belt is inferred to have formed by rifting of the Archean continent, which produced true oceanic crust, followed by southward-directed compression leading to ocean closure, over a time span of 1999 to 1922 Ma (Parrish, 1989; Lewry and Stauffer, 1990). Paleomagnetic studies
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Fig. 2. Sketch map of northeastern Canada and adjacent Greenland showing the Circum-Superior belt (shaded) and other rift-related Proterozoic features. A = Aillik Group; AM = Amundsen Embayment; B = Belcher Islands; BU = Borden Basin; CR = Coppermine River Group; CS = Cape Smith Fold Belt; D = Dubawnt Group; FH = Fury and Hecla Basin; FR = Fox River belt; K = Ketilidian Province; LT = Labrador Trough; 0 = Ottawa Islands; P = Payne River dikes; T = Thompson belt; W = Wopmay orogen. Dotted lines show limit of Phanerozoic cover. After Pharaoh et al. (1987) and other sources.
have not been able to determine the amount of opening (ocean width) of this rift (Irving and McGlynn, 1981). In the Cape Smith belt (CSB; Francis et al., 1983; Arndt et al., 1987; Picard et al., 1990; St-Onge and Lucas, 1990), the oldest Proterozoic rocks (Watts Group) constitute an ophiolite sequence which is in tectonic contact with the rest of the belt, but implies the complete early rifting of the Archean basement. It is composed of layered ultramafic cumulates, layered gabbros and anorthosite, sheeted dikes and pillowed tholeiitic basalts. This is one of the world’s oldest ophiolites. Succeeding the Watts Group is the Povungnituk Group, which contains sandstones, conglomerates, dolostones, iron-formations and shales deposited in a continental rift to shelf environment and is overlain by tholeiitic basalts. These are interbedded with more sandstone and siltstone and intruded by mafic and ultramafic sills. Some of the latter contain Cu, Ni, and Pt-group element
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concentrations (Giovenazzo and Picard, 1989). A few alkaline rocks are also present. Geochemically the tholeiites are continental-rift basalts showing light rare-earth (LREE), Fe, and Ti enrichment. The overlying Chukotat Group is composed principally of olivine-, pyroxene-, and plagioclase-phyric pillowed and massive basalts and basaltic komatiites (with up to 17% MgO), with low Ti and Al, and approximately level to LREE-depleted chondrite-normalized REE patterns, geochemically similar to N-MORBs. Spinifex texture is common in the olivine-phyric lavas. These rocks represent a maturing oceanic crust which resulted from rifting of a transitional continental margin. The apparently youngest rocks in the CSB are the Parent Group, which includes clastic sedimentary rocks, subaqueous basalts, and a calc-alkaline volcanic sequence ranging from andesite to rhyolite. %ace-element characteristics include low Cr and imply an origin as a magmatic arc, the result of subduction to the north as the ocean basin began to close. All of these volcanic and sedimentary units were intensely imbricated by south-directed thrust faults during the 1.89-Ga Hudsonian orogeny. The Ottawa and Belcher Islands in eastern Hudson Bay represent the southwestern continuation of the Circum-Superior belt. The Belcher Islands are made of terrestrial and shelf sedimentary rocks interbedded with subaerial, low-Mg, continental tholeiitic basalts and basaltic andesites of the Eskimo and Flaherty Formations, which unconformably overlie Archean basement gneisses and are thought to be approximately equivalent to the Povungnituk Group of the CSB (Arndt et al., 1987). The Ottawa Islands to the north consist largely of subaqueous pillowed and massive and spinifex-textured high-Mg basalts and mafic pyroclastic rocks, rather like the Chukotat Group in geochemistry and appearance. West of Hudson Bay the Fox River and Thompson (Nickel) belts form the further continuation of the Circum-Superior belt (Baragar and Scoates, 1981, 1987). These sequences also include shelf-type metasedimentary rocks and tholeiitic and later komatiitic basalts, many of them pillowed. They are intruded by mafic sills including the 2 km thick, 275 km long, differentiated Fox River sill. Archean gneisses near the Thompson belt are cut by the large, probably comagmatic, NE-trending Molson dike swarm (Ermanovics and Fahrig, 1975). At the east end of the CSB, the Circum-Superior rift bends sharply to the southeast and reappears as the Labrador Bough (New Quebec orogen of Hoffman, 1989). The intervening Payne River dike swarm (Fahrig, 1987) may represent a linkage between these two major rift segments. Subsequent convergence may not have been coeval along the CSB and the New Quebec orogen (Hoffman, 1989). The Labrador %ough/New Quebec orogen (Dimroth, 1972, 1981; Wardle and Bailey, 1981) is similar to the rest of the Circum-Superior belt. It consists of sedimentary rocks deposited unconformably on Superior Province basement, interbedded with mafic volcanic rocks and sills, all metamorphosed and thrust to the southwest during the Hudsonian orogeny. The southern branch of the Archean Rae Province (Hoffman, 1988) lies to its east, and it is bordered to the
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c Presently
preserved part of Trough
EMBRYONIC RIFT
CONTINENTAL CRUST
Gp
Fig. 3. Model for initial rifting phase, Labrador Trough; view to northwest. After Wardle and Bailey (1981).
south by the younger metamorphic and structural Grenville Front. The Labrador Pough consists principally of a continental shelf-slope-rise assemblage, which began with subaerial, coarse clastic deposition in a graben system (Lower Seward Subgroup) (Fig. 3). This was accompanied by minor alkaline volcanism, and was followed by transgression and a complex sequence of shallow marine shelf and basin deposition with thicker, deep-water slope-and-rise turbidite accumulation to the east. The shelf sedimentary rocks include dolostones and the economically important Sokoman Iron-formation, as well as clastic rocks. Mafic volcanic rocks, predominantly submarine, became more abundant and culminated in the Doublet Group which contains over 3 km of pillow basalts. These metabasalts, as well as the abundant mafic intrusions of the Montagnais Group, are all low-K tholeiites of oceanic affinity (according to their Ti/K/P contents and Pearce et al.’s 1977 discriminant), but there are no genuine ophiolite sequences and no evidence for a fully-developed oceanic crust before Hudsonian convergence commenced. Animikie rift
Evidence of rifting is also present along the southern margin of the Superior Province in Minnesota, Wisconsin, Michigan, and Ontario (Fig. 4). This extension, which developed about 2.0-1.9 Ga, is localized roughly along the Great Lakes Tectonic Zone (GLTZ), which is thought to be a suture between an older (3.6-2.7 Ga) gneiss terrane, on the south, and a Late Archean (2.7 Ga) greenstone-granite terrane to the north - the Superior Province proper (Sims and Peterman, 1983; Morey, 1983a,b; 1989). Sims et al. (1951) suggest that the Animikie rifting was the final stage of a long extensional history that began about 2.5 Ga northeast of Lake Huron, led to the ensialic deposition of the thick Huronian Supergroup (2.5-2.2 Ga), and progressively moved to the west to central Minnesota.
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Fig. 4. Sketch map of Appalachian and midcontinent region, North America. Inset shows Newfoundland at same scale. Random dash pattern: Grenville Province. Vertical line pattern: Midcontinent Rift System. Dot pattern: Lower Proterozoic rocks, Great Lakes area. Diagonal line pattern and black areas: Proterozoic rocks of Laurentia (not accreted) of Appalachian orogen. Long dashes: Grenville Front; short dashes: Blue-Green-Long axis of Appalachians; dotted lines: edge of Paleozoic overlap (Great Lakes area) and edge of Cretaceous overlap (Gulf and Cretaceous Coastal Plain).A = Animikie Group; BR = Blue Ridge; GF = approximate position of covered Grenville Front; GLTZ = Great Lakes Tectonic Zone; GM = Grenville metasedimentary rocks; GP = Grenville Province; H = Huronian Supergroup; JB = James Bay; LN = Lake Nipigon; LR = Long Range; S = Sibley Group. States and Provinces: G = Georgia; L4 = Iowa; K = Kansas; L = Labrador; MI = Michigan; MN = Minnesota; N = Newfoundland; NC = North Carolina; 0 = Ontario; Q = Quebec; TN = Tennessee; V = Vermont; I/A = Virginia; W = Wisconsin. After Green (1983), Rankin (1976), Rast (1989) and other sources.
The Animikie Group (in east-central Minnesota, northern Wisconsin, and Thunder Bay district, Ontario) and its correlatives in the Menominee and Baraga Groups of northern Michigan overlie Archean basement or earlier Huronian sedimentary formations. The Animikie Group consists of a deepening-shelf sequence of quartzite, banded iron-formation (BIF) (including the major Lake Superior iron ores), and slate/graywacke, which in Michigan was deposited in narrow rift-basins (Marquette, Republic troughs) and contains abundant (up to 3 km) metabasaltic flows and subvolcanic intrusions. Some of these mafic rocks (e.g., Hemlock Formation and Kiernan sills) have a distinct geochemical signature characteristic of high-Ti mid-ocean ridge basalts, contaminated by continental crust during erup-
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tion through the rifted margin (Ueng et al., 1988). A few ophiolite remnants are preserved along the suture (Schulz, 1987) but only a narrow oceanic crust appears to have been produced. Fahrig (1987) suggests that several mafic dike swarms that cut Superior Province Archean rocks to the north of this belt were associated with this incomplete cratonic rifting event. The Animikie sedimentary rocks and their correlatives thus record the development of a rifted continental margin, followed by a foreland basin during convergence (Southwick et al., 1988). Except for their northern edge in northeastern Minnesota and adjacent Ontario, the Animikie and Baraga Groups subsequently were deformed in the northward-directed folding and thrusting of the Penokean orogeny (1900-1800 Ma; Sims and Peterman, 1983; Holm et al., 1988; Morey, 1989) that welded a calc-alkaline volcanic arc (the Wisconsin Magmatic Terrane, 1890-1830 Ma) and the older Archean gneiss terrane onto the Superior craton. Svecokarelian belt
The Baltic Shield (Fig. 5 ) has undergone a similar history to that of the Laurentian shield, and plate reconstructions show these cratons to have been connected through Scotland and Greenland during the Proterozoic (e.g., Patchett and Bridgwater, 1984; Gower and Owen, 1984; Gad1 and Gorbatschev, 1987; Gorbatschev and GaA1, 1987; Pharaoh and Brewer, 1990; Park, 1991). Early Proterozoic rifting along NW to E-W trends developed about 2.3-2.0 Ga within the Archean Kola and Karelian blocks in northern Norway, Sweden and Finland, and adjacent Russia producing tholeiitic volcanism in marine and continental basins and rifts. The metabasalts are interbedded with continental and shallow marine metasedimentary rocks of the Sumian-Sariolian Group and Upper Lapponian and Jatulian sequences. In the Karasjok belt in north Finland the sequence includes basaltic and ultramafic komatiites (Saverikko et al., 1983) and in the Vetreny Poyas synclinorium, SE of the White Sea, more than 2 km of komatiites and basaltic komatiites are preserved, along with basaltic andesites, conglomerates, quartzites, and marbles in the Karelian sequence (Ryabchikov et al., 1988) (Fig. 6). The Jormua complex in northeastern Finland may be the oldest known ophiolite and has MORB geochemistry (Kontinen, 1987) (Fig. 6). Svetov (1979) and Golubev and Svetov (1983) describe and discuss the volcanology, petrology, and geochemistry of the extensive Jatulian plateau volcanism in Russian Karelia. These basalts have tholeiitic compositions (Fig. 6), including strong iron-enrichment trends. In Ostrobothnia, northwest Finland, a sequence dated at about 2.1-1.9 Ga contains abundant pillowed metabasalts and marine metasedimentary rocks. The mafic rocks are olivine and quartz tholeiites with MORB-like compositions, not continental geochemistry, and they may have been deposited in a rift marginal to the Karelian block (Honkamo, 1987). Most of the metabasalts in the region, however, exhibit within-plate tholeiitic geochemistry (Fig. 6), enriched in Fe, Ti, V, and LREE (Pharoah and Pearce, 1984; Pharaoh et al., 1987; Pharaoh and Brewer, 1990). All of these volcanic rocks thus indicate a history of continental rifting and
Proterozoic rifts
107
Fig. 5. Geologic sketch map of Baltic Shield and environs. A = Alta; B L = Bergslagen district; BP = Belomorian Province; GP = Transscandinavian granite-porphyry and rapakivi belt; H = Hedmark Group; J = Jormua complex; K = Kautokeino; KB = Karasjok Belt; LL = Lake Ladoga; 0 = Oslo graben; OB = Ostrobothnia; R = Raipas Group; S = Seiland complex; SF = Svecofennian Province; SS = Southwest Scandinavian Province; T = Telemark supracrustals; VP = Vetreny Poyas synclinorium. After Gorbatschev and Gas1 (1987) and other sources.
production of a narrow belt of new oceanic crust on the east and south of the craton. This rifting was followed by convergence, which involved subduction to the northeast under the Kola plate and more subduction, arc volcanism, arc accretion, and cratonic enlargement in the subsequent Svecokarelian orogeny (1.9-1.8 Ga). In northernmost Norway, a 200 km long, roughly N-S belt of mafic volcanic rocks (Komagfjord-Repparfjord window, Alta-Kautokeino belt) contains thick calc-alkaline volcanic rocks and volcaniclastic debris at the base of the sequence at Repparfjord, apparently representing a continental-margin arc and deep back-arc basin (Bergh and Torske, 1988). These rocks are overlain by the Raipas Group, which contains at its base a 1.5 km thick, well-preserved succession of submarine to
J. C.Green
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Fig. 6. Geochemical diagrams for rift-related Proterozoic suites of the Baltic Shield. A. Spidergram of Upper Lapponian metabasalts, from Pharaoh and Brewer, 1990. B. Ti/Cr diagram (after Pearce, 1975) for Early Proterozoic greenstones of Lapland (squares, North Noway; stars, Kiruna region, Sweden) and Jormua ophiolite complex, NE Finland (circles). From Pharaoh and Pearce (1984) and Kontinen (1987). VAB: volcanic arc basalt. C, D. Proterozoic metabasalts of Karelian Russia (Golubev and Svetov, 1983): C, diagram after Miyashiro, 1974; D, diagram after Mullen, 1983. Symbols: triangles, early Jatulian; squares, middle Jatulian; circles, late Jatulian; x , north Onega volcanic complex.
subaerial tholeiites (Kvenvik Greenstones) with MORB-like geochemistry. Though not well constrained by radiometric dating, these basalts are interpreted to have been erupted in a passive rift (aulacogen) at a high angle to the northern Karelian cratonic margin as a result of and contemporaneous with the Svecokarelian collision (Bergh and Torske, 1988). In south-central Sweden the Bergslagen area contains a Lower Proterozoic supracrustal complex, dated at 1.9 Ga, that also appears to have been deposited in a continental rift environment, perhaps in a back-arc situation (Oen, 1987; Parr and Rickard, 1987; Valbracht et al., 1991). The rocks include carbonates, ironformations, and Mn-rich sediments and other exhalite mineral deposits interpreted
Proterozoic rifts
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to be lacustrine, along with ‘abundant bimodal volcanic rocks remarkable for the great dominance of felsic compositions (“leptites” and “halleflintas”) and their volcaniclastic debris. They are intruded by essentially coeval granitic rocks with alkaline to subalkaline compositions and trace-element enrichments characteristic of anorogenic suites, The mafic volcanic rocks are continental tholeiites (Valbracht et al., 1991). Mafic rocks from the Bothnian Basin, just to the northeast of the Bergslagen district, show slightly positive ENd values, implying a nearly undepleted mantle source (Claesson, 1987). Park (1991) gives a more detailed interpretation of the complex Proterozoic tectonic history in Fennoscandia. To the west of the Svecokarelian belt, certain rocks in northwestern Scotland may represent a tectonic connection to the Circum-Superior belt. The Scourie dikes (2.4-2.2 Ga), of continental basalt composition (Weaver and Brney, 1981), cut the Archean basement and are analogous to the coeval mafic intrusions of the northern Baltic shield. The Loch Maree Group is a fault-bounded sequence in the Lewisian complex in Scotland (Johnson et al., 1987) (Fig. 17) with a Sm/Nd model age of 2.0 Ga (O’Nions et al., 1983). It consists of amphibolites and metasedimentary rocks of an original sequence at least 3 km thick. The metasedimentary rocks were dominantly graywacke, much of it volcaniclastic, with some chemical shelf deposits (marble, iron-formation). The meta-igneous rocks include MORB-type and continental tholeiitic flows and sills, and mafic intrusions that show some evidence of crustal assimilation. Johnson et al. (1987) infer that these rocks record an episode of continental rifting during dextral transtension. They were subsequently deformed and metamorphosed during the Laxfordian orogeny (1.8-1.6 Ga). In southern Greenland, at least four swarms of tholeiitic dikes of similar age to the Loch Maree Group (2.15-1.95 Ga) may represent a further connection between rifting in the Baltic and Laurentian cratons (Hall and Hughes, 1988).
Wopntny orogen A major rifted continental margin developed along the western part of the Archean Slave craton in northwestern Canada at about 2.0-1.9 Ga (Hoffman, 1980, 1987, 1988) (Figs. 2, 7). The resulting Coronation Supergroup includes initial rift products of the Akaitcho Group, overlain by slope-rise and shelf clastics and carbonates (Epworth Group), in turn overlain by foredeep turbidites and shales (Recluse Group) as the continental margin collapsed. All of these rocks subsequently became involved in eastward thrusting as the Great Bear volcanic arc became accreted to the Slave craton during the Hudsonian orogeny. The Akaitcho Group (Easton, 1980, 1981) contains up to 10 km of rift-fill clastic and volcanic rocks (Fig. 7). The basal portion includes a t least 500 m of pillowed and massive metabasalt and tuff, interfingering upward with pelites. These are overlain by 3 to 4 km of arkosic and subarkosic turbidites intruded by rhyolitic sills, which are in turn overlain by over 2 km of bimodal volcanic rocks and conglomerates. The upper 3-4 km of the Akaitcho Group consists of pelitic
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Fig. 7 . Cross-section model for the depositional framework of the Coronation Supergroup, Wopmay orogen, looking north. After Hoffman (1987). Akaitcho Group filled initial cratonic rifts.
rocks, bimodal tuffs, and volcaniclastic rocks, with local submarine metabasalt sequences up to 400 m thick. All are intruded by gabbroic sills. Although there is some geochemical overlap with arc tholeiites, the lower basalts, erupted during rift initiation, resemble continental tholeiites with moderately enriched LREE's (Ce/Yb = 10-15); as the crust thinned the later basalts show more marginal-basin character with less-enriched LREE's (Ce/Yb = 2-5). The rhyolites, which show more steeply sloping R E E patterns (Ce/Yb = 20-25) are thought to represent crustal melts (Easton, 1981).
EARLY TO MIDDLE PROTEROZOIC RIFTS: 2.0-1.5 Ga
Canadian Shield sequences Several rift-related volcano-sedimentary sequences were emplaced in separate parts of the Canadian Shield during the period 1.9-1.76 Ga, during or immediately following the Hudsonian orogeny. Among these are the Dubawnt Group northwest of Hudson Bay and the Aillik Group in central Labrador (Fig. 2). The Dubawnt Group in the District of Keewatin, about 1850-1760 Ma, is a sequence of continental sedimentary and volcanic rocks that occupy NE-trending depressions within the Archean Rae Province (Blake, 1980; LeCheminant e t al., 1987; Hoffman, 1988). At its base near Baker Lake it contains redbeds in the Kazan and South Channel Formations. Overlying these are the volcanic Kunwak
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and Christopher Island Formations, which extend for at least 400 km along strike. These are overlain by quartzose sediments and rhyolites, and are intruded by comagmatic high-F granites about 1.76 Ga old. The Christopher Island volcanic rocks are a strongly alkaline suite of trachyandesite and trachyte subaerial lavas and pyroclastic rocks and are accompanied by an even more extensive suite of alkaline stocks and NE- to N-trending dikes composed of syenite and lamprophyre. Geochemically these alkaline rocks are highly enriched in K/Na and LREE (ave. L a m b = 69) as well as in ,'F F, Ba and other incompatible elements except Ti. LeCheminant et al. (1987) attribute this magmatic event to melting of enriched mantle (apparently in a plume) with associated continental uplift and rifting. A tectonic explanation is as yet elusive (Hoffman, 1988). The Makkovik orogen in central Labrador is an Archean to Lower Proterozoic terrane that lies between the Archean Nain Province to the northwest and the Middle Proterozoic Grenville Province to the south (Gower and Owen, 1984; Scharer and Gower, 1988). In pre-Atlantic reconstructions it appears to correlate with the Ketilidian mobile belt of southern Greenland and foreland zones of the Svecononvegian orogen in southern Scandinavia. Within this mobile belt are remnants of a supracrustal sequence of volcanic and sedimentary rocks (Aillik Group, Moran Lake Group) that appear to have been deposited in an incipient continental rift that developed at 1.86 Ga before the onset of the 1.8 Ga Makkovikian orogeny (Sharer and Gower, 1988). Volcanism was bimodal, and included submarine metabasaltic eruptions early in the cycle and more voluminous, younger, subaerial felsic volcanism accompanied by continental and shallow-shelf sedimentation. The youngest rhyolites appear to be comagmatic and coeval with the nearby 1.80-Ga granites. Geochemical data are difficult to interpret because of metasomatism (Gower and Ryan, 1987).
Southern African basins Situated in the northernmost part of the Transvaal, South Africa (Fig. l), the Soutpansberg Group is a sequence of mafic volcanic and clastic rocks as much as 7 km thick (Jansen, 1975; Bnkard et al., 1982; Crow and Condie, 1990). It occupies an E-W trough about 40 km wide and 300 km long (450 km long including outliers to the west in Botswana), and developed by infilling of a gradually subsiding, graben-like basin that cut across the structural grain of the underlying Lower Proterozoic Limpopo mobile belt (Fig. 8). The rocks have undergone only mild burial metamorphism, and have been dated by Rb/Sr at about 1770-1750 Ma (Barton et al., 1979). The lower part of the Soutpansberg Group (Sabasa Fm) is primarily subaerial basalts; higher in the section mature, fluvial, quartzose sandstones and argillites predominate. There are a few mafic and felsic pyroclastic beds, but no significant intrusions. Rifting subsidence began at the east end and progressed westward and also to the north, as indicated by onlapping upper stratigraphic units. Many faults were reactivated during and after Karoo times in the Jurassic (Jansen, 1975). The volcanic rocks are continental
J.C. Green
112
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Fig. 8. Model for rift evolution of the Soutpansberg Trough, northern Transvaal. From Tankard et al. (1982), after Jansen (1975).
tholeiites geochemically (Barton et al., 1979; Crow and Condie, 1990), with moderately enriched Fe/Mg and LREE, but some suggestion of subduction-related enrichment in their mantle source. The Waterberg Group is another thick sequence of mostly subaerial redbeds that are preserved in two basins (Warmbaths, Middleberg) in the northwestern and central Transvaal Province (Tankard et al., 1982). They overlie Transvaal Supergroup and older rocks (Fig. 1). The Waterberg Group may have covered an area approximately 300 x 300 km originally, and a section about 7 km thick is preserved (the top is erosional). There is a small amount of subaerial felsic volcanic rock a t the base, but the remainder is red, fluvial sandstone and conglomerate with
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minor siltstone, shale, and shallow tidal-shelf deposits. Although the Waterberg sequence is approximately coeval with the nearby Soutpansberg Group, it does not show the latter’s clear evidence for active faulting of the basement associated with the subsidence. Australian basins The three Archean cratons (Yilgarn, Pilbara, Gawler blocks) in Australia are associated with several irregularly shaped inliers of Proterozoic igneous and metamorphic rocks, all separated and covered to varying degrees by younger Proterozoic and Phanerozoic sedimentary rocks (Fig. 9). Many of these Proterozoic inliers contain evidence of localized crustal rifting, followed soon by compression. The separate sequences are in general not aligned with each other, nor can strata be correlated between them, and they cannot be interpreted as one or more coherent, linear continental rifts. Although evidently formed as separate rift basins, their history has been similar (Wyborn et al., 1987; Etheridge et al., 1987). Local stretching and thinning of the Archean basement began sometime in the period 2000-1880 Ma, with deposition of quartzose, fluvial sediments and eruption of bimodal volcanic rocks (mainly mafic) and intrusion of large dolerite
Fig. 9. Archean cratons and rift-related Proterozoic terranes in Australia. Cratons (shaded): PB = Pilbara Block; GC = Gawler Craton; YB = Yilgarn Block; pre-1870 Ma basins: AB = Arunta Block; AT = Ashburton Trough; HC = Halls Creek inlier; MI = Mt. Isa Inlier; PC = Pine Creek inlier; H B = Willyama Block. AG = Late Proterozoic Adelaide Geosyncline and Flinders Range; dashed line, western limit of Paleozoic Tasman geosyncline/fold belt. From several sources.
114
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sills and layered mafic intrusions. The mafic rocks are geochemically similar to both continental tholeiites and MORBs, suggesting considerable thinning of the crust (Wyborn et al., 1987). This intense stretching was followed by more gentle subsidence, with deposition of shallow-shelf sediments which include pelites, carbonates, and iron-formations. More active subsidence followed, with deposition of turbidites in what may have been foreland basins related to the immediately succeeding Barramundi orogeny at 1870 f 20 Ma. This deformation produced local fold and thrust belts which, like the supracrustal sequences, cannot be traced between the separate Proterozoic domains. The Barramundi orogeny was characterized by high-temperature, low-pressure metamorphism, felsic calc-alkaline volcanism, and intrusion of voluminous synorogenic, I-type granites (1870-1840 Ma). Etheridge et al. (1987) attribute this tectonic cycle to crustal underplating (about 2.2-1.9 Ga) from mantle plumes, subsequent crustal attenuation and local basin formation, then later delamination which led to extensive crustal melting associated with the compression (Fig. 10). McCulloch, (1987) however, concludes, from a study of Sm/Nd isotopes, associated U/Pb analyses, and other arguments, that there is very little evidence for the involvement of Archean sialic crust in this voluminous magmatic episode at 1.87 Ga. In fact, more than one-third of the Australian crust appears to have been formed between about 2.2 and 1.8 Ga. McCulloch infers that initial 2.2-Ga plumes could have underplated mafic (oceanic) crust; the underplated material in turn became the source for Barramundian anatectic melting at 1.87 Ga. Thus, these Lower Proterozoic sequences may not represent rifting of sialic Archean basement. Another problem with the multiple-plume model for these Australian rift basins (e.g., Etheridge et al., 1987) lies in the size of the plumes themselves. According to the model (Fig. 10) each basin (now roughly 200-500 km across) overlies a single plume. These basins would be much smaller than the plume heads in the models of Courtney and White (1986), White and McKenzie (1989), and Campbell and Griffiths (1990). Different assumptions as to Early Proterozoic mantle rheology and/or plume potential temperature might help to reconcile these models. A second, extended period of crustal formation in the Early-to-Middle Proterozoic (1820-1670 Ma) also apparently involved local crustal rifting (Wyborn et al., 1987). The Willyama Complex (Supergroup) in New South Wales and South Australia (James et al., 1987), dated at about 1820 Ma and metamorphosed a t 1660 Ma, begins with immature, terrigenous, possibly volcaniclastic sediments (now gneisses), interpreted as rift-basin fill. These are overlain by a bimodal volcanic suite interbedded with semipelitic metasandstones and pelites of a deepening shelf environment, and are succeeded by continental-slope and abyssal pelites without volcanic rocks. The amphibolites are inferred to be partly lavas and tuffs, and partly dikes and sills (Stroud et al., 1983), whereas the felsic metavolcanic gneisses are dacitic and rhyolitic in composition and are thought to represent lavas and/or ash-flow tuffs. The mafic rocks are tholeiitic, enriched in Fe and Ti. These rocks have been interpreted as forming either in a strongly attenuated crustal setting or a rifted margin.
115
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Fig. 10. Model for evolution of Early Proterozoic intracratonic basins, Australia. After Etheridge et al. (1987). Cratonic extension and thinning are driven by mantle upwelling; subsequent orogeny results from delamination and sinking of lithosphere and rise of hot asthenosphere.
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Evidence of rift volcanism is also found in the Mt. Isa inlier in north-central Australia (Wilson, 1987). Here the Eastern Creek Volcanics of the Harlingden Group (about 1750 Ma) and Soldiers Gap Group (probably of similar age) contain tholeiitic mafic metavolcanic rocks with MORB-like compositions. The Eastern Creek Volcanics also include some continental tholeiites. Several tholeiitic dike swarms of mid-Proterozoic age also intrude the Australian basement, indicating incipient rifting (Wyborn et al., 1987; Parker et al., 1987; Parker, 1988).
MIDDLE PROTEROZOIC RIFTS: 1.5-1.0 Ga
North American Midcontinent RiftSystem Although it is covered by Paleozoic and Pleistocene deposits over most of its length, the Midcontinent Rift System (MRS) is one of the world’s largest and best-preserved continental rifts (Wold and Hinze, 1982; Van Schmus and Hinze, 1985; Cannon et al., 1989). The MRS forms a 2000 km long loop that extends from northeastern Kansas ENE through Iowa and eastern Minnesota, eastward through the Lake Superior area, and south and southeast through Michigan approximately to the western end of Lake Erie (Fig. 4). The principal exposures are along the shorelands of Lake Superior, to the northwest at Lake Nipigon, Ontario, and to the southwest in east-central Minnesota and adjacent Wisconsin. Elsewhere it is recognized by its gravity and magnetic signatures (among the largest anomalies and gradients in North America) and from several seismic reflection and refraction studies (e.g. King and Zietz, 1971; Ocola and Meyer, 1973; Oray et al., 1973; Serpa et al., 1984; McSwiggen et al., 1987; Chandler et al., 1989; Cannon et al., 1989; Hinze et al., 1990). The crust is anomalously thick (up to 58 km) along the MRS under central Lake Superior (Behrendt et al., 1990). The MRS tapped enormous volumes of mantle melts (estimated at 1.3 x lo6 km3; Hutchinson et al., 1990) and was active magmatically over a limited period, between 1109 and 1086 Ma (U/Pb; Davis and Sutcliffe, 1985; Davis and Paces, 1990; Palmer and Davis, 1987). Franklin et al. (1980) proposed that the Sibley sedimentary basin to the northwest of Lake Superior was an aulacogen related to the initiation of the MRS, but the Rb/Sr age of the Sibley Group (1340 Ma) appears much too old to be related to the Midcontinent Rift. The stratified rocks of the MRS are known as the Keweenawan Supergroup. Stratigraphic and structural evidence from the Keweenawan exposures show that rifting produced several successive accumulations of subaerial plateau lavas, each several km thick, centered over different segments of the rift (BVSF’, 1981; Green, 1977, 1982,1983). Each plateau subsided centrally. Multichannel shipboard seismic profiling (e.g., Behrendt et al., 1988; Cannon et al., 1989) has recently shown that major half-grabens developed along the rift axis, in addition to the warping subsidence, and that up to 20 km or more of volcanic rocks accumulated in these
Proterozoic rifts
117 100
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160 K M
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Fig. 11. Interpreted, migrated seismic reflection profile (line F, roughly N-S) across Midcontinent Rift, east-central Lake Superior. Random dash pattern, Archean and Lower Proterozoic basement; M , seismic Moho. Shaded zone, inferred volcanic rocks; unpatterned area above volcanics represents clastic sedimentary rift fill. Heavy lines show inferred faults. Vertical exaggeration 1.1x assuming seismic velocity of 6 km/s. After Behrendt et al. (1988).
basins. The rifting appears to have been segmented by accommodation zones, with adjacent half-grabens bounded by normal faults of alternating polarity, but some graben segments are symmetrical (Cannon et al., 1989; Dickas and Mudrey, 1989) (Fig. 11). These normal faults are probably rooted in the ductile lower crust, which was highly attenuated; only approximately 10-14 km of older crust is preserved above the seismic Moho under Lake Superior (Behrendt et al., 1988), and some of this appears to be intruded and underplated by additional mafic igneous rocks (Behrendt et al., 1990). Along strike to the southwest, buried rift segments are offset by transform-like faults (Chase and Gilmer, 1973). No new ocean was produced in this aborted rift, but gravity and seismic refraction studies show that anomalously dense and high-velocity rocks (probably abundant mafic intrusions) underlie the rift axis. At about 1094 Ma the active rifting and volcanism began to wane, but subsidence continued. Up to 10 km of mostly red, fluvial and deltaic and gray lacustrine sediments of successively increasing maturity and wider provenance were deposited over the volcanic rocks (Ojakangas and Morey, 1982). The lower sequence of redbeds (Oronto Group) is preserved on top of a large horst that trends along the rift axis from the Lake Superior Basin southwestward into Iowa. This horst resulted from a late compressional phase of the MRS that caused high-angle reverse movement, apparently in part at least along the older, grabenforming normal faults. The upper clastic sequence (Bayfield Group, Jacobsville Sandstone, and subsurface equivalents) is subhorizontal and fills marginal basins along the flanks of the horst as well as the central Lake Superior Basin. The volcanic rocks constitute a subaerial, bimodal, tholeiitic suite, dominated by olivine tholeiites along with abundant transitional basalts and basaltic andesites (e.g., BVSP, 1981; Green, 1982, 1983; Brannon, 1984; Paces, 1988; Lightfoot et
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MG"
0.4
Fig. 12. Geochemical diagrams for volcanic rocks of Midcontinent Rift System. A. AlzOs/Mg# for North Shore Volcanic Group, showing high-Al character of more primitive basalts. R = rhyolites; I = icelandites and ferroandesites; BA = basaltic andesites; TB = transitional basalts; OT = olivine tholeiites. B. REE diagram for representative mafic and intermediate lavas, NSVG, from BVSP (1981). C. Spidergram for average high- and low-Ti basalts of Portage Lake Volcanics, Michigan (Nicholson and Shirey, 1990). D. AFM diagram for mafic and intermediate rocks. Symbols: triangles = Portage Lake Volcanics, Michigan (Nicholson and Shirey, 1990); circles = representative samples, Mamainse Point Volcanics, eastern Lake Superior (Klewin and Berg, 1991); x = Keweenawan reference suite (NSVG), BVSP (1981). Some samples fall in calc-alkali field because of metasomatism related to burial metamorphism.
al., 1991; Klewin and Berg, 1991; Fig. 12). Rhyolites are relatively uncommon except in northeastern Minnesota (the North Shore Volcanic Group, NSVG) where they constitute between 10% and 25% of the section and were erupted mostly as very large, hot lava flows and rheoignimbrites (Green and Fitz, 1992). The basalts and interflow conglomerates have been a world-class source of native copper. Several mafic dike swarms, roughly parallel to the trend of the rift, cut the older lavas and the surrounding pre-rift basement (Green et al., 1987). Plutonic
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119
rocks, notably the 210 km long Duluth Complex, make up a large proportion of the Keweenawan igneous suite in northeastern Minnesota. The Cu-Ni bearing Duluth Complex, which generally underlies and intrudes the NSVG, comprises many separate plutons of various sizes and shapes. An earlier major phase (the “anorthositic series”) was intruded as a crystal mush (Miller and Weiblen, 1990); this was followed by a later group of layered, olivine-rich cumulate rocks (the “troctolitic series”) and several large granophyre bodies (Weiblen, 1982). Evidence from chilled margins and cumulus mineral compositions implies that the Duluth Complex magmas were moderately evolved and were derived from differentiating chambers much deeper in the crust. Similar but smaller cumulate plutons across the MRS in Wisconsin are known as the Mellen Complex (Olmsted, 1968; Klewin, 1990). Higher-level cumulate gabbro to ferrodiorite plutons and large diabase sills intrude higher levels of the NSVG (Green, 1972; Miller, 1987) and Lower Proterozoic (Animikie) strata in the Thunder Bay-Lake Nipigon area, Ontario (“Logan sills”; Sutdiffe, 1987). Geochemically the Keweenawan olivine tholeiites, which are generally aphyric and Al-rich, resemble E-MORB and Icelandic basalts. On discriminant diagrams the lavas plot mainly in the within-plate, ocean-island, or mid-ocean rift fields (Fig. 12). Rare-earth element patterns are undepleted, with approximately level chondrite-normalized trends for the most primitive olivine tholeiites (Mg# 6672) and increasing in R E E content and La/Yb to the Fe-enriched andesites, icelandites, and rhyolites. Incompatible element ratios (Th/Tm = 2.12 - 2.16; ZrfY = 4.3 - 4.4) and initial isotope ratios at 1.1 Ga ( 8 7 ~ r / 8 6 ~ =r 0.7038; i ENd = o f 2; pI = 8.2) of the main basaltic units imply magma origin from a large, enriched, asthenospheric mantle plume centered beneath Lake Superior (Massey, 1983; Nicholson and Shirey, 1990; Hutchinson et al., 1990). These magmas subsequently evolved by a combination of processes including fractionation in deep and shallow crustal chambers, replenishment and mixing, and only minor or local crustal assimilation (Brannon, 1984; Sutcliffe, 1987; Paces, 1988; Klewin and Berg, 1991; Lightfoot et al., 1991). The rhyolites that have been studied reflect considerable partial melting/assimilation of Archean crust or of earlier MRS basalts (Nicholson and Shirey, 1990). Gordon and Hempton (19%) have proposed that the MRS developed as a series of pull-apart basins, related to a set of NW-trending sinistral strikeslip faults in the craton, that resulted from the contemporaneous Grenville continental collision to the east (Fig. 4). Several lines of evidence suggest otherwise, however, as outlined here: (a) Both limbs of the MRS are nearly parallel to the Grenville Front as it is traced from Lake Huron through Michigan, Ohio, and Kentucky (Green et al., 1988; Pratt et al., 1989), whereas most of the major collision-generated grabens used for analogy by Gordon and Hempton (1986) (including the Alpine, Himalayan, and VariscadHercynian collisions) are more nearly perpendicular to the related deformation fronts. (b) Although there appear to be transform faults separating several of the Keweenawan lava basins, there is no regional system of NW-trending sinistral strike-slip faults documented
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J. C. Green
that might be construed as related to the Grenville collision front. (c) Keweenawan diabase dike swarms in the Lower Proterozoic and Archean basement outside of the rift are generally parallel to the rift, including its inverted-U bend in the Lake Superior area (Green et al., 1987); they d o not show any stress orientation related to the Grenville Front. (d) No early marginal faults have been found along the MRS; the first Keweenawan, rift-related deposits are either thin quartz-arenites or mafic lava, not fanglomerate derived from basement rocks. (e) The volume of magma involved is much greater than in most passive or collision-related rift basins. (9Both geophysical (Hutchinson et al., 1990) and geochemical (Nicholson and Shirey, 1990) evidence indicate the presence of a superheated mantle plume beneath Lake Superior operating from the beginning of magmatism. Although these considerations make it unlikely that the Grenville collision caused the Keweenawan rift, it might be argued that the collision was responsible in some way for the compression in late Keweenawan time that produced the central horst along much of the MRS, after magmatism had ceased. Further definitive radiometric ages for Grenville events will help to resolve this uncertainty.
Belt-Purcell- Wernecke basins Along the western margin of the Lower-Middle Proterozoic Laurentian continent large, aulacogenic, evidently fault-controlled basins developed on stabilized Hudsonian basement and on or outboard of the tectonically thinned (rifted) continental margin (Price, 1984). Thick successions of deep-water turbidites accumulated, eventually succeeded by thick cycles of shallow-water marine clastic and carbonate sediments as the subsidence became less intense. The main depocenter was a NW-trending trough in Montana, Idaho, northeastern Washington, and British Columbia (Fig. 13), where the strata are known as the Belt (U.S.), the Purcell (Canada), and to the northwest in the Yukon and North West Territories (N.W.T.) the Wernecke Supergroups (Stewart, 1972; McMechan, 1981; McMechan and Price, 1982; HOy, 1982; Young, 1984). The east edge of the Belt Basin was probably as far east as central Montana and southwestern Alberta, and the southwestern margin is unknown (Reynolds, 1984). Delaney (1981) correlates the Wernecke Supergroup with the lower three groups of the Belt/Purcell. The timing of the onset of this rifting subsidence is not well constrained, being estimated by various authors at about 1500 Ma; a gabbroic sill intruded well above the base is datcd at 1433 f 13 Ma (Zartman et al., 1982). On the basis of similar geology and truncated structural grain, Sears and Price (1978) suggest that this cratonic rifting was associated with the breaking away of a Siberian craton from the western margin of Laurentia at about 1.5 Ga. Sediment thicknesses of up to 12 km or more (e.g., Wcrnecke Supergroup, Wernecke Mts: Delaney, 1981; 10 to 20 km, Purcell: Price, 1984) accumulated before the end of this cycle. The Purcell Supergroup contains a relatively thin sequence (up to 400 m) of basalt to basaltic andesite lavas (Purcell lavas, Nichol Creek Fm; McMechan 1981; HOy, 1982, 1984) in the upper-central part of the sequence along the northeastern
Proterozoic rifts
121
Fig. 13. Sketch map of North American cordilleran area showing localities of Proterozoic rift-related sequences. Dashed line: approximate western limit of mid-Proterozoic craton. T = Tindir Group; 0 = Ogilvie Mountains; W = Wernecke Mtns; MM = Mackenzie Mtns; BP = Belt-Purcell basin; U = Uinta Mtn. Group; GC = Grand Canyon. States and provinces: AK = Alaska; A1 = Alberta; A 2 = Arizona; BC = British Columbia; CA = California; ID = Idaho; Mi" = Montana; NUT = Northwest Territories; UT = Utah; WA = Washington. After Harrison et al. (1974) and other sources.
portion of its extent, and one of the last events in the cycle was a renewal of rifting on the adjacent platform at about 1220 Ma, with deposition of the >4 km thick Coppermine River Group in the NWT. This sequence consists of continental tholeiitic plateau basalts and redbeds, and may be correlated with the nearby tholeiitic Musk Ox intrusion, the Hart River volcanic rocks in the Wernecke Mountains (Bell and Jefferson, 1987) and the widespread Mackenzie dike swarm (Gibson et al., 1987). Geochemistry of the Coppermine River basalts (Dostal et al., 1983) indicates a source similar to that of P-MORB, with some
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. C. IGreen
contamination as well as considerable crystal fractionation within the crust before eruption. The Wernecke Supergroup was disrupted by large megabreccias, and eventually was deformed and locally metamorphosed in the Racklan orogeny at about 1200 Ma (Young, 1984). McMechan and Price (1982) have suggested that the East Kootenay orogeny, which terminated deposition in the Belt/Purcell Basin, can be correlated with the Racklan orogeny, but definitive evidence is lacking. Obradovich et al. (1984) believe that K/Ar and Rb/Sr ages of sedimentary rocks of the Missoula Group (upper part of the Belt) show it to be as young as 1100-900 Ma, in disagreement with paleomagnetic studies of Elston and Bressler (Elston, 1984). Riphean aulacogens of the former US.S.R. Several elongate rifts developed in the East European and Siberian cratons during the Riphean (Middle to Late Proterozoic) (Zonenshain et al., 1990). Flood basalts as well as clastic rift fill are found in most of them. In the East European Platform the rifts form a roughly rectangular pattern trending NE and NW. Where some of these reached the cratonic margin, they may have been the failed arms of triple-junction systems which developed into ocean basins. Examples include the northeastern terminus of the Sredne-Russky aulacogen and the southwestern teiminus of the Volyno-Orshano-Krestsovsky rift. Other grabens of this system (e.g. the Kandalasksha and Mezen grabens) are parallel to the rifted catonic border. The closure of the eastern ocean during Paleozoic time was responsible for construction of the Urals orogen. In Siberia, Riphean rifting is exemplified by the Udzh aulacogen, which lies between the Oleniok uplift and the Anabar massif (Zonenshain et al., 1990). This rift involved considerable thinning of the continental crust and deposition of 7 to 9 km of sediment. Alkaline-ultramafic magmatism, including kimberlites, was also associated with this aulacogen (Trushkov et al., 1974). This period of cratonic rifting may have been related to the breaking away of Siberia from western Laurentia and the development of the Belt-Purcell-Wernecke aulacogens of that continent (Sears and Price, 1978). Grenville and Teleniark supracrustals At about 1200 Ma, a series of metasedimentary basins developed on older Proterozoic crust in both the Grenville Province (eastern Canada) and Telemark (southern Norway) (Figs. 4, 5). The Telemark supracrustal rocks are mainly sandstones and conglomerates, now in greenschist facies; the Grenville metasedimentary rocks include amphibolite-grade shallow-marine carbonates as well as psammitic and pelitic rocks. This sedimentation was accompanied by the intrusion of mafic dikes and sills with a wide range of geochemical character, including continental tholeiites and calc-alkaline suites (Smith and Holm, 1990; Brewer and Atkin, 1987). The N-S trending “Protogine Zone” dolerites of southern Sweden
Proterozoic rifts
123
(Johansson and Johansson, 1990) appear to be part of this rift-related magmatic event. After clockwise rotation of the Baltic shield with respect to Laurentia, these supracrustal sequences were involved in a major collisional orogeny at about 11001050 Ma (Grenville in North America, Sveconorwegian in the Baltic Shield). The sedimentary accumulations are interpreted as rift-basin fill, accompanied by mafic intrusions injected into an attenuated crust in a back-arc environment, and an adjacent continental arc associated with the collision (Falkum and Petersen, 1980; Smith and Holm, 1990; Brewer and Atkin, 1987; Gorbatschev and GaB1, 1987). Kalahari copper belt A northeasterly-trending rift system developed along what is now the western and northern margins of the Kalahari craton of southern Africa (Fig. 14) during the Late-Middle Proterozoic (Borg and Maiden, 1987; Borg, 1988). Available age
Medit. Sea
Fig. 14. Sketch showing Proterozoic rift-related features of Africa and adjacent South America. BA = Bou Azzer; BB = Bangwelulu Block; BNS = Benin-Nigeria shield; C = Chad craton; CC = Congo craton; D = Dahomeyan belt; DB = Damara belt; G = Gariep Province; GT = Gourma trough; K = Kalahari copper belt; KB = Kibaran belt; KC = Kalahari craton; PB = Pharusian belt; RB = Ribiera belt; SFC = SBo Francisco craton; TC = Bnzania craton; TS = Tuareg shield; VB = Volta Basin; WAC = West Africa craton. After many sources.
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determinations suggest that rifting and sedimentation may have lasted from 1300 to 950 Ma (SACS, 1980). Products of this rifting are contained in an 1800 km long series of block-faulted basins in Lower-Middle Proterozoic basement, trending from central Namibia through northern Botswana into northern Zimbabwe. The general sequence in these basins contains a thick basal accumulation of high-K rhyolites, including pyroclastic rocks, with minor coarse clastic sediments. Active block-faulting then produced a rift-fill sequence of up to 3 km of continental redbeds and interbedded basalts (as much as 2.5 km locally). These have associated native Cu mineralization. The basalts and redbeds are overlain by shallow-marine and possibly lacustrine sediments with stratabound Cu in the argillites. All of these rocks were subsequently metamorphosed to greenschist facies in the latest Proterozoic Damara orogeny. The basalts are tholeiites having within-plate geochemistry (Borg and Maiden, 1987; Borg, 1988). Kibaran belt
The Kibaran belt of eastern and central Africa trends in a NNE direction between the Congo craton on the west and the Bangewelulu block and nnzania craton on the east (Fig. 14). It is one of the many Proterozoic mobile belts that separate several stable Archean cratons of Africa, and their tectonic origin has been a matter of considerable controversy (e.g. Kroner, 1977). In their analysis of the Burundian portion of the Kibaran belt, Klerkx et al. (1987) propose that crustal extension, related to voluminous (plume-generated?) mafic intrusions in the lower crust, was the initiating phase of the Kibaran orogenic cycle. Evidence for rifting includes a thick sequence (11-14 km) of clastic metasedimentary rocks and bimodal magmatism. The sedimentary rocks (Burundi Supergroup) consist mainly of quartzitic and pelitic rocks, with a concentration of poorly sorted and immature deposits, including conglomerates with angular fragments of underlying sedimentary deposits in the upper division. This is interpreted to imply active rifting producing fault scarps at this stage. Mafic and minor felsic volcanic rocks are widespread in the upper-middle portion of the sequence, especially in western Burundi, but volcanic rocks are not abundant. Granitic intrusions and mafic and ultramafic plutons (especially in east-central Burundi and trending north to Lake Victoria in Tanzania) were emplaced into the deeper, more ductile portion of this sedimentary pile during extension from about 1330 to 1250 Ma. Subhorizontal ductile shear zones and foliations, with which the granites are intimately associated, are interpreted to be related to detachment/decollement as the lower crust was being attenuated. The granitic rocks are strongly peraluminous, which is normally suggestive of collisional tectonics, but Klerkx et al. (1987) attribute this to assimilation of pelitic Burundi metasedimentary rocks as the anatectic magmas rose in the middle crust. By about 1200 Ma extension had ceased and the major compressional phase began. Open, upright folding accompanied by additional peraluminous granitic magmatism was followed by late-tectonic shearing and post-tectonic bimodal intrusive magmatism at about 1100 Ma.
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MIDDLE TO LATE PROTEROZOIC RIFTS: 1.0-0.6 Ga
Damara-Ribiera Province In western South Africa, Namibia, southwestern Angola, and Brazil, thick and highly deformed shelf and slope/trough metasedimentary rocks of the Damara orogen appear to record a major episode of continental rifting followed by convergence (Porada, 1979; Kroner, 1980; Barnes and Sawyer, 1980; Zinkard e t al., 1982; Miller, 1983b; Bernasconi, 1987; Breitkopf and Maiden, 1987). Several deep grabens initially developed in a triple-junction pattern at about 1.0 to 0.8 Ga, in the then contiguous Brazilian (SBo Francisco), Congo, and Kalahari cratons. Two branches, in total about 1200 km long, were approximately coincident with the modern continental margins (Northern Coastal Branch; Gariep Province), and the third trends ENE through Namibia (Damara Belt Central Zone) (Fig. 14, 15). Although this rifting was possibly initiated by spreading over a mantle plume, only minor volcanism accompanied the rifting. The early Nosib grabens were filled with up to 6 km of almost exclusively coarse clastic sediments (up to 5 km in the northern branch), with some carbonates and evaporites. Only a few late rhyolites were erupted, some of which are peralkaline, though more andesites, rhyolites, and quartz latites are included in the Gariep Province graben (Stinkfontein Fm.). A major N-S mafic dike swarm (Gannakouriep), implying E-W extension, cut the adjacent western Kalahari craton. Henry et al. (1990) suggest that the strong asymmetry of the inland branch of the orogen is the result of the involvement of large, NW-dipping detachment faults during the early rifting (Fig. 15). At about 830 Ma the rifting stage evolved into a broader subsidence, producing shelf sedimentation (Otavi Group) on the margin of the Congo craton and in the
CONGOCRATON Northern Zone
KALAHARI CRATON Central Zone
Southern Zone
Southern Margin Zone
? INW
, Sea
SE
7 CONTINENTAL CRUST
Fig. 15. TWOmodels for early rifting stage of Damara orogen, Namibia, looking northeast: A, after Miller (1983b); B, after Henry et al. (1990).
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Gariep Province, and very thick flysch (deepwater fans, Swakop Group) in the Central Damara Belt and Gariep Province. Mafic volcanism is represented in the Central Damara Basin by the locally pillowed Matchless belt amphibolites (350 km long, up to 3 km thick) and in the Southern Marginal Zone by the metabasalt flows and sills of the Chuos Formation and Vaalgras Subgroup (at least 400 km along strike). The Grootderm Formation, an assemblage of 4 to 5 km of mafic volcanic rocks possibly representing ocean crust, accompanies the submarine-fan turbidites in the Gariep Province. The Brazilian0 cycle in the Ribiera belt also contains coarse clastics and cross-bedded quartzites in its early, N-S rift basin, overlain by carbonates and rythmites interbedded with metabasaltic amphibolites. All of these rocks were subsequently intruded by calc-alkaline granitoid rocks, and strongly deformed and metamorphosed in the Damaran orogeny at roughly 660-550 Ma. Thrusting was principally to the southeast in the Damara belt but to the east along coastal Africa. The Chuos metavolcanics show major- and trace-element signatures of withinplate basalts, whereas the Matchless belt amphibolites have regional geochemical differences that include both ocean ridgeMoor and continental flood basalt types (Breitkopf and Maiden, 1987; Miller, 1983a). These, and a belt of serpentinites in the Southern Margin Zone that have geochemical characteristics of alpine ultramafics (Barnes, 1983) suggest that oceanic crust of some uncertain width was probably produced. Current models (e.g., Xinkard et al., 1982; Miller, 1983b) suggest that the original rift-forming mantle plume evolved to delaminate the crust from mantle lithosphere, producing crustal heating that resulted in ductile spreading, subsidence, and eventual continental separation. Kukla and Stanistreet (1991) interpret the thick pelitic rocks (Kuiseb Fm, Swakop Group) as an accretionary prism overlying ocean crust, supporting the full-ocean model. Recent paleomagnetic analysis (Renne et al., 1990) also suggests that there was an oceanic separation between the Kalahari and Congo cratons (Adamstor Ocean, Hartnady et al., 1985), which was closed by oblique convergence during the Damara orogeny. However, Rb/Sr isotopic analyses (Hawkesworth et al., 1983) of Damaran orogenic plutonic rocks indicate little if any separation between the Kalahari and Congo cratons. This problem currently remains unresolved.
Dahomeyan-Pharusian (Dans-Sahara) belt Extending roughly N-S through Algeria, Niger, Mali, and Nigeria in northwestern Africa is the Dahomeyan-Pharusian belt, another mobile zone activated during the broadly defined Pan-African orogenic episode of Late Proterozoic to Cambrian age (Bertrand and Caby, 1978; Caby et al., 1981; Black, 1984; Ajibade et al., 1987). About 500-800 km wide, it lies between the West African craton, to the west, and the East Saharan/Chad craton (Tuareg, Benin-Nigeria shields) and Congo craton to the east and southeast (Fig. 14). As in the Damara belt to the south, to which it may have been connected, the Dahomeyan-Pharusian belt records a history of rifting of older basement gneisses and migmatites
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(Archean and Lower Proterozoic: Eburnean, about 2 Ga), sedimentation and volcanism, followed by strong convergence leading to continental stabilization. Similar early Upper Proterozoic epicratonic sedimentary rocks on both the West African and Tuareg shields imply that there was continuous cratonic crust across the west-central Sahara before this rifting event (Caby et al., 1981). According to Ajibade et al. (1987), extension in Nigeria (Dahomeyan belt) began at about 1 Ga at the east margin of the West African craton. Depositional basins (now represented by at least ten “schist belts”) were either separate grabens or larger, now-disrupted complex basins with a variety of depositional facies. The dominant sedimentary fill was pelitic and semipelitic, with minor conglomerate, quartzite, and volcanic rocks; the Anka belt (Nigeria) contains more volcanic rocks, both mafic and felsic. These rocks were subsequently metamorphosed to greenschist or low amphibolite facies and complexly de€ormed in the Pan-African E-W convergence, and were intruded by syn- and post-tectonic granitoid rocks at about 600 Ma. TI the north in the Pharusian belt, Algeria (Bertrand and Caby, 1978; Caby et al., 1981) extension began about 800 Ma. It is represented by N-S dike swarms of diabase, andesite, and later alkaline rocks that cut the basement gneisses in the western Hoggar (Algerian Sahara). Platform sedimentary rocks <1 Ga old that overlie the West African craton rocks are intruded by a wide variety of mafic and ultramafic rocks, including a layered ultramafic intrusion >4 km thick. These rocks were then cut by diabase dikes. This mafic magmatism appears to be related to crustal rifting; actual separation to produce oceanic crust is suggested by the occurrence of ophiolites along strike to the north in southern Morocco and to the south in eastern Mali (Leblanc, 1981; Bodinier et al., 1984). The NE-SW Gourma Trough in Mali contains >8 km of mostly clastic sedimentary rocks, and is interpreted as an aulacogen in the West African craton (Caby et al., 1981). This is associated with eastward-thickening continental-margin sedimentation to the south along the east edge of the Volta Basin in Togo and Benin. An Andean-type continental-margin volcanic arc then developed, producing a thick sequence of calc-alkaline volcanic rocks, volcaniclastic deposits, and calc-alkaline batholiths during the subsequent Pan-African convergence.
Iapetus rift Several intriguing models have recently been proposed for the rifting and breakup of a supercontinent in latest Proterozoic (Riphean, Vendian) and Early Cambrian time (e.g. Bond et al., 1984; Dalziel, 1991; Hoffman, 1991). One of the best documented portions of this worldwide rift complex is the segment along what is now the southeastern margin of Laurentia, a zone that anticipated the Mesozoic rifting that led to the formation of the North Atlantic Ocean. This early, temporary sea, later destroyed by a series of Paleozoic convergence episodes, is known as the Proto-Atlantic or Iapetus Ocean. The width of this ocean has been estimated to have been at least 2000 km (Williams, 1980). Because of the complex
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subsequent history of long-distance drift, subduction, obduction, arc accretion, collision, and renewed rifting, the evidence for the continental dismemberment that produced the Iapetus Ocean is fragmentary and widespread. The following sketch by no means attempts to cite all of this evidence. Much of the evidence for the lapetus rift is found in the Appalachian orogen in the southeastern U.S. (Bird and Dewey, 1970; Rankin, 1976; Rankin et al., 1989; Thomas, 1991) (Fig. 4). The opposite side of the rift was for a long time assumed to be Africa (e.g. Bird and Dewey, 1970; Hatcher, 1987) but recent interpretations have proposed the west side of South America (Bond et al., 1984; Dalziel, 1991; Hoffman, 1991). The cratonic basement that was rifted consists of Grenville-age (1.1 Ga) gneisses and plutonic rocks. They were intruded in the Virginia area by the alkalic-peralkaline Crossnore felsic plutonic complex (Rankin, 1975), dated at 690 f 10 Ma (U/Pb, Odom and Fullagar, 1984), which may represent crustal melts from the initial rifting process. These give a maximum age for the extension, as Crossnore debris is found in rift-basin fill. Timing of the breakup is well constrained at about 625 f Ma in the central Appalachians (Fichter and Diecchio, 1986; Bond et al., 1984) by tectonostratigraphic analysis, subsidence models and paleontology. Rifting may have begun as early as 760 Ma in the Maritime Provinces (Strong et al., 1975) or 700 Ma in Quebec (Seguin, 1982), and as late as Early to Middle Cambrian in the Ouachita rift in Oklahoma and Texas (Thomas, 1991). The principal continental separation apparently took place along a zone that is now in the Appalachian Piedmont Province, buried beneath younger thrust sheets. It appears as a nearly continuous gravity gradient and high representing the attenuated cratonic margin and possibly mafic volcanics (Cook and Oliver, 1981). Landward rift basins, analogous to the Triassic/Jurassic basins associated with the Atlantic opening, formed along what is now the eastern flank of the allochthonous Blue Ridge anticlinorium. Associated with this extension were aulacogenic depressions at the Sutton Mountain (Quebec), South Mountain (Pennsylvania) and Mount Rogers (Virginia) salients (Rankin, 1976; Cook and Oliver, 1981; Thomas, 1991). These rift basins of the “Blue-Green-Long axis” (northeast Georgia to Quebec) contain 1 to 4 km thick sequences of a variety of subaerial and subaqueous clastic rocks, including tillite, with associated bimodal volcanic rocks. These volcanic-dominated formations include the Catoctin Fm in northern Virginia and Maryland (dated at 570 f 36 Ma, Rb/Sr, Badger and Sinha, 19S8; Reed and Clarke, 1989), the Mt. Rogers Fm (Virginia, North Carolina, Tennessee) and the Grandfather Mtn Fm (North Carolina), all of which contain tholeiitic metabasalts and rhyolites along with clastic rocks. In western North Carolina, amphibolites of the Ashe and Alligator Back Formations appear to represent oceanic crust, with geochemistry that resembles both N- and T-type MORB and implies several sources (Misra and Conte, 1991) (Fig. 16). Greenstones in the Camels Hump Group, Vermont have geochemical signatures of within-plate basalts and E-MORB (Coish et al., 1985). Associated with these volcanic rocks, and probably feeders for them, is a suite of
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Proterozoic rifts Till00
Hi/
A
A
3
B
Th TiO,
Nbl 2
D
zri
4
Ta
Y
A 10
Fig. 16. Geochemical diagrams for basaltic rocks related to Iapetus rifting. Symbols: X = Dalradian, Scotland (Graham and Bradbury, 1981); solid squares = Tibbit Hill and Huntington greenstones, Vermont (Coish et al., 1985); diamonds = Hancock and Gillett Pond greenstones, Vermont (ibid);open squares and open circles = Saw dolerite dikes, Norway and Sweden (Solyom et al., 1984). A. After Pearce and Cann, 1973 (A = island arc tholeiites; B = same plus MORB and calc-alkali basalts; C = calc-alkali basalts; D = within-plate basalts). B. After Wood (1980): N = normal MORB; E = E-type MORB; W = within-plate basalts; D = basalts of destructive plate margins. C. After Meschede, 1986: A = within-plate basalts; B = P-type MORB; C = within-plate tholeiite and volcanic-arc basalt; D = arc basalt. D. After Mullen, 1983: OIT = oceanic island tholeiites; OIA = oceanic-island alkali basalts; M = MORB; L4T = island-arc tholeiites; CAB = calc-alkali basalts. Note lack of orogenic characteristics.
NE-trending high-Ti tholeiitic to transitional diabase dikes that extends from Newfoundland to North Carolina (Rankin et al., 1989). In the Great Smoky Mountains in western North Carolina, Rnnessee, and Georgia, the southwesternmost rift basin contains up to 12 km of exclusively clastic rocks of the Ocoee Supergroup. These rift-fill sequences were subsequently overlain by the siliciclastic, transgressive, latest Proterozoic (Vendian) or Early Cambrian Chil-
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howee Group (and correlatives in the northern Appalachians), which marked the initial subsidence of the new continental trailing margin and which was succeeded by Cambrian marine sedimentary rocks deposited from the southeastern U.S. to Scotland. At the northeastern end of the Appalachian orogen proper, mafic rocks present a similar record of the initiation of Iapetus rifting. Small, undeformed remnants of tholeiitic plateau basalts occur on the Humber Peninsula, northern Newfoundland, and on islands in the Strait of Belle Isle, between Newfoundland and Labrador (Strong and Williams, 1972; Strong, 1974) (Figs. 4, 16). The basalts overlie arkosic sandstone which is unconformable on Grenville basement. These lavas are associated with, but chemically distinct from, a large suite of NE-trending, tholeiitic mafic dikes that cut the Grenville rocks in the Long Range of the Humber Peninsula and were probably feeders for now-eroded flows (Strong, 1974). These dikes have been dated at 605 Ma (40Ar/39Ar,Stukas and Reynolds, 1974); similar NE-trending dikes to the north in the Long Range of eastern Labrador are 615 f 2 Ma old (U/Pb: Kamo et al., 1989). To the northeast in Scotland and Ireland, the Dalradian Supergroup consists of Late Proterozoic and Cambrian sedimentary and volcanic rocks deposited on Lewisian basement gneisses (Fig. 17). The Argyll Group (Riphean to Lower Cambrian?) and the Southern Highlands Group (L. Cambrian?) in Scotland constitute a sequence of clastic and mafic volcanic and intrusive rocks that similarly record the initiation and development of continental rifting along the (now) northwest flank of the nascent Iapetus Ocean (Graham and Bradbury, 1981). The earliest extensional phase followed deposition of the Port Askaig
Fig. 17. Map of northwestern British Isles showing rift-related Proterozoic rock localities. Patterned area: Dalradian rocks; TO = fault-bounded Torridonian Basin; LMG = Loch Maree Group. After several sources.
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Tillite, and resulted in deposition of the thick Jura Quartzite and the metabasaltic Killiecrankie Schist in the mid- to late Vendian, at about 630-625 Ma (Anderton, 1982). Although up to 5 km of metabasalts of MORB geochemical affinity (Fig. 16) are present in the Early Cambrian(?) Tmyvallich Volcanics in the Southwest Highlands, clasts of Lewisian-type gneiss entered the deepening rift basin from the southeast, suggesting that, as in the Blue Ridge, the main Iapetus rift and new continental margin were farther to the southeast (Graham and Bradbury, 1981; Anderton, 1982). Cambrian to Ordovician trailing-margin shelf sedimentation followed, with subsequent Caledonian convergence and thrusting. In northwestern Ireland Dalradian amphibolites show geochemical affinities to both within-plate and ocean-floor basalts (Fig. 16), consistent with a rift origin (Winchester et al., 1987). In western Scotland a thick sequence of fluvial and lacustrine clastic rocks (Stoer Group, about 1 km thick; Sleat and Torridon Groups, about 6 km) unconformably overlie Archean and Early Proterozoic basement gneisses. These “Torridonian” redbeds occupy a NNE-trending basin about 80 km wide east of the Outer Hebrides, that was subsequently disrupted by Caledonian Moine thrusting to the northwest (Stewart, 1982). Allen et al. (1974) conclude that the redbeds were shed off a highland source to the northwest (Greenland?) into the expanding Iapetus sea. Shales in these two sequences (Stoer, Sleat/Torridon Groups) have been dated at 968 and 777 Ma respectively by Rb/Sr (Moorbath, 1969), presumably reflecting sedimentation/diagenesis ages, but it is possible that they are younger and belong to the Iapetus rifting episode. Stewart (1982) suggests that they represent early, pre-Iapetus extension. Plate reconstructions place the Baltic Shield on the east side of the Iapetus Ocean, partly at least on the basis of the marked contrast in early Paleozoic fauna across the North Sea. Evidence for Iapetus rifting is found in rocks that have been caught up in southeastward Caledonian thrusting. The uppermost Proterozoic (Riphean, Vendian) Hedmark Group, central Norway (Fig. 5), is a 34 km sequence dominated by fluvial, shallow-marine, and turbiditic sandstones and conglomerates (“sparagmites”) (Bjorlykke et al., 1976; Nystuen, 1982). They were deposited in N- to NNE-trending grabens, and were succeeded by latest VendianEarly Cambrian continental-shelf sedimentary rocks. Within the Hedmark Group, and underlying the Moelv Tillite, is a thin sequence of metabasaltic flows (Furnes et al., 1983), and mafic dikes of similar age are contained in several other Caledonian nappes. For example, the 100 km long swarm of Ottfjallet Dolerites (dated at 665 Ma, 40Ar/39Ar;Claesson and Roddick, 1983) cut continental and shallow-marine sedimentary rocks including tillites in the Middle and Upper Allochthons (Kumpulainen and Nystuen, 1985). These dikes are also referred to as the Baltoscandian Dike Swarm (BDS) (Solyom et al., 1984; AndrCasson, 1987). Although the Hedmark greenstones are strongly metasomatized, immobile trace elements imply a within-plate tholeiitic to MORB character (Fig. 16), whereas the BDS, which in places constitutes well over 50 percent of the exposed rock, is dominated by P-MORB geochemistry with minor within-plate alkalic
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character. These suites show strong similarities to mafic rocks associated with the modern Red Sea Rift (Andreasson, 1987). The large, tholeiitic-to-alkaline Seiland Complex in northwestern Norway, which includes peridotites, a tholeiitic layered series, syenites and carbonatites, may also have been intruded into this rifting continental margin (Krill and Zwaan, 1987).
Australo-American Dough, Proto-Arctic and Proto-Pacific rifting Strikingly similar sedimentary sequences along the northern Rocky Mountains, the Canadian Arctic borderlands, eastern China and central Australia suggest that these areas were all adjacent and shared general sedimentological and tectonic histories during the Late Proterozoic, preparatory to and during breakup of the supercontinent (e.g., Preiss and Forbes, 1951; Young, 1984; Eisbacher, 1985; Bell and Jefferson, 1987). In northwestern North America two major successive sequences are found: the Mackenzie Mountains Supergroup (MMSg) and the overlying Windermere/Ekwi Supergroup (Fig. 13). The MMSg (Yukon and Northwest Territories, NWT) is generally interpreted as a subsiding cratonic shelf sequence over attenuating basement, but its base is rarely exposed. Where it is, it overlies Wernecke Supergroup rocks (see above) deformed in the Racklan orogeny (Young, 1984) at about 1220 Ma. These sequences have been sliced by eastward-directed Laramide thrusting. The MMSg consists of 5-7 km of shales, mature sandstones, limestones and dolostones (Aitken, 1981; 1982). Near its top the Little Dal Group includes basalt lavas along with renewed clastic deposits and evaporites, recording the onset of the next phase of extensional tectonics. The MMSg was then locally block-faulted and eroded before renewed deposition of the “copper cycle” (Redstone River, Coppercap Formations). These contain lensoidal accumulations of conglomerate, redbeds, evaporites, marine shales and carbonates, and debris-flow deposits that were emplaced in an actively rifting environment. This disturbance eventually led to uplift and erosion at approximately 850-800 Ma, in what is known as the Hayhook orogeny (Aitken, 1982; Young, 1984). The Mackenzie Mountains Supergroup can be correlated with similar thick shelf sequences across the Arctic borderlands of Canada, from Victoria Island to northwest Greenland (e.g., Shaler and Rae Groups, Bylot Supergroup). These deposits accumulated in a series of generally northwest-trending, fault-controlled embayments (aulacogens?) in the craton (Amundsen Embayment, Fury and Hecla Basin, Borden Basin, Thule Basin; Fig. 2), which record several cycles of subsidence and shallowing (Jackson and Ianelli, 1981; Young, 1981). In northern Baffin Island, for instance, the 5.6 km thick Bylot Supergroup contains basal coarse clastics overlain by a few subaerial tholeiitic plateau basalt flows, and higher in the sequence a large graben and horst complex developed, producing thick conglomerates (Jackson et al., 1980). This was probably due to the same rifting event that affected the Little Dal and Copper Cycle beds to the west, and is thought to have culminated in continental separation to form the Proto-Arctic
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(Poseidon) Ocean (Jackson and Ianelli, 1981). The rocks are cut by the 750 Ma Franklin diabase dikes (Fahrig, 1987). To the south of the Mackenzie Mountains, equivalent strata record unstable shelf sedimentation in a series of westward-facing embayments down the Rocky Mountains into the northwestern U.S. (upper part of the Belt Supergroup), in the Wasatch and Uinta Mountains of northeastern Utah (5 km thick Big Cottonwood Formation, 7 km thick Uinta Mountain Group), in northern Arizona (Unkar and Chuar Groups of the Grand Canyon Supergroup), in southeastern Arizona (Apache Group), and in southeastern California (Crystal Spring, Beck Spring Formations) (Stewart, 1972). In the Grand Canyon, these rocks were affected by a terminal rifting/block-faulting event known as the Grand Canyon Disturbance at about 820 Ma (Elston and McKee, 1982). In South Australia, the Warrina Supergroup of the Adelaide Geosyncline, centered in the Flinders Range, records sedimentation on an unstable, rifting platform very similar to that in northwest Canada (von der Borch, 1980; Preiss and Forbes, 1981; Bell and Jefferson, 1987) (Fig. 9). The Burra Group, consisting of deltaic sandstones, evaporites, tidal and subtidal dolostones and shales, resembles the MMSg and overlies the Callanna Group rift-related sequence which includes graben-fills and volcanic rocks similar to the upper Wernecke Supergroup and its Belt/Purcell equivalents. Large breccia zones are also found in this sequence (Bell and Jefferson, 1987). Somewhat more stable basins developed in the interior of the Australian craton at the same time as the Adelaide geosyncline and were probably connected to it, and they more sensitively record a history of more and less intense rifting, starting at 900 Ma (Lindsay et al., 1987). As in North America, this cycle ended at about 800 Ma. Renewed rifting, block faulting and subsidence all along the North American Cordillera (Stewart, 1972) initiated the final Proterozoic cycle of sedimentation. Known as the Windermere Supergroup (up to 9 km thick) near the U.S.-Canada border, this cycle is correlated with the 5 km thick Ekwi Supergroup in northwestern Canada and the upper Tindir Group in eastern Alaska (Young, 1984). In northern British Columbia, basal Windermere conglomerate nonconformably overlies granite dated at 728 f 8 Ma ( U P b , Evenchick et al., 1984). The preWindermere tectonic disturbance is known as the Goat River orogeny in British Columbia. This cycle begins with an abrupt increase in tectonic instability and subsidence as well as climatic change, producing a locally discontinuous mixture of tillites and glaciomarine deposits, jaspilitic iron-formations, and basaltic rocks in the Rapitan Group. l h o periods of glaciation are indicated in some sections (Aitken, 1982; Crittenden et al., 1983). The WAr and Rb/Sr systems of the basalts from the basal Windermere in eastern Washington are too disturbed by burial metamorphism to provide dependable ages, but immobile-element geochemistry shows similarities to both ocean-floor and intra-plate basalts (Devlin et al., 1985), consistent with a rifting environment (Fig. 18). Higher-grade mafic and ultramafic Windermere rocks from southeastern British Columbia (Sevigny, 1988) are of both alkaline and
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Fig. 18. Geochemical diagrams for rift-related Late Proterozoic basaltic rocks, North American cordillera. A. Northern Utah and SE Idaho (Harper and Link, 1986); note clear within-plate signature. Diagram after Pearce and Nony (1979). B. Huckleberry volcanics (Windermere), NE Washington (after Devlin et al., 1985); fields as in A. C. Windermere amphibolites, SE Canadian cordillera (Sevigny, 1988). MORB fields after Le Roex et al. (1983). D. Mt. Harper volcanics, Ogilvie Mts, Yukon (asterisks) and Huckleberry volcanics, Washington (squares), after Roots (1988). Diagram from Mullen (1983), fields as in Fig. 16.
tholeiitic character (Fig. 18), and apparently were derived from a heterogeneous source. Basalts associated with diamictites in basal Windermere equivalents in southeast Idaho and adjacent Utah are also both alkaline and tholeiitic, with transitional MORB to within-plate trace-element signatures (Harper and Link, 1986) (Fig. 18). In the Ogilvie Mountains, Yukon, the correlative, bimodal Mt. Harper Volcanic Complex (Fig. 18) was deposited in active graben and contains both submarine and subaerial tholeiitic basalts and rhyolites dated at about 751 f 20 Ma (Roots, 1988). Mafic intrusions dated at about 770 Ma (Rb/Sr; Armstrong et al., 1982) that intrude underlying MMSg strata in the Mackenzie Mountains are also thought to be related to this initial extensional event. These basal, glaciation- and rift-related rocks of the Windermere cycle are succeeded by clastic and carbonate rocks typical of a somewhat more stable but subsiding platform (Stewart, 1972; Aitken, 1982; Young, 1984; Ross, 1991). Again,
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a similar sequence, starting with glacial deposits and iron formations, is recorded in the Adelaide Geosyncline, the Stuart Shelf to its west, and in the thinner successions of the more stable interior Australian basins (von der Borch, 1980; Preiss and Forbes, 1981; Bell and Jefferson, 1987; Lindsay et al., 1987). The position of these latest Proterozoic depositories in North America and Australia with respect to the actual continental margins has been a matter of considerable uncertainty. The strong resemblance of these sequences to typical trailing-margin successions suggests that the continental separation occurred earlier, perhaps at the time of the Hayhook/Grand Canyon Disturbance, but the occurrence of glacial deposits, evaporites, and other rocks of continental character on both sides of the Proto-Pacific also indicate that separation did not commence until the end of the Proterozoic. Thus Bell and Jefferson (1987) have proposed the existence of an “Australo-American Ttough”, over an unstable, tectonically thinned strip of the supercontinent, that lasted for over two hundred million years up until final continental breakup. Paleomagnetic analysis indicates that East Antarctica lay to the south of Australia and also was across the trough from southwestern North America at this time (Dalziel, 1991), possibly replacing the Siberian craton that had occupied that position before about 1500 Ma (Sears and Price, 1978). Cambrian marine strata conformably or disconformably overlie the latest Proterozoic beds along this entire trough, as the North American Cordilleran Geosyncline and Australian basins continued to fill. Therefore the best evidence for the actual time of continental separation comes from calculations of tectonic subsidence curves (Bond and Kominz, 1984; Lindsay et al., 1987). According to this record, final cratonic rift breakup along the western margin of Laurentia occurred essentially at the Proterozoic/Phanerozoic transition (600-555 Ma). Apparently, minimal igneous activity was associated with this major breakup. Whatever rift-related basalts were erupted, or new oceanic crust formed, have since been subducted or dismembered and drifted to unknown destinations in the Pacific region.
PROTEROZOIC RIFTS AND PLATE TECTONICS
Nearly all studies of the Proterozoic features reviewed above have described and interpreted these features in terms of the plate tectonics paradigm. Certainly the evidence for continental rifting and mobility is overwhelming; in many places fully convincing evidence of complete Wilson cycles can be found. Yet there remain several aspects of Proterozoic geology related to continental rifting for which a conventional plate-tectonic explanation is not straightforward, as has been pointed out by several authors. Among these are the many “mobile belts” of Africa: to what extent did they involve crustal thinning and/or separation and ocean formation before (re)convergence, and what sort of thermotectonic regime in the mantle was responsible? The same questions could be asked of the several Early-to-Middle Proterozoic basins in Australia that subsequently became sites
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of orogenesis, as well as the question why they do not appear to be connected structurally to each other though undergoing more or less contemporaneous rifting and deformation. What was the literally underlying cause of the large, epicratonic but nonorogenic basins in southern Africa? How could the (proposed) AustraloAmerican Trough last for hundreds of millions of years without actually separating and allowing drift of the opposite sides? Finally, why did the North American Midcontinent Rift, one of the largest and longest of all, abruptly stop just short of producing a new ocean, and revert to a compressional regime? What actually was its relation to the coeval Grenville orogeny? These and many other questions will keep isotope geochemists, paleomagnetists, field geologists, geophysicists, and geochronologists well occupied for the forseeable future. REFERENCES Aitken, J.D., 1981. Stratigraphy and sedimentology of the Upper Proterozoic Little Dal Group, Mackenzie Mountains, Northwest Territories. In: EH.A. Campbell (Editor), Proterozoic Basins of Canada. Geol. Surv. Can. Pap., 81-10: 47-71. Aitken, J.D., 1982. Precambrian of the Mackenzie fold belt - A stratigraphic and tectonic overview. In R.W. Hutchinson, C.D. Spence and J.M. Franklin (Editors), Precambrian Sulphide Deposits. Geol. Assoc. Can., Spec. Pap., 25: 149-161. Ajibade, A.C., Woakes, M. and Rahaman, M.A., 1987. Proterozoic crustal development in the PanAfrican regime of Nigeria. In A. Kroner (Editor), Proterozoic Lithospheric Evolution. Am. Geophys. Union, Geodyn. Ser., 17: 259-271. Alapieti, TT, Filen, B.A., Lahtinen, J.J., Lavrov, M.M., Smolkin, V.E and Voitsekhovsky, S.N., 1990. Early Proterozoic layered intrusions in the northeastern part of the Fennoscandian Shield. Mineral. Petrol., 4 2 1-22. Allen, P., Sutton, J. and Watson, J.V., 1974. Torridonian tourmaline-quartz pebbles and the Precambrian crust northwest of Britain. J. Geol. SOC.London, 130: 85-91. Anderson, J.L. and Morrison, J., 1992. The role of anorogenic granites in the Proterozoic crustal development of North America. In: K.C. Condie (Editor), Proterozoic Crustal Evolution. Developments in Precambrian Geology, 10, Elsevier, Amsterdam, pp. 263-299 (this volume). Anderton, R., 1982. Dalradian deposition and the late Precambrian-Cambrian history of the North Atlantic region: a review of the early evolution of the Iapetus Ocean. J. Geol. Soc. London, 139: 421-431. Andrkasson, P.-G., 1987. Early evolution of the Late Proterozoic Baltoscandian margin: inferences from rift magmatism. Geol. Foren. Stockholm Forh., 109(4): 336-340. Armstrong, R.L., Eisbacher, G.H. and Evans, P.D., 1982. Age and stratigraphic-tectonic significance of Proterozoic diabase sheets, Mackenzie Mountains, northwest Canada. Can. J. Earth Sci., 1 9 316323. Arndt, N.T, Brugmann, G.E., Lehnert, K., Chauvel, C. and Chappell, B.W., 1987. Geochemistry, petrogenesis and tectonic environment of Circum-Superior Belt basalts, Canada. In: TC. Paraoh, R.D. Beckinsdale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC., Spec. Publ., 3 3 133-146. Badger, R.L. and Sinha, A.K., 1988. Age and isotopic signature of the Catoctin volcanic province: Implications for subcrustal mantle evolution. Geology, 16: 692-695. Baker, B.H+and Morgan, P., 1981. Continental rifting: Progress and outlook. EOS, Trans. Am. Geophys. Union, 62 (29): 585-586.
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Baragar, W.R.A., 1987. Volcanic geochemistry of the northern segments of the Circum-Superior Belt of the Canadian Shield. In: TC. Pharaoh, R.D. Beckinsdale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC.,Spec. Publ., 33: 113-132. Baragar, W.R.A. and Scoates, R.EJ., 1981. The Circum-Superior Belt: a Proterozoic plate margin? In: A. Kroner (Editor), Precambrian Plate Tectonics. Elsevier, Amsterdam, pp. 297-330. Barnes, S.-J., 1983. Pan-African serpentinites in central South West Africamamibia and the chemical classification of serpentinites. In: R.M.Miller (Editor) Evolution of the Damara Orogen of South West Africamamibia. Geol. SOC.S. Afr., Spec. Publ., 11: 147-155. Barnes, S.-J. and Sawyer, E.W., 1980. An alternative model for the Damara Mobile Belt: Oceanic crust subduction and continental convergence. Precambrian Res., 13: 297-336. Barton, J.M., Ryan, B., Fnpp, R.E.P. and Horrocks, P., 1979. Effects of metamorphism on the Rb-Sr and U-Pb systematics of the Singelele and Bulai gneisses, Limpopo Mobile Belt, southern Africa. Geol. SOC.S. Afr. Trans., 8 2 259-269. Behrendt, J.C., Green, A.G., Cannon, W.F., Hutchinson, D.R., Lee, M.W., Milkereit, B., Agena, W.E and Spencer, C., 1988. Crustal structure of the Midcontinent rift system: Results from GLIMPCE deep seismic reflection profiles. Geology, 16: 81-85. Behrendt, J.C., Hutchinson, D.R., Lee, M., Thornber, C.R., Trehu, A,, Cannon, W and Green, A., 1990. GLIMPCE seismic reflection evidence of deep-crustal and upper-mantle intrusions and magmatic underplating associated with the Midcontinent Rift system of North America. Tectonophysics, 173: 595-615. Bell, R.T. and Jefferson, C.W., 1987. An hypothesis for an Australian-Canadian connection in the Late Proterozoic and the birth of the Pacific Ocean. Pacific Rim Congress 87 Proc., pp. 39-50. Bergh, S.G. and Torske, T, 1988. Palaeovolcanology and tectonic setting of a Proterozoic metatholeiitic sequence near the Baltic Shield margin, northern Norway. Precambrian Res., 39: 227-246. Bernasconi, A,, 1987. The major Precambrian terranes of eastern South America: A study of their regional and chronological evolution. Precambrian Res., 37: 107-124. Bertrand, J.M.L. and Caby, R., 1978. Geodynamic evolution of the Pan-African orogenic belt: A new interpretation of the Hoggar Shield (Algerian Sahara). Geol. Rundsch., 67 (2): 357-383. Bird, J.M. and Dewey, J.E, 1970. Lithosphere plate-continental margin tectonics and the evolution of the Appalachian orogen. Geol. SOC.Am. Bull., 81: 1031-1060. Bjorlykke, K., Elvsborg, A. and Hoy, T, 1976. Late Precambrian sedimentation in the central sparagmite basin of south Norway. Nor. Geol. Tidsskr., 5 6 233-290. Black, R., 1984. The Pan-African Event in the geological framework of Africa. Pangaea, 2 6-16. Blake, D.H., 1980. Volcanic rocks of the Paleohelikian Dubawnt Group in the Baker Lake-Angikuni Lake area, District of Keewatin, N.W.T Geol. Surv. Can. Bull., 309, 39pp. Blake, TS. and Groves, D.I., 1987. Continental rifting and the Archean-Proterozoic transition. Geology, 15: 229-232. Bodinier, J.B., Dupuy, C. and Dostal, J., 1984. Geochemistry of Precambrian ophiolites from Bou h e r , Morocco. Contrib. Mineral. Petrol., 87: 43-50. Bond, G.C. and Kominz, M.A., 1984. Construction of tectonic subsidence curves for the early Paleozoic miogeocline, southern Canadian Rochy Mountains: Implications for subsidence mechanisms, age of breakup, and crustal thinning. Geol. SOC.Am. Bull., 9 5 155-173. Bond, G.C., Nickeson, P.A. and Kominz, M.A., 1984. Breakup of a supercontinent between 625 Ma and 555 Ma: new evidence and implications for continental histories. Earth Planet. Sci. Lett., 70: 32.5-345. Borg, G., 1988. The Koras-Sinclair-Ghanzi Rift in southern Africa. Volcanism, sedimentation, age relationships and geophysical signature of a Late Middle Proterozoic rift system. Precambrian Res., 38: 75-90.
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Borg, G. and Maiden, K.J., 1987. Alteration of late Middle Proterozoic volcanics and its relation to stratabound copper-silver-gold mineralization along the margin of the Kalahari Craton in SWA/ Namibia and Botswana. In: TC. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC.,Spec. Publ., 3 3 347-354. Brannon, J.C., 1984. GeochemistIy of Successive Lava Flows of the Keweenawan North Shore Volcanic Group. Ph.D. Dissertation, Washington University, St. Louis, Mo., 312 pp. Breitkopf, J.H. and Maiden, K.J., 1987. Geochemical patterns of metabasites in the southern part of the Damara Orogen, SWA/Namibia: applicability to the recognition of tectonic environment. In: TC. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. SOC.,Spec. Publ., 3 3 355-361. Brewer, TS. and Atkin, B.P., 1987. Geochemical and tectonic evolution of the Proterozoic Telemark supracrustals, southern Norway. In: T.C. Pharaoh, R.D. Beckinsale and D. Rickard (Editors), Geochemistry and Mineralization of Proterozoic Volcanic Suites. Geol. Soc., Spec. Publ., 33: 471-487. Burke, K. and Dewey, J.E, 1973. Plume-generated triple junctions: Key indicators in applying plate tectonics to old rocks. J. Geol., 81: 406-433. BVSP (Basaltic Volcanism Study Project), 1981. Basaltic Volcanism on the Terrestrial Planets. Pergamon, New York, N.Y., 1286 pp. Caby, R., Bertrand, J.M.L. and Black, R., 1981. Pan-African ocean closure and continental collision in the Hoggar-Iforas segment, central Sahara. In: A. Kroner (Editor), Precambrian Plate Tectonics. Elsevier, Amsterdam, pp. 407-434. Campbell, I.H. and Griffiths, R.W., 1990. Implications of mantle plume structure for the evolution of flood basalts. Earth Planet. Sci. Lett., 99: 79-93. Cannon, W.E, Green, A.G., Hutchinson, D.R., Lee, M., Milkereit, B., Behrendt, J.C., Halls, H.C., Green, J.C., Dickas, A.B., Morey, G.B., Sutcliffe, R. and Spencer, C., 1989. The North American Midcontinent rift beneath Lake Superior from GLIMPCE seismic reflection profiling. Tectonics, 8 (2): 305-332. Chandler, V.W., McSwiggen, P.L., Morey, G.B., Hinze, W.J. and Anderson, R.R., 1989. Interpretation of seismic reflection, gravity, and magnetic data across middle Proterozoic Mid-Continent Rift system, northwestern Wisconsin, eastern Minnesota, and central Iowa. Am. Assoc. Pet. Geol. Bull., 73 (3): 261-275. Chapman, H.J., 1979. 2,390 m.yr. Rb-Sr whole rock age for the Scourie dykes of north-west Scotland. Nature, 271 642-643. Chase, C.G. and Gilmer, 7: H., 1973. Precambrian plate tectonics: the Midcontinent gravity high. Earth Planet. Sci. Lett., 21: 70-78. Claesson, S., 1987. Nd isotope data on 1.9-1.2 Ga old basic rocks and metasediments from the Bothnian Basin, Central Sweden. Precambrian Res., 35: 115-126. Claesson, S. and Roddick, J.C., 1983. 40Ar/39Ardata on the age and metamorphism of the Ottfjallet Dolerites, Saw Nappe, Swedish Caledonides. Lithos, 1 6 61-73. Coish, R.A., Fleming, ES., Larsen, M., Poyner, R. and Seibert, J., 1985. Early rift histoly of the ProtoAtlantic Ocean: geochemical evidence from metavolcanic rocks in Vermont. Am. J. Sci., 285: 351378. Condie, K.C., 1989a. Geochemical changes in basalts and andesites across the Archean-Proterozoic boundary: identification and significance. Lithos, 23: 1-18. Condie, K.C., 1989b. Plate Tectonics and Crustal Evolution. Pergamon, New York, N.Y., 3rd ed., 476 PP. Cook, EA. and Oliver, J.E., 1981. The Late Precambrian-Early Paleozoic continental edge in the Appalachian orogen. Am. J. Sci., 281: 993-1008.
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Chapter 4
GEOCHEMISTRY AND SIGNIFICANCE OF MAFIC DYKE SWARMS IN THE PROTEROZOIC J. TARNEY
INTRODUCTION
Mafic dyke swarms are an important feature of the Proterozoic and in parts of some stabilised Archaean cratons may be the only significant geological event in perhaps 2 Ga. Elsewhere, in less stable regions, the dyke swarms are affected by Proterozoic orogenic activity and can potentially be important time markers. Proterozoic swarms tend to be voluminous in terms of the number, width and length of dykes. The genesis of each swarm clearly constitutes a major thermal event affecting the Earth’s mantle. Moreover, because dyke swarms are often parallel to major transpressional shear zones, and the dykes may be subsequently affected by these shear zones, there is clearly an important tectonic control on their genesis. What is more surprising is that recent careful and precise U-Pb dating (e.g. LeCheminant and Heaman, 1989; Heaman and Tarney, 1989) is beginning to indicate that emplacement of any particular swarm took place over a very restricted time period. So that not only has a large amount of thermal energy to be concentrated in order to melt the mantle, but the energy has to be delivered quickly and then apparently shut-off. To accomplish this rapid burst of activity is in itself an important constraint on mantle processes. The aim of this chapter is to summarise some of the more significant features of Proterozoic mafic dyke swarms, and try to account for these features in the context of mantle evolution. There are some clear similarities - but also some differences - with Phanerozoic continental flood basalt provinces, and comparisons will be made with both continental and oceanic flood basalt provinces, where relevant. The injection of mafic dyke swarms at intervals throughout the Proterozoic provides a useful window to monitor mantle evolution, particularly the subcontinental lithosphere, which appears to be the dominant source component of most dyke swarm magmas. There is always the question, whether or not the lithosphere is the dominant source (Weaver and ’hrney, 1981), of the extent to which the magma compositions have been modified by other processes such as fractional crystallisation en route to the surface (”hrney and Weaver, 1987), fractional crystallisation in RTF (periodically replenished, tapped and fractionated) magma chambers (Cox, 1988), assimilation with fractional crystallisation (‘‘AFC”: DePaolo, 1981), thermal erosion of deep crust by mafic magmas (Huppert and Sparks, 1985) or more substantial
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ponding of magmas near the Moho with melting, assimilation, storage and homogenisation (“MASH’: Hildreth and Moorbath, 1988) before emplacement. Even without crustal involvement it is potentially possible to account for the range of compositional variation in terms of some combination of the following: partial melting, dynamic melting (Langmuir et al., 1977; Wood, 1979), disequilibrium melting (Bkdard, 1989), veined mantle sources (Brney et al., 1950), mantle lithosphere enrichment processes (Hawkesworth et al., 1990), or contributions from heterogeneous mantle reservoirs (MORB, HIMU, EM1, EM2, PREMA: Hart and Zindler, 1989). Moreover these mantle sources could be further modified by subduction zone processes (Saunders et al., 1980, 1991; Sun et al., 1989) or subducted sediment contamination (Weaver et al., 1986; Hergt et al., 1989). All the above are possible factors to consider, but not all are necessarily likely in the context of Proterozoic crust-mantle evolution.
FORM AND FEATURES
From a physical viewpoint there is nothing unusual about dykes (Emerman and Marrett, 1990): low-viscosity magmas will naturally form sheets and it takes less energy to propagate a fracture than to deform host country rocks to accommodate rounded diapiric forms. Fractures can propagate rapidly and opening fractures can be rapidly filled with fluid magma. A recurrent problem is the extent to which intruding dykes have propagated laterally or vertically. We know that dykes can penetrate laterally for many tens of kilometers from recent magmatic centres in Iceland (Sigurdsson, 1987) and as much as 200 km from the Tertiary centres in NW Scotland, and that their compositions can remain essentially constant throughout this length. Similarly the 2150-Ma “Long Dyke” in West Greenland is compositionally uniform for ca. 400 km (Kalsbeek and ’Bylor, 1956). But the spectacular 1270 Ma Mackenzie Swarm in Canada (Fahrig, 1987; LeCheminant and Heaman, 1989) radiates outwards for almost 2500 km from the “centre” marked by the Muskox Intrusion, and it is conceptually difficult to imagine why magma should penetrate laterally for such huge distances. Conversely, if the dyke magmas are penetrating vertically it is necessary to specify that the source compositions and magma-generating processes must remain constant over similar distances, which is equally difficult to conceive. The volume of magma involved in the Mackenzie Swarm is estimated at 90000 km3 (Fahrig, 1987), comparable with some Phanerozoic continental flood basalt provinces. If this volume of magma were held in some central magma chamber it would be expected to assimilate crust or differentiate to produce more silicic magmas, which are absent. Cadman et al. (1990) noted that there is evidence of lateral propagation around the Tkrtiary plutonic centres of NW Scotland, which have fractionated/assimilated to produce silicic compositions (and dykes), but Proterozoic dykes in the same area (and elsewhere) show none of these features. Hence, if there are central igneous complexes feeding Proterozoic dyke swarms,
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they must be located at depth - but they have never been uncovered by erosion. The related question of whether the extensive Proterozoic dyke swarms were matched by equally extensive eruptive volcanic sequences is usually circumvented by assuming that these ancient Proterozoic terrains are deeply eroded and the evidence removed. However, major volcanic sequences (not connected with dyke swarms) are commonly preserved in both the Archaean and the Proterozoic. Cases like the Coppermine River Lavas, which are closely associated with the Muskox Intrusion at the centre of the Mackenzie Swarm, are rare. B r n e y and Weaver (1987) noted that dykes were more abundant and thicker in granulite terrains than in juxtaposed amphibolite facies terrains. This suggests an alternative explanation: simply that the dykes did not usually reach the surface. This carries a number of further implications: (a) that the source of the dyke magmas is shallow, thus lacking the hydraulic “head” to reach the surface, or (b) that, being Fe-rich, the magmas were too dense to reach the surface, or (c) that the magmas were underplated into deep sub-crustal Moho magma chambers, from which they had only limited opportunity to ascend (as in (a)). These all have a bearing on models for Proterozoic dykes (see below), but it should be noted that this is a major difference with continental flood basalts which are dominantly extrusive and appear to have a limited number of associated dykes. In a recent study of Proterozoic dyke magnetic anisotropy flow fabrics in Labrador, Cadman et al. (1992) found that initial flow fabrics were vertical, and then later replaced by horizontal fabrics. This would seem to imply that each dyke fracture initially fills vertically, then propagates laterally: if there was an escape route to the surface, vcrtical flow would dominate. Dyke swarms reflect significant extension of the continental crust, and dyke densities indicating extension of the order of 5-10% are not unusual (Cadman et al., 1990). Fahrig (1987) suggested that the Proterozoic dyke swarms around the Canadian Shield could represent “failed arm” extensional rifts in modern plate tectonic parlance. However, the width of many dyke swarms exceeds several hundred kilometers, even over 1000 km in the case of the Mackenzie Swarm (Fahrig, 1987), which is wider than most modern failed arms. Of course in the Basin and Range Province and in the Aegean, extension and crustal thinning occurs over lateral distances of several hundred kilometers, but these are not failed arms. So while the cause of the extension may be related to mantle processes, the tectonic environment has yet to be established. LeCheminant and Heaman (1989) have proposed that the Mackenzie Swarm is centred over a 1000 km diameter mantle plume head, following the model of White and McKenzie (1989) for continental flood basalts. For Early Proterozoic dyke swarms in Greenland and Scotland (Nielsen, 1987; Hall et al., 1990), it is not so easy to link them to possible failed arm rifts or plume heads. They are more closely associated with shear zones which became the focus of later Mid-Proterozoic orogenic activity. It is to be noted that shear zones may have a major transtensional as well as a transpressional component, and that in some cases (e.g. the Gulf of California) they can be associated with large degrees of extension.
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Finally, an observation worth making is that in some cases Proterozoic dykes are so abundant that extension exceeds 30%. This usually occurs near major shear zones. Despite the fact that such dykes may be emplaced into hot high-grade granulite or amphibolite facies gneisses, the country rock never shows signs of melting, even when the dykes are high-temperature picrites. This cautions against assumptions that mafic magma injection into the deep crust can cause crustal melting and severe contamination of the mafic magmas. It would appear that Proterozoic dykes rarely have sufficient superheat to promote such crustal melting.
CHRONOLOGY
Although dyke swarms have been exceptionally useful tools in separating major phases of Proterozoic orogenic activity, achieving precise dates has proved to be much more difficult. With K-Ar dating there are many problems of argon loss and excess argon (e.g. Evans and Tarney, 1964); with Rb-Sr and Sm-Nd whole-rock dates there are uncertainties over inherited source characteristics, crustal contamination and metamorphic re-equilibration; and the latter can also affect U-Pb zircon or rutile dates. Even with very careful and detailed work (e.g. Chapman, 1979; Sheraton e t al., 1990), age uncertainties of several tens or even hundreds of m.y. can result. Fortunately, U-Pb baddeleyite dates (e.g. Heaman and Thrney, 1989) and SmNd mineral isochrons (Waters et al., 1990) now seem able to provide dyke ages with much higher precision. So it has now been established that the Scourie dyke swarm in NW Scotland comprises two distinct phases of emplacement, close to 2.42 and to 2.00 Ga, but the proportion of dykes in each phase is not known. In Canada the Matachewan Swarm has now been similarly constrained to 2.45 Ga and the Mackenzie Swarm to 1.27 Ga (LeCheminant and Heaman, 1989). Halls (1987) has made a compilation of dyke ages worldwide and has identified concentrations of dyke ages at certain periods, as well as periods of apparent inactivity. It is difficult to know whether the spectrum of dyke ages (?errorchrons) that has been reported from many cratons - covering almost the whole of Proterozoic time - will eventually be narrowed down to discrete pulses with global significance. For instance the 2.4-Ga-suite is also represented in Antarctica (Sheraton and Black, 1981), by the Great Dyke in Zimbabwe and the Jimberlana Dyke in Western Australia (Hatton and Von Gruenewaldt, 1990). It is interesting to note that the high-Mg noritic compositions are dominant at this age.
PETROLOGICAL CHARACTERISTICS
Amongst most Proterozoic dyke swarms it is possible to recognise several petrological types. In NW Scotland, for instance, four main types are readily distinguished (Xirney, 1973): quartz dolerites, olivine gabbros, norites and bronzite
Geochemistryand significance of mafic dyke swarms in the Proterozoic
155
picrites, with some minor more alkalic types, but with the dolerites (tholeiites) forming over 80% of all dykes. The noritic and picritic rock types represent an important magma type in the Late Archaean-Early Proterozoic, and Hall and Hughes (1987) have stressed their possible petrogenetic affinities with modern island arc boninites (Crawford et al., 1989). Xrmed “siliceous high-magnesian basalts” (SHMB) by Sun et al. (1989) to distinguish them from komatiites, this magma type seems to be more common in the Early Proterozoic, whereas more alkalic dykes appear in the later Proterozoic. The majority of Proterozoic dykes worldwide are aphyric iron-rich tholeiites comprising augite, plagioclase and titanomagnetite and minor hypersthene. Most are oversaturated quartz tholeiites, but they range to olivine tholeiites that are still quite Fe-rich. Some swarms, like those in southern Greenland (Nielsen, 1987) and Labrador (Cadman et al., 1990) are plagioclase phyric, some spectacularly so. Tholeiitic dykes that were emplaced at considerable depth, such as those in NW Scotland and (some in) West Greenland, may have large, though variable, amounts of primary hornblende (5 biotite), and kaersutite in the case of the olivine gabbros, indicating high p~~~ conditions during crystaIlisation. As these dykes were emplaced into “dq” granulite-facies host rocks, they cannot have acquired their fluids locally so their high water contents must have been inherited at the source. Interestingly, despite such high water contents, Proterozoic dykes never follow calc-alkaline fractionation trends (i.e. co-magmatic dykes of andesitic or dacitic character are absent), with the implication that high P H ~ O is not the only factor determining fractionation trends. Indeed the mantle sources beneath the old Archaean cratons seem to be quite reduced (Daniels and Gurney, 1991). H i g h p ~ does , ~ suppress plagioclase crystallisation, so it may be that the rarity of plagioclase phenocrysts in most dyke swarms is in part attributable to h i g h p ~ , ~ . Conversely shallow level dykes that have lost water could precipitate plagioclase in profusion (cf. Phinney et al., 1988); in any case high water pressures could generate basic melts with high normative plagioclase contents (Yoder and Tilley, 1962). The noritic and picritic dykes of the Scottish Lewisian are notable for their coarse grainsize and strong across-dyke symmetrical and asymmetrical petrological variations (Tmrney and Weaver, 1987), which are attributable to flowage in turbulent low-viscosity magmas and crystal settling in inclined dyke-sheets that were cooling slowly in hot country rocks. The mineralogy of these high-magnesian, low-alumina dykes is dominated by olivine and/or orthopyroxene, with smaller amounts of clinopyroxene, plagioclase but always with significant amounts of phlogopitelbiotite. Hornblende is quite rare. It is clear that these two types of dyke magma represent separate lineages, although both range from olivine-rich through to olivine-poor and silica-saturated. One is more Fe-rich and “fertile” in terms of major element components; the other is Fe-poor and refractory. Hornblende occurs in one, phlogopitebiotite in the other. Interestingly this difference is commonly apparent throughout both space and time in continental igneous sequences (e.g. Ellam and Cox, 1989), the
156
1 Tamey
phlogopite reflecting the fact that SHMB have enhanced levels of potassium and other lithophile elements (e.g. Sun et al., 1989). However, a potentially important observation (Hall and Hughes, 1990a) is that the high-Mg picritic and noritic suites are dominant in the Early Proterozoic, and are petrologically similar to many of the large layered intrusions such as Stillwater, Bushveld, Great Dyke, Jimberlana, etc. (Hatton and Von Gruenewaldt, 1990) that are found in the Late Archaean-Early Proterozoic. The tholeiitic magma type seems to have been available contemporaneously with the noritic type, and indeed may have been locally present as one of the inhomogeneous magma pulses that characterises these large intrusions. Conversely alkaline magma types are quite rare in Early Proterozoic dyke suites, but on almost every craton (e.g. North America, Condie et al., 1987; Greenland, Nielsen, 1987; Antarctica, Sheraton et al., 1990; see other compilations in Halls and Fahrig, 1987) they become much more important in the later Proterozoic and particularly about 1.1 Ga. This results in a much more diverse assemblage of magma types in the later Proterozoic (see Sheraton et al., 1990). Tarney and Weaver (1987) suggested that this resulted from continual additions of plume material to the base of the lithosphere, providing a greater diversity of mantle compositions for later thermal events. These systematic changes in dyke types throughout the Proterozoic must reflect in some way the processes of mantle evolution. It is important to note, of course, that large noritic intrusions of late Mesozoic age occur in the Himalayas (Chilas Complex: Khan et al., 1989), so it could be that noritic magmatism is more strictly linked to regions of recent active crustal growth rather than to an absolute time scale.
GEOCHEMISTRY
A large body of geochemical and isotopic data has accumulated for Proterozoic dyke suites worldwide. It is probably easiest to summarise these data by comparison with the Early Proterozoic Scourie dyke suite which has a large petrological diversity (Tamey, 1973), but is quite well characterised chronologically (Heaman and Tarney, 1989) and its trace element (Weaver and Tarney, 1981, 1983; Wood et al., 19Sl), Sr-, Nd- and Pb-isotopic (Waters et al., 1990) and O-isotopic (Cartwright and Valley, 1991) composition is very well known. The trace element geochemistry of these dykes is summarized by the mantle-normalised spiderdiagrams (Fig. 1) and chondrite-normalised R E E plots (Fig. 2). The Scourie quartz dolerites and olivine gabbros show a clear Fe-enrichment trend, and this is matched by a wide range of REE and trace element abundances, though Weaver and Tarney (1981) argued that this range could not have resulted from simple magma chamber crystal fractionation. Indeed they suggested that the greater H RE E depletion and the more picritic nature of the olivine gabbros was a result of melting at greater depth, of an essentially similar source, but with small amounts of garnet in the residue. The shapes of the REE and spider patterns
Geochemistry and signijicance of mafic dyke swarms in the Proterozoic
157
SCOURIE QUARTZ DOLERITE and OLIVINE GABBRO DYKES
)
Olivine gabbros
500-"
"
1 (b)
Rb
Ea
K
Nb
. ' SCOURIE BRONZITE PICRITE and NORITE DYKES "
La
Ce
"
Sr
Nd
"
P
Zr
Ti
Y
Fig. 1. a. Mantle-normalised trace element patterns for Scourie Iow-Ti quartz dolerite dykes (open symbols) and olivine-gabbro dykes (closed symbols), after Weaver and Tarney (1981). The patterns of the two dyke types are very similar, but the olivine gabbros lack Sr anomalies but show strong Y depletion, possibly attributable to garnet, instead of plagioclase, being residual in the source. b. Patterns for norites and bronzite picrites are similar to each other, but appear more fractionated than those in (a) because of incompatible element enrichment of a more refractory (hanburgitic) host (Weaver and Tarney, 1981). Small negative Sr and P anomalies probably reflect source as plagioclase and apatite are only late-crystallising phases in these dyke types.
are thus representative of the source. The REE patterns tend to be sigmoidal and convex-upward, and in fact not too dissimilar to those of Icelandic basalts (Thrney e t al., 1980). The patterns show moderate enrichment in the LREE and lithophile elements, distinct negative Nb anomalies and smaller negative Sr and Eu anomalies, and while the dolerites have small negative Ti anomalies, the olivine gabbros do not. By contrast the bronzite-picrite and norite REE patterns are concave-upwards (indeed more like patterns of ocean basalts from the FAMOUS area: Tarney
J. Tarney
158 loot
50
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and NORITE DYKES
1.o
La Ce
Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb
Lu
Fig. 2. a. REE plots of olivine gabbro dykes show greater HREE depletion than the low-Ti quartz dolerites, reflecting residual garnet in source (see Fig. 1). b. Norite and picrite R E E patterns are strongly concave-upwards and more fractionated compared with those in (a). The lack of fractionation between the MREE and HREE suggests residual garnet in source is unlikely.
et al., 1980). While alumina and the more compatible trace elements such as Zr, Y, Tb and Ti are much lower than in the dolerite dykes, as would be expected if the picro-norites were derived from a more refractory source or through much higher degrees of mantle melting (or both), there is much greater relative enrichment in the LREE and the lithophile elements K, Rb, Th, Ba, etc. The negative Nb anomaly is much more pronounced than in the dolerites. Weaver and Tarney (1981) argued that the picrite and norite dykes could not be related to each other by fractional crystallisation, but that they were both derived by partial melting of a similar refractory mantle source, the picrites resulting from melting at greater depth. The consistently greater K20/A1203 ratios in the picro-norites relative to the dolerites accounts for the almost ubiquitous presence of phlogopitebiotite rather than hornblende in these norites. Interestingly this
Geochemistryand SigniJcanceof ma$c dyke swarms in the Proterozoic
1.59
PROTEROZOIC NORlTlC DYKES OOA Scotland V A E.Antarctica
SEGreenland SWGreenland
1
:
Y 1
Rb
Ba
K
Nb
La
Ce
Sr
Nd
P
Zr
TI
Y
Fig. 3. Multielernent patterns for Early Proterozoic noritic dykes from Scotland (Weaver and Tarney, 1981), Greenland (Hall and Hughes, 1987, 1990a, b) and East Antarctica (Sheraton and Black, 1981), showing that chemical characteristics are veIy similar. Negative Nb, Sr, P and Ti anomalies most likely reflect the source, as mineral phases containing these elements are never on the liquidus of this dyke type.
is a feature of all noritic magmas (even modern ones such as Chilas), and the accompanying high Rb/Sr ratio, which is probably inherited from the source, means that there is rapid growth in s7Sr, and hence it is not surprising that most of the large noritic or SHMB intrusions such as Bushveld, Stillwater, Great Dyke, etc., have high (though variable) initial 87Sr/86Srratios, a feature which Hatton and Von Gruenewaldt (1990) ascribed to contamination of the mantle source with subducted sediment. Several pertinent petrogenetic observations can be made at this stage: (1) Hall and Hughes (1987, 1990a, b) have shown that abundant norite (BN) dykes accompany or predate the 2.1 Ga “MDl”, “MDY, “MD3” and slightly younger Kangamiut dolerite swarms in both West and East Greenland, and that the norites essentially have the same distinctive major and trace element compositions (cf. Fig. 3) as the Scourie norites. Moreover, most other Early Proterozoic norite dykes, from Antarctica (Sheraton and Black, 198l), South America (Wirth et al., 1990) and North America (Hall et al., 1987) are very similar. l3king account of the large noritic layered intrusions, this is an abundant magma type in the Early Proterozoic. But it was always available at the same time as tholeiite magma. How can two very different magma types be available essentially contemporaneously? How is it that they do not mix? It is very difficult to reconcile this observation with models of massive basaltic magma chambers underlying the lower crust, which should homogenize such diverse magma types. (2) Tholeiite dykes are rather uniform in composition throughout the Proterozoic. Surprisingly they are also similar to many tholeiitic flood basalts, from as far apart in space and time as Bsmanian and Karoo dolerites. This is il-
L Tamey
160 500-
'
'
'
1 (a)
'
'
'
'
EARLY PROTEROZOIC DYKES & CONTINENTAL FLOOD BASALTS 1
1
10 Karoo Central Dolerites
Rb
Ba
K
Nb
La
Ce
Sr
Nd
P
Zr
Ti
Y
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Sr
Nd
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'
'
t
t 1-
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Rb 500L
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N. INDIAN MAFIC ROCKS (Early Proterozoic)
I + Aravalli Basal Tholeiite 1 Rb
Ea
K
Nb
La
Ce
Sr
Nd
P
Zr
TI
Y
Fig. 4. a. Multielement diagram showing close chemical similarities of Proterozoic low-Ti dykes and Phanerozoic low-Ti continental flood basalts, in contrast to (b) wide diversity of patterns shown by modern oceanic lavas and possible Archaean crust and post-Archaean terrestrial shale contaminants. Diagram (c) shows that Early Proterozoic extrusive suites also share the low-Ti basalt characteristic of prominent negative Nb and Sr and small negative P and Ti (after Ahmad and Tarney, 1991).
Geochemistry and significance of mafic dyke swarms in the Proterozoic
161
lustrated in Fig. 4a (after Ahmad and Drney, 1991), and contrasts with the spectrum of other magma types shown in Fig. 4b. This magma type is not found beneath the oceans, but clearly is continually available beneath the continents from the Early Proterozoic (Fig. 4c) onwards. Hergt et al. (1989, 1991) persuasively argued, on trace element grounds, that the Tasmanian and Ferrar dolerites were derived from a mantle source pervasively contaminated by ca. 3% subducted sediment. But similar arguments would apply to the mantle source of the norites, which otherwise have quite different major element compositions. In theory it is possible to reconcile these two models by arguing that the dolerites are derived from sediment-contaminated asthenosphere and the norites from sediment-contaminated refractory lithosphere, but this fails to explain the bimodal distribution and the lack of intermediate members. ( 3 ) Students of continental flood basalts and of Proterozoic dyke swarms are largely split into those who want massive crustal contamination of magmas and those who prefer to contaminate the mantle source with sediment or subduction-derived fluids in order to account for the “continental” trace element characteristics. Because Lewisian country rock gneisses have such an anomalously low-Rb, -U, -Th composition, Weaver and Tmrney (1981) were able to rule out completely, on trace element grounds, any significant country gneiss contamination of any of the four Scourie dyke magmas. This conclusion has been fully sustained by the detailed Nd and Pb isotopic studies of Waters et al. (1990), and the oxygen isotopic data of Cartwright and Valley (1991). Similarly, Hall and Hughes (1990a) have argued that to achieve the composition of the Greenland norite dykes through contamination of a simple MORB-like magma would require impossibly large amounts (ca. 70%) of gneiss contaminant. Sheraton et al. (1990) have forcefully argued that crustal contamination is not an important factor in the petrogenesis of East Antarctic Proterozoic dykes. The argument then resolves itself if the wide range of Scourie dyke compositions have been generated without crustal contamination of magmas, there must be other processes which can produce such coexisting diverse magma types on a global scale. (4) From the limited Sr isotopic data then available, Weaver and Brney (1981) suggested a lithosphere mantle source for the Scourie dyke suite, and that its continental signature had developed at the same time as the Lewisian crust (i.e. at about 2.9 Ga). The recent Sm-Nd and U-Pb whole rock isotopic data for the Scourie dykes (Waters et al., 1990) now confirm that the trace element characteristics were established in the lithosphere source at about 3.0 Ga, some 0.6 to 1.0 Ga before the dykes were emplaced. ( 5 ) The oxygen isotope data of Cartwright and Valley (1991) provide an important new key to the whole problem (Fig. 5). They show that the wholerock S ” 0 values for the Scourie dykes are rather uniform at ca. 2%0, which is significantly below the “upper mantle” values of 6%0 normally seen in basalts. Because high-temperature magmatic 6l80 distributions are preserved in the primary minerals, and because the Sl8O values for the enclosing gneisses or adjacent shear zones are not anomalous, secondary effects cannot be responsible.
. l Tamey
162
t
Ibas u-basi
I SCOURIE DYKES Fig. 5. Summary of oxygen isotope data for nine Scourie dolerite dykes (after Cartwright and Valley, 1991). Seven fresh d kes have very low whole-rock 6"O values of ca. 2%0, and these dykes preserve high temperature 6l 0 mineral distributions (see inset), implying whole-rock values are primary. The two sheared dykes show partial re-equilibration of oxygen isotopic compositions with enclosing gneisses. SMOW = Standard Mean Ocean Water.
& 7 .
Hence these low SlSO values characterize the dyke magmas and must have been inherited from the source. Further, it is argued that the only way of achieving such low 6l80 values in the source is to subduct hydrothermally altered oceanic crust into the source regions of the Scourie dyke magmas. There are some further consequences arising from these data which are explored below, but the immediate implication is that the volume component that was being added to the lithosphere to become the major material contributor to the Scourie tholeiitic magma, was subducted hydrothermally altered'mafic material (?amphibolite). This new information begins to offer a solution to the intractable problem of contemporaneous tholeiitic and noritic dyke magmas. For instance, it permits the dominant Fe-rich quartz tholeiites to be generated from mafic material (or mantle highly permeated by subducted mafic material), and the Cr- and Ni-rich noritic magmas to be derived through melting of more refractory lithosphere.
Later Proterozoic dykes It is not possible in this short review to synthesise all the information on Middle to Late Proterozoic dyke swarms, but better to focus on the more important petrogenetic features. For instance, in the Southern Superior Province five dyke swarms (Condie et al., 1987), were emplaced over a 1.5 Ga period from the Matachewan plagioclase-phyric dykes at 2.45 Ga to the thick 700 km long Abitibi Swarm at 1.1 Ga. They show a systematic change in chemistry with time from typical Fe-tholeiites, like those described above, to distinctly alkaline dolerites with high levels of Ba and Sr, much more fractionated REE and positive rather than negative Nb anomalies. This distinctly alkaline characteristic at about 1.3-1.1 Ga is even more evident in the Gardar Province in southern Greenland (Nielsen, 1987). Just how complex the dyke chemistry can get, even in a small area, in the
Geochemistry and signijicance of ma$c dyke swarms in the Proterozoic
163
later Proterozoic is well shown by the dyke swarms in the Bunger Hills area of East Antarctica (Sheraton et al., 1990) most of which were emplaced at about 1.1 Ga, but the latest at 0.5 Ga. Five different suites of dykes are recognised, which range from quartz and olivine tholeiites through to alkaline dolerites and picrite ankaramites. They display a surprisingly wide range of initial Sr (0.703 - 0.717) and Nd EN^ = +6.3 to -18.6) isotopic compositions, and an equally broad range of spidergram patterns. The latter vary from patterns typical of Early Proterozoic tholeiites (moderately fractionated with small negative Nb, Sr and P anomalies) to highly fractionated patterns with large negative Nb anomalies, to patterns with distinctly positive Ba, Nb, Sr, P and Ti anomalies. A selection is shown in Fig. 6. At least three source components need to be involved to explain the dyke 500 .
'
BUNGER HILLS, E. ANTARCTICA -
: (4
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n
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Fig. 6. Multi-element patterns for Late Proterozoic dyke suites from Bunger Hills, after Sheraton et al. (1990), who divided the dykes into 5 groups and several sub-groups. This illustrates much greater diversity of patterns and much higher abundances of incompatible elements in the later Proterozoic compared with the Early Proterozoic (see Fig. 1),arguably a consequence of continual additions of enriched material to the base of the lithosphere.
J. Tamey
164
trace element and isotopic compositions: a depleted asthenospheric component, a lithosphere component (possibly including subducted sediment) and a Nb-rich alkaline OIB component. As with some of the Early Proterozoic dykes discussed above, there is a requirement for some of these mantle components to be added to the sub-continental lithosphere at the time of continent generation. A recurring theme in this and many recent studies of Proterozoic dykes is the need to develop some chemical characteristics of the mantle sources supplying the dyke magmas at quite an early stage of crustal evolution - often many hundreds of Ma before the dykes themselves were emplaced. This is also the case with Phanerozoic continental flood basalts: for instance the characteristic Parana chemistry can be recognised in Proterozoic mafic suites in Brazil (Oliveira and Tmrney, 1989), and the Karoo source may also have been initiated in the Proterozoic (Ellam and Cox, 1989). The next section attempts to reconcile some of the problems of Proterozoic dyke generation with some general aspects of mantle evolution.
MANTLE EVOLUTION
The Earth's mantle is now known to consist of a number of chemically distinct reservoirs which have been isolated from one another for periods in excess of 1.8 Ga, but which are nonetheless contributing to the spectrum of basalt compositions through plume activity, including continental flood basalts and dyke swarms. On the basis of isotopic diagrams such as 87Sr/s6Sr vs. 143Nd/144Nd,87Sr/86Sr vs. 206Pb/204Pb,87Sr/s6Sr vs. 208Pb/204Pbor 207Pb/m4Pbvs. m8Pb/204Pb,a t least four different end-member mantle components have been invoked to explain the variations amongst ocean island basalts (e.g. Hart and Zindler, 1989): DMM
(MORB reservoir): high 143Nd/144Nd,low s7Sr/86Srand 208Pb/204Pb, very low 'OSPb/ '06Pb, and low Ba/Nb, Th/Nb and K/Nb.
HIMU
(e.g. Mangaia, St. Helena): high 'O'Pb, '07Pb and 20sPb,low 207Pb/206Pb,208Pb/206Pb and s7Sr/s6Sr,high U/Pb, low LIL/Nb.
EM1
(e.g. Walvis Ridge): low 143Nd/144/Nd and 87Sr/86Sr,high 207Pb/206Pb and 20sPb/206Pb; generally high LIL/Nb.
EM2
(e.g. Samoa): as EM1 but with high 87Sr/86Srand higher Rb/Nb, K/Nb.
Some of these relationships are illustrated in simplified form in Fig. 7. Many islands have intermediate compositions which can be regarded as mixtures between these different components. The HIMU and EM1 components have low Nd ratios relative to their Sr ratios on the familiar ESr-ENd diagram, and these, along with several islands with similar low Nd characteristics, maintain reasonable consistency on all isotopic plots (the "LoNd array"). However, this consistency does not hold when other components are considered, so it is likely that the assumption
Geochemistry and significance of mafic dyke swarms in the Proterozoic I
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165
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&Nd c16-
0.5134 +12-
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IDMM B I I
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Fig. 7. Summary diagrams illustrating the compositional characteristics of main mantle components in Sr-Nd and Sr-Pb isotopic space (simplified after Hart and Zindler, 1989). PUM = primordial uniform mantle; DMM = depleted MORB mantleA and B; HIMU component with high U/Pb ( p ) ratio; “enriched” EM1 and EM2, not strictly defined but with low €Nd but variable esr; PREMA, or prevalent mantle which typifies major hotspots like Iceland and Hawaii, but could be a mix of several components.
of only four discrete components is an oversimplification. It is important to note that there are many subtle trace element differences between these mantle components, which correlate with the isotopic differences (Weaver et al., 1987; Weaver, 1991) and which were therefore established when the isotopic systems were set. There is, in addition, a rather common mantle composition represented by basalts from Iceland, Hawaii, many oceanic flood basalts and oceanic plateaus (and which is therefore volumetrically abundant) that could be regarded as a mixture of perhaps three components. Alternatively it can also be regarded as a discrete compositional entity; this has been termed PREMA (PREvalent MAntle) by Hart and Zindler (1989). This mantle composition in particular has high
166
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3He/4He ratios, compatible with a deep, rather primitive source, and consistent with deeply anchored hotspot plumes. This mantle component is not so depleted as DMM and could represent material from the lower mantle that has been isolated from the convecting upper mantle for a significant proportion of Earth history. Invoking this origin for PREMA begins to have some interesting consequences for Proterozoic dyke swarms. Providing explanations for the formation and preservation of the various mantle reservoirs (or components) is more difficult. We know that the earliest Archaean mantle products were moderately isotopically depleted, and that the early “planetary” stage of Earth evolution may have left the whole mantle slightly depleted (e.g. PREMA-like). Since then the convecting upper mantle above the 670km discontinuity has become progressively more depleted in incompatible elements to form the DMM reservoir. To a first approximation this correlates with the growth of continental crust, but strictly (cf. Saunders et al., 1988; Sun and McDonough, 1989) it results from basalt extraction at ocean ridges, from which crustal components are extracted at subduction zones, and the residues (which then have many of the compositional characteristics of the HIMU OIB component) are then removed down the subduction zone. So the complement to the continental crust is not just the DMM reservoir, but DMM plus an OIB component. It has been suggested (e.g. Ringwood, 1985, 1990; Ringwood and Irifune, 1988) that the subducted slab residues have been stored a t the 670 km discontinuity, from which they may rise (giga years later) as plumes to feed the ocean island hotspots. Accounting for EM1 and EM2 compositions is not so easy because their Pb isotope compositions require their chemical parameters to be set or acquired in the Late Archaean or Early Proterozoic, or it requires contamination with material of this age. It is possible to account for some of the isotopic characteristics of EM1 or EM2 by metasomatism of, and storage in, the sub-continental lithosphere (e.g. McKenzie and O’Nions, 1983; Hawkesworth et al., 1986). Alternatively, Weaver et al. (1986) and Weaver (1991) have shown that it is possible to account quantitatively for many of the isotopic and trace element characteristics of EM1 and EM2 by contaminating the HIMU residue with abyssal or terrigenous sediment respectively during subduction (in the Precambrian). However as recent models of subduction zones (Peacock, 1991; Davies and Bickle, 1991; Saunders et al., 1991; Davies and Stevenson, 1992) require induced convection of the mantle wedge in order to achieve thermal and material balance in magma genesis, and therefore entrainment and progressive downdrag of the subcontinental lithosphere, then lithosphere fractionation and contamination processes can be employed too, i.e. competing models are not so far apart. So are the enriched characteristics of most Proterozoic dykes and continental flood basalts the result of progressive lithosphere enrichment processes, or of plume additions from the mesosphere?
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PREMA mantle and Proterozoic dykes At this stage it is necessary to consider further the status of PREMA (1) If predominantly residing in the lower mantle, the PREMA reservoir has escaped the depletion processes that have affected the convecting DMM upper mantle since the early Archaean. When it rises as plumes to form ridge-centred hotspot islands like Iceland, it not only has more potential thermal energy, but is also more fertile in terms of major elements (cf. Brooks et al., 1991), hence it is able to generate ocean crust two or three times thicker than normal (cf. White and McKenzie, 1989); indeed ocean crust that may become sub-aerial. (2) Larson (1991) has recently shown that there was a major spurt in oceanic crust production between 120 and 80 Ma, but this resulted not in a significant increase in global spreading rate, but was manifest in production of oceanic plateaus (e.g. Ontong Java, Manihiki Rise) with considerably over-thickened ocean crust like Iceland. Moreover, because this period coincides almost exactly with the Cretaceous magnetic quiet zone (normal polarity), he argues that this represented a major release of material j?om the lower mantle that actually affected the convective behaviour of the outer core and inhibited magnetic reversals for 41 Ma. ( 3 ) Many of these ocean plateaus and oceanic flood basalt provinces still exist in the western Pacific, where some have been sampled via ocean drilling. However, in the eastern Pacific, where their counterparts suffered attempted subduction along the Cordillera of South America, this was clearly difficult because large volumes were obducted along the coastal belt of Colombia. Indeed some refused to be subducted and were carried through to form the floor of the Caribbean. The implication is that these oceanic plateaus are dificult to subduct. (4) Geochemical studies of these Colombian volcanics (Millward et al., 1984; Guevara, 1987) and the drilled ocean plateaus (Saunders, 1986) show that their closest geochemical counterparts (both trace element and isotopic) are with Icelandic/Reykjanes Ridge basalts. As stressed above, the closer geochemical counterparts of Proterozoic tholeiitic dykes amongst oceanic basalts are Icelandic basalts. Interestingly, some of the few modern high-Mg counterparts of Archaean komatiites occur within the Colombian and Caribbean obducted volcanic sequences (Gorgona, Curacao, Romeral), as emphasised by Storey et al. (1991). There are two very important points here which bear upon the problem of accounting for the oxygen isotope data for the Scourie dykes, reported above, which required that their source be hydrothermally altered ocean crust injected into the Lewisian sub-continental mantle. The first difficulty with the model is in explaining why oceanic crust should be injected into the lithosphere rather than being subducted in the normal fashion. The observations above, on Colombian/Caribbean volcanic sequences, suggest that thick PREMA-type crust may be rather resistant to subduction, simply because it is warmer (McKenzie and Bickle, 1988), less dense and less likely to transform to eclogite to provide the “slab-pull” force; hence it is more likely to underplate into a rheologically weak young lithosphere, as required.
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The second problem is in accounting for the low S1'0 values in the proposed underplated ocean crust, because low values of the order of ~ 2 % 0in equilibrium with seawater can only occur in the deeper sections of ocean crust - therefore it is necessary to strip off, and dispose elsewhere, the whole upper section of ocean crust during underplating to maintain these low values. However, if the PREMA ocean crust was sub-aerial, like Iceland, this problem disappears because the isotopic exchange is with meteoric water which (particularly in polar climates) can result in very low exchange S1'0 values. In the event it would not be necessary to strip off all but the low S1'0 value rocks, but simply homogenize them. Note that with a smaller continental crust volume in the Archaean, and hence an average shallower ocean, it is more likely that elevated plateaus could become subaerial (cf. Abbott, 1954; Schubert and Reymer, 1985; Galer 1991). However, because of the latitude dependence of the oxygen isotopic composition of meteoric water, we should not expect all underplated altered ocean crust to have low S " 0 values. There is a fairly wide range in S1'0 reported from pyroxenite veins and eclogites thought to be derived from the lithosphere (Pearson et al., 1991), and which are interpreted as subduction components, and which reflect the diverse nature of material being subducted. This model is useful in quite a number of respects: (1) Because the Earth's upper mantle has evolved from PREMA composition to DMM with time, there is more likelihood of thick sub-aerial crust like Iceland being generated in the late Archaean, and therefore potentially more thick crust that would underplate rather than subduct. If dolerite dykes are linked to this underplate, is this why there is an abundance of dykes near old potential sites of plate subduction like the Nagssugtoqidian Belt (= Kangamiut dykes)? (2) The material is already a low-melting component when underplated. It is also hydrous. It would therefore be very vulnerable to flushing out in huge volumes with the development of any major thermal anomaly. This would explain the high primary hornblende contents of some Proterozoic dykes. It also accounts for the fact that they are Fe-rich since during melting they would not necessarily be in equilibrium with Mg-rich mantle. It obviates the need for sub-crustal magma chambers to hold the mafic magmas while they fractionate to Fe-rich compositions before emplacement. Some of the observed compositional variation in Proterozoic dolerites (which is always difficult to account for by fractional crystallisation) could, in fact, have been inherited from the underplated oceanic crust. (3) The isotopic data for the Scourie dykes (Waters et al., 1990) which give Sm-Nd whole rock "isochrons" of ca. 3.0 Ga, and indicate that the U-Pb systems were disturbed perhaps as much as 3.0 Ga ago, are perfectly consistent with the model as these would indicate the age of processes associated with the formation and emplacement (underplating) of the source. (4) If Iceland can be regarded as a typical product of a PREMA source, it should be noted that, despite the source being moderately 'depleted' with respect to Sm-Nd and Rb-Sr isotopic systems, the erupted lavas are dominantly LREE-enriched (see compilation in Walker, 1991, fig. 13.11). Hence an underplate
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of such material would experience retarded growth of 143Ndand evolve towards low 6Nd compositions (and high cSr) depending upon the time stored. Such characteristics are common in Proterozoic dykes (e.g. Waters e t al., 1990) as well as continental flood basalts. (5) The source of the dykes is relatively shallow, which is one of the requirements outlined earlier to account for the apparent lack of co-genetic volcanic suites. (6) It is now possible to explain the norite dykes as partial melting products of the sub-crustal mantle lithosphere. This source may have been harzburgitic because of previous melt extraction; alternatively, as silica is always mobile in the subduction environment, and as excess silica is liberated when low-silica hornblende- or garnet-assemblages develop in subducted mafic rocks, the mantle may have become harzburgitic because of silica metasomatism (olivine > orthopyroxene).
THERMAL PROBLEMS IN DYKE GENERATION
With dyke swarms representing volumes of mafic magma of the order of 50000 to 100000 km3, there is a need for a major thermal source which has to be focused to provide the energy for melting. Moreover, if dyke swarms are emplaced over a very short time interval of not much more than 2-3 Ma, as implied by recent U-Pb dating (LeCheminant and Heaman, 1989), then it is necessary to turn the thermal tap on and off very quickly. Within the scenario outlined above, there are two possible ways in which this might be done.
Mantle pluvltes It has been common to appeal to mantle plumes to supply this energy for melting (White and McKenzie, 1989; Campbell and Griffiths, 1990). The difficulty is that hotspots like Iceland, Hawaii, Kerguelen or Cape Verde tend to remain active for many tens of Ma. How then can the intrusive pulse be shortened? First, many of the Earth’s major hotspots that are thought to represent deep mantle plumes, such as Iceland, Kerguelen, and Hawaii, as well as many or all of those which gave rise to the ocean plateaus (Larson, 1991), seem to have been initiated near spreading ridges, though, like Hawaii, they may later migrate off the ridge provided the mechanical boundary representing the base of the lithosphere can be raised and the magma conduit kept open. There is no inherent reason why plumes initiated in the lower mantle should be constrained by shallow-level plate boundaries near the Earth’s surface. Why do they not burn their way through the middle of plates? The reason must be that the mechanical boundary layer (MBL) beneath is too thick, and that it is only plumes that are rising near ridges that can reach the surface and turn their potential thermal energy into extrudable magma. This is not surprising: Watson and McKenzie
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(1991) have calculated that the MBL beneath Hawaii is 72 km thick and the melt-producing region only a further 55 km in vertical extent. The likelihood is that even major plumes rising beneath continents or thick ocean plates with a MBL perhaps well over 100 km thick may be unable to penetrate to the surface and simply contribute to a more fertile asthenosphere, with their energy converted into small degree melts which permeate and fertilise the lithosphere. In fact Storey et al. (1989) have suggested that the major Kerguelen plume has effectively contaminated much of the Indian Ocean asthenosphere in this way, and has contributed to the rather distinctive composition of Indian Ocean basalts. Hill (1991) has similarly argued that plumes cannot provide the ultimate driving force for continental break-up. This may be one reason why many major Proterozoic dyke swarms are closely linked to transtensional/transpressional shear zones: these provide vertical access. Without tectonic assistance, extensive adiabatic melting cannot take place, and the magma cannot penetrate upwards as dykes. The second and rather surprising point, well exemplified by studies of Mesozoic radial dyke swarms around the Cape Verde hotspot (Oliveira et al., 1990) or the development of Kerguelen (Storey et al., 1988), is that the material input to magma from the plume itself is very minor. The magmas, particularly those emplaced in the early stages, carry a strong lithospheric signature. Hence a very large proportion of the available energy is converted to lithosphere melting. There must be some very strong controlling factor here. A possible reason follows from the model discussed above. If the basaltic component in the lower lithosphere or upper asthenosphere is largely held in hydrous minerals, phlogopite, kaersutite or K-richterite, which have probably formed close to their stability limit just below the MBL, then any thermal perturbation may produce dehydration, and the fluids released so alter the rheological properties in the region below or even within the MBL so that advection replaces conduction, the whole zone becomes unstable, and large amounts of melt are available if tectonic conditions (shear zones or stretching) permit rise to the surface. The useful feature of this model is that it relies on hydrous phases to create the instability, and structural control to deliver the magma. Once the fluids are expelled along with the magma, the system becomes anhydrous, stability returns and it would take a great over-supply of thermal energy to de-stabilise it again. The dyke phase is short.
Sinking (negative)plumes What comes up must go down, and as it is density contrast that determines whether material will rise or sink, it is perfectly possible that inherently dense, but warm, mantle material might sink (cf. Griffiths and Turner, 1988). McKenzie and O’Nions (1983) suggested that portions of the subcontinental lithospheric keel might detach themselves and sink to provide ultimately an enriched source reservoir for alkali basalts. More particularly, Kroner (1981) has suggested that
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delamination and sinking of subcontinental lithosphere in a n ensiulic environment (“A-subduction”) may account for the apparent differences between Proterozoic and Phanerozoic orogenic styles, the latter being controlled by normal “Bsubduction” of oceanic lithosphere a t Benioff zones. The attractive feature of A-subduction is that the orogenic compression and crustal shortening can be achieved without having to account for the absence of ophiolites, blueschists and continental margin sedimentary sequences that normally characterise the Wilson Cycle. The further attractive feature is that as hot asthenosphere eventually rises to replace the sunken A-subducted lithosphere, a ready supply of thermal energy is provided to generate voluminous post-orogenic Proterozoic granites through melting of lower crust (cf. Houseman et al., 1981). The additionally useful factor as far as dykes are concerned is that there is a ready parallel between the boninites and island arc tholeiites commonly associated with the initial stages of B-subduction and the norites and tholeiites associated with A-subduction. However, a major problem with the A-subduction concept is that subcontinental lithosphere beneath A ch a e a n cratons is thought to be very refractory (Boyd, 1989) and inherently buoyant. Ellis (1992) has demonstrated that such lithosphere would need to be some 700°C cooler than the underlying asthenosphere for it to sink spontaneously. This effectively rules out A-subduction in normal circumstances; and as it is also argued that transformation to eclogite is unlikely to occur in any mafic material in the lower crust, there is little potential help from a mafic underplate either. However, if mafic material is emplaced into a rheologically weak lithosphere during the crustal growth phase, as implied by the Scourie dyke S ” 0 data, then transformation to eclogite is potentially possible at any later date (perhaps several hundred m.y. later), thus providing the enhanced density contrast necessary to initiate A-subduction (or a sinking plume). With A-subduction, any hornblende present in the assemblage would suffer pressureinduced breakdown, with release of fluids (as opposed to the temperature-induced breakdown caused by an uprising plume). It is this fluid, in combination with the hot uprising asthenosphere replacing the A-subducted material, which initiates melting. In physical terms the conditions are not unlike that of melting with induced convection in the mantle wedge of a modern subduction zone (Saunders et al., 1991), though the fluids would be less oxidised, and hence the magmas tholeiitic rather than calc-alkaline. Recent work by Foley (1991) has demonstrated that fluorine can substantially enhance the depth range over which hydrous minerals like pargasite, K-richterite and phlogopite are stable in the mantle, hence greatly increasing the amount of fluid that could be available for melting. With this model, the suggested correspondence between Proterozoic continental boninites (Hall and Hughes, tholeiite + norite dykes and modern arc tholeiite 1987) becomes rather more plausible, as is the fact that Proterozoic dykes share many chemical features (high LIL, low-Ti02, negative Nb anomalies) with subduction zone magmas. Equally importantly, because the amount of lithosphere that could be consumed by A-subduction is limited, this would constrain the time period of associated mafic dyke magmatism. The comparable situation in
+
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a modern arc system like the Marianas (cf. Tirney et al., 1981) is the period following initiation of subduction, where the rate of magma production is very high, and the volume of magma produced (Kyushu-Palau Remnant Arc) similar to that of a typical dyke swarm.
COMPARISON WITH CONTINENTAL FLOOD BASALTS
The discussion above explores a number of different ways in which the compositional, volume and temporal characteristics of Proterozoic dyke swarms might be accounted for. Because of the close compositional similarities between Proterozoic dolerite dykes and Phanerozoic low-Ti02 continental flood basalts (CFB) it is interesting to compare petrogenetic models. There has been a progressive shift over the last decade away from petrogenetic schemes involving crustal contamination of magmas, and two of the most recent papers on the Gondwana CFB provinces (Hergt et al., 1991; Ellam and Cox, 1991) employ subducted sediments and lamproite liquids respectively, as lithosphere contaminants before extracting the CFB magmas. In neither the dominant low-Ti02 (Ferrar, Karoo, Parana, etc.) quartz tholeiites, nor the Karoo picrites can a plume “OIB” component be recognised, though it is apparent in Deccan basalts. The compositions of the predominant uniform low-Ti basalts are consistent with moderately high degrees of melting of a relatively fertile source at moderate water pressures, and a t shallow depths (no garnet), but the consistently low ENd and variably high csr (Hergt et al., 1991) suggest either that the enriched source is old or that the contaminant is old. Contamination of the lithosphere source with subducted sediment is convenient, but difficult to prove, in that where abundant sediment is being subducted beneath arcs, very little appears in the arc magmas (Hole et al. 1984), and subducted sediment can be used equally convincingly to produce other basalt compositions (Weaver, 1991). The fact that these low-Ti basalts have spidergram patterns very similar to average post-Archaean upper crust can be interpreted in two ways: either that their source is contaminated by continental sediment (= granite), or that Proterozoic granitoids were derived from sub-continental lithosphere with a trace element composition close to low-Ti basalts. The latter is a t least consistent with the basaltic underplating model for Australian Proterozoic granitoids of Etheridge et al. (1987). If not a result of sediment subduction, the problem remains how this high Rb/Ba, high Rb/Nb and low Ti/Y component of the low-Ti basalt source is generated. It is not the lamproite component of Ellam and Cox (1991), which has high Sr and Ba, and is a more suitable end-member component for high-Ti Gondwana basalts. However, the high Rb/Ba characteristic is typical of subduction zones, a result of selective transport of LIL elements by subduction fluids (Saunders et al., 1980). If the basaltic components are largely held in hydrous phases such as phlogopite, Krichterite or hornblende (Sudo and Titsumi, 1990), these normally exert a strong control on chemistry, particularly if fluids allow some open system behaviour.
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If slab dehydration continues to much greater depths than previously thought (Ringwood, 1990) then it is possible that hydrous fluids may permeate upwards scavenging the mantle in the back-arc region, ultimately to become trapped in phlogopite or K-richterite in the subcontinental tectosphere. The distinctive negative Nb anomalies may result from a titanite phase being stable in equilibrium with and during the migration of these hydrous fluids, thus sequestering Nb, lh, Ti, etc. No real explanations are yet forthcoming to explain the location, the size and volume, the timing and the thermal causes of Gondwana CFB volcanism. Cox (1978) noted the Parana, the Karoo and the Ferrar CFB provinces reside in the back-arc region along the active margin of the reconstructed Gondwana continent, forming a semi-continuous belt some 10000 km long (see Hergt et al., 1991, fig. 1). It is difficult to envisage how rising mantle plumes could account for this distribution, and indeed a plume signature is not much in evidence in the basalt chemistry, as noted above. However, it is not difficult to imagine large segments of hot over-thickened ocean plateau crust being injected into and beneath the immature lithosphere of that Gondwana margin in perhaps the Late Proterozoic. This material is then available to be mobilised some 0.6 Ga later by plumes, rising or sinking, or during the general disturbances associated with the breakup of Gondwanaland.
CONCLUSIONS
Dyke swarms in the Early Proterozoic include mainly low-Ti quartz dolerites and Mg-rich norites, both of rather consistent composition, which have been generated from two different sources. Crustal contamination does not seem to be an important factor in their petrogenesis, nor can a plume o r asthenospheric source component be recognised except in later Proterozoic dyke swarms. Proterozoic dyke magmas share with Phanerozoic continental flood basalts the severe thermal and tectonic problems of generating huge volumes of uniform and distinctive melts from the mantle system in a relatively short time span. Data for the Scourie dykes suggest that the source for the dolerite magmas may be slivers of warm over-thickened ocean plateau crust that were too buoyant to subduct but were injected into the lithosphere beneath the newly developing continent, and mobilised some 0.6-1.0 Ga later when thermal and tectonic conditions were favourable. The norites are products of melting of silica-metasomatised or refractory harzburgitic mantle. Hydrous minerals in the source (amphibole f phlogopite for the dolerites and phlogopite for the norites) are important in controlling the chemistry and in providing the mechanism to generate large volumes of melt relatively quickly. Later Proterozoic dykes reflect the addition of more alkaline components to the lithosphere. Whereas uprising deep mantle plumes can provide the thermal energy to mobilise the dyke magmas, they must entrain major amounts of lithosphere to satisfy the compositional constraints, and it is not easy to turn the thermal
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tap off. An alternative mechanism of foundering and sinking of stored mafic ocean plateau underplate (as it converts to eclogite) could also provide short-term energy for melting as fluids are released and the sinking mass is replaced by hot asthenosphere.
ACKNOWLEDGEMENTS
The ideas in this paper arose from discussions over many years with colleagues and students, notably Barry Weaver, Andy Saunders, Mike Norry, Elson Oliveira, B l a t Ahmad and Andy Cadman. Very helpful comments on the manuscript were provided by Andy Cadman, Kent Condie, Peter Hall, Ray Kent, Andy Saunders, Shen-su Sun and John Sheraton.
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Oliveira, E.P. and Tarney, J., 1990. Petrogenesis of the Canindt do SLo Francisco Complex: a major late Proterozoic gabbroic body in the Sergipe Foldbelt, northeastern Brazil. J. S. Am. Earth Sci., 3: 125-140. Oliveira, E X , 'Grney, J. and JoSo, X.J., 1990. Geochemistry of the Mesozoic AmapA Dyke Swarm, N. Brazil: plume related magmatism during opening of the Central Atlantic. In A.J. Parker, P.C. Rickwood and D.H. Tucker (Editors), Mafic Dykes and Emplacement Mechanisms. A.A. Balkema, Rotterdam, pp. 173-183. Peacock, S.M., 1991. Numerical simulations of subduction zone pressure-temperature-time paths: constraints on fluid production and arc magmatism. Philos. Trans. R. SOC.London, Ser. A, 335: 341353. Pearson, D.G., Davies, G.R., Nixon, P.H., Greenwood, P.B. and Mattey, D.P., 1991. Oxygen isotope evidence for the origin of pyroxenites in the Beni Bousera pendotite massif, North Morocco: derivation from subducted oceanic lithosphere. Earth Planet. Sci. Lett., 102 289-301. Phinney, W.C., Morrison, D. and Manuga, D.E., 1988. Anorthosites and related megacystic units in the evolution of Archaean crust. J. Petrol., 29: 1283-1323. Ringwood, A.E., 1985. Mantle dynamics and basalt petrogenesis. Tectonophysics, 112 17-34. Ringwood, A.E., 1990 Slab-mantle interactions, 3. Petrogenesis of intraplate magmas and structure of the upper mantle. Chem. Geol., 8 2 187-207. Ringwood, A.E. and Irifune, T, 1988. Nature of the 650 km seismic discontinuity: implications for mantle dynamics and differentiation. Nature, 331: 131-136. Saunders, A.D., 1986. Geochemistry of basalts from the Nauru Basin, Deep Sea Drilling Project Legs 61 and 89: Implications for the origin of oceanic flood basalts. Init. Rep. DSDP, 89. U.S. Govt. Printing Office, Washington, D.C., pp. 653-678. Saunders, A.D., Tarney, J. and Weaver, S.D., 1980. Transverse geochemical variations across the Antarctic Peninsula: implications for the genesis of calc-alkaline magmas. Earth Planet. Sci. Lett., 46: 344360. Saunders, A.D., Norry, M.J. and Tarney, J., 1988. Origin of MORB and chemically-depleted mantle reservoirs: trace element constraints. J. Petrol., Special Lithosphere Issue, pp. 415-445. Saunders, A.D., Nory, M.J. and Tarney, J., 1991. Fluid influence on the trace element composition of subduction zone magmas. Philos. Trans. R. SOC.London, Ser. A, 335: 377-392. Schubert, G. and Reymer, A.P.S., 1985. Continental volume and freeboard through geological time. Nature, 316: 336-339. Sheraton, J.W. and Black, L.P., 1981. Geochemistry and geochronology of Proterozoic tholeiite dykes of East Antarctica: Evidence for mantle metasomatism. Contrib. Mineral. Petrol., 78: 305-317. Sheraton, J.W., Black, L.P., McCulloch, M.T. and Oliver, R.L., 1990. Age and origin of a compositionally varied mafic dyke swarm in the Bunger Hills, East Antarctica. Chem. Geol., 85: 215-246. Sigurdsson, H., 1987. Dyke injection in Iceland: a review. In: H.C. Halls and W.F. Fahrig (Editors), Mafic Dyke Swarms. Geol. Assoc. Can., Spec. Pap., 34: 55-64. Storey, M., Mahoney, J.J., Kroenke, L.W. and Saunders, A.D., 1991. Are ocean plateaus sites of komatiite formation? Geology, 1 9 376-379. Storey, M., Saunders, A.D., Tarney, J., Gibson, I.L., Norry, M.J., Thirlwall, M.F., Leat, P., Thompson, R.N. and Menzics, M.A., 1989. Contamination of Indian Ocean asthenosphere by the KerguelenHeard mantle plume. Nature, 338: 574-576. Storey, M., Saunders, A.D., Tarney, J., h a t , P., Thirlwall, M.F., Thompson, R.N., Menzies, M.A. and Marriner, G.F., 1988. Geochemical evidence for plume-mantle interactions beneath Kerguelen and Heard Islands, Indian Ocean. Nature, 336: 371-374. Sudo, A. and Tatsumi, Y., 1990. Phlogopite and K-amphibole in the upper mantle: implications for magma genesis in subduction zones. Geophys. Res. Lett., 1 7 29-32.
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Sun, S-S. and McDonough, W.F., 1989. Chemical and isotopic systematics of oceanic basalts: implications for mantle composition and processes. In: A.D. Saunders and M.J. Norry (Editors), Magmatism in the Ocean Basins. Geol. SOC.London, Spec. Publ., 4 2 313-345. Sun, S-S., Nesbitt, R.W and McCulloch, M.T, 1989. Geochemistry and petrogenesis of Archaean and early Proterozoic siliceous high-magnesian basalts. In: A.J. Crawford (Editor), Boninites. Unwin Hyman, London, pp. 149-173. Tarney, J., 1973. The Scourie dyke suite and the nature of the Inverian event in Assynt. In: R.G. Park and J. Tarney (Editors), The Early Precambrian of Scotland and Related Rocks of Greenland. University of Keele, pp. 105-118. Tarney, J. and Weaver, B.L., 1987. Mineralogy, petrology, and geochemistry of the Scourie dykes: petrogenesis and crystallization processes in dykes intruded at depth. In: R.G. Park and J. Tarney (Editors), Evolution of the Lewisian and Comparable Precambrian High Grade Terrains. Geol. Soc. London, Spec. Publ., 27: 217-233. Tarney, J., Wood, D.A., Saunders, A.D., Cann, J.R. and Varet, J., 1980. Nature of mantle heterogeneity in the North Atlantic: evidence from deep sea drilling. Philos. Trans. R. Soc. London, Ser. A, 297 179-202. Tarney, J., Saunders, AD., Mattey, D.P., Wood, D.A. and Marsh, N.G., 1981. Geochemical aspects of back-arc spreading in the Scotia Sea and Western Pacific. Philos. Bans. R. SOC.London, Ser. A, 300: 263-285. Walker, C., 1991. North Atlantic ocean crust and Iceland. In: P.A. Floyd (Editor), Oceanic Basalts. Blackie, Glasgow, pp. 311-352. Waters, EG., Cohen, A.S., O’Nions, R.K. and O’Hara, M.J., 1990. Development of Archaean lithosphere deduced from chronology and isotope chemistry of Scourie Dykes. Earth Planet. Sci. Lett., 97: 241255. Watson, S. and McKenzie, D., 1991. Melt generation by plumes: a study of Hawaiian volcanism. J. Petrol., 3 2 501-537. Weaver, B.L., 1991. The origin of ocean island basalt end-member compositions: trace element and isotopic constraints. Earth Planet. Sci. Lett., 104: 381-397. Weaver, B.L. and Tarney, J., 1981. The Scourie dyke suite: petrogenesis and geochemical nature of the Proterozoic sub-continental mantle. Contrib. Mineral. Petrol., 78: 175-188. Weaver, B.L. and Tarney, J., 1983. The chemistry of the sub-continental mantle: inferences from Archaean and Proterozoic dykes and continental flood basalts. In: C.J. Hawkesworth and M.J. Norry (Editors), Continental Basalts and Mantle Xenoliths. Shiva Publications, Nantwich, pp. 209-229. Weaver, B.L., Wood, D.A., Tarney, J. and Joron, J-L., 1986. Role of subducted sediment in the genesis of ocean-island basalts: geochemical evidence from South Atlantic Ocean islands. Geology, 14: 275278. Weaver, B.L., Wood, D.A., Tarney, J. and Joron, J-L., 1987. Geochemistry of ocean island basalts from the south Atlantic: Ascension, Bouvet, St. Helena, Gough and Tristan da Cunha. In: J.G. Fitton and B.G.J. Upton (Editors), Alkaline Igneous Rocks. Geol. SOC.London, Spec. Publ., 30: 253-267. White, R.S. and McKenzie, D.P., 1989. Magmatism at rift zones: the generation of volcanic continental margins and flood basalts. J. Geophys. Res., 94: 7685-7730. Wirth, K., Oliveira, E.P., Sa, H.S. and Tarney, J., 1990. Early Precambrian basic rocks of South America. In: R.P. Hall and D.J. Hughes (Editors), Early Precambrian Basic Magmatism. Blackie, Glasgow, pp. 379-404. Wood, D.A., 1979. Dynamic partial melting: its application to the petrogenesis of basalts erupted in Iceland, the Faroes, the Isle of Skye (Scotland) and the 2 o o d o s Massif (Cyprus). Geochim. Cosmochim. Acta, 43: 1031-1046.
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Wood, D.A., Tarney, J. and Weaver, B.L., 1981. Trace element variations in Atlantic Ocean basalts and Proterozoic dykes from northwest Scotland: their bearing upon the nature and geochemical evolution of the upper mantle. Tectonophysics, 75: 91-112. Yoder, H.S. and Tilley, C.E., 1962. Origin of basaltic magmas: An experimental study of natural and synthetic rock systems. J. Petrol., 3 342-532.
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Chapter 5
TECTONIC SETTING OF PROTEROZOIC LAYERED INTRUSIONS WITH SPECIAL REFERENCE TO THE BUSHVELD COMPLEX G. VON GRUENEWALDT and R.E. HARMER
INTRODUCTION
Layered intrusions are a feature of several tectonic environments and are known from the oldest greenstone belts of the Yilgarn, Zimbabwe and Kaapvaal Cratons (Hatton and Von Gruenewaldt, 1990) to the TertiaIy layered intrusives associated with continental rifting, such as Skaergaard and Rhum (Wager and Brown, 1968). These intrusions are of great economic significance as they host virtually all the world’s resources of the platinum-group elements and chromite, as well as important deposits of vanadium, Cu and Ni. Consequently, many of these intrusions have been studied in great detail. Most of these investigations however, have focussed on magma chamber crystallization processes in order to explain igneous layering and associated mineralization. Recently, interest has shifted somewhat to questions concerning the tectonic setting of layered intrusions and the source regions of their parental magmas. In this paper, the more important Proterozoic layered complexes are discussed in terms of their tectonic setting, postulated parental magmas, and internal stratigraphy. The primary aim of is to evaluate whether a correlation exists between tectonic setting, intrusion form and source of parental magma for the more important Proterozoic layered intrusions. Crystallization processes within magma chambers to produce igneous layering are not discussed and interested readers are referred to papers in Parsons (1987). Emphasis is placed on the Bushveld Complex, where recent investigation of its postulated parental magmas and some associated volcanic and plutonic rocks has led to debate concerning its tectonic setting. Although this paper focusses on Early Proterozoic layered intrusions, many features which characterize these intrusions also apply to many well known Late Achaean complexes, such as the Stillwater Complex in Montana and the Windimurra layered intrusion of Western Australia (Hatton and Von Gruenewaldt, 1990). Excluded from this paper is the 1850 Ma old Sudbury irruptive, since numerous lines of evidence have been cited in favour of an extra-terrestrial origin. For a detailed discussion on this topic the reader is referred to several papers in Pye et al. (1984).
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TECTONIC SETTING AND PARENTAL MAGMAS OF RIFT-RELATED PROTEROZOIC LAYERED COMPLEXES
The Great Qyke of Zimbabwe The Great Dyke extends for 550 km in a NNE direction across the entire Zimbabwe Craton (Fig. 1) and varies in width from 4 to 12 km (Worst, 1960). Together with its associated satellite dykes, the Umvimeela Dyke to the west and the East Dyke, the Great Dyke intruded 2461 f 16 Ma ago into Archaean granites and gneisses, and associated greenstone belts of the Zimbabwe Craton (Wilson, 1982; Wilson and Prendergast, 1989). Its northern extremity extends into
KAAPVAAL PROVINCE
0
\
,-
,'-I
GROENFONTEIN ANOMALY
Fig. 1. Tectonic setting of the major Proterozoic layered intrusions of southern Africa. Compiled from various sources, including Van Biljon and Legg (1983), Meixner and Peart (1984), Wilson and Prendergast (1988) and Von Gruenewaldt et al. (1988).
Tectonic setting of Proterozoic layered intrusions with reference to the Bushveld Complex
183
the Zambezi mobile belt, where dyke rocks have undergone intensive deformation, in that parts of the dyke have become dislodged, folded and rotated during an orogeny about 500 Ma ago (Wiles, 1968). In the south the dyke rocks terminate just short of the northern marginal zone of the Limpopo Province, although the southern satellite dykes continue from the Great Dyke for SO km into the mobile belt, where they are truncated by the Tuli-Sabi shear zone (Wilson and Prendergast, 1989) (Fig. 1).The tectonic setting of the dyke was recently evaluated by Wilson and Prendergast (1989) and there seems little doubt that the dyke and its satellite were emplaced during a period of crustal extension. Detailed mapping of the Great Dyke led Worst (1960) to interpret the dyke as four contiguous canoe-shaped layered complexes. These four complexes, named from north to south, the Musengezi, Hartley, Selukwe and Wedza complexes, each consist of a thick succession of ultramafic rocks and are capped in the central areas by gabbroic rocks which are considered by Worst to overlie the feeders to each of the complexes. A recent gravity survey (Podmore and Wilson, 1987) indicates that a dyke-like feeder underlies the intrusion over most of its length. The gravity survey has also revealed a noticeably shallower succession in the dyke near the junction of the Hartley and Selukwe complexes where there is no evidence for a dyke-like feeder zone. As a result, the Great Dyke has been subdivided into a Southern Chamber and a Northern Chamber by Prendergast (1987) and by Wilson and Prendergast (1989) who consider these as the remains of two major magma chambers. The stratigraphy in the two chambers is broadly similar in that in each a thick, cyclically layered sequence of ultramafic rocks is capped by a comparatively homogeneous succession of gabbroic rocks. The ultramafic sequence is subdivided into a lower chromite cyclic succession of predominantly serpentinised dunite and chromitite layers, and an upper, bronzite cyclic succession consisting of a basal chromitite layer, followed by a dunite or harzburgite, which grades upward into olivine bronzitite and bronzitite at the top (Worst, 1960; Wilson, 1982; Wilson and Prendergast, 1989). The ultramafic sequence is best developed in the Hartley Complex where fourteen cyclic units with a thickness in excess of 2000 m are developed (Wilson, 1982). The uppermost cyclic unit is remarkably similar in all four complexes, and differs from the underlying cycles in that cumulus clinopyroxene appears in a websterite layer at the top of this cycle, and because the platiniferous Main Mineralized Zone is located in bronzitite directly below this websterite. The thickness of the preserved overlying gabbroic rocks varies considerably from one complex to the next, but is best preserved in the Hartley Complex where it is estimated to be 1150 m thick. Here the mafic sequence commences with an olivine gabbro near the base, which grades into norite, the dominant rock type, and to magnetite- and quartz-bearing pigeonite gabbros at the top (Wilson and Prendergast, 1989). A liquid with about 15% MgO, and similar in composition to the chill of one of the satellite dykes (lhble l), is in close agreement with the observed and modelled
G. von Gruenewaldt and R.E. Harmer
184 TABLE 1
Postulated parental magma compositions of Proterozoic layered intrusions 1
2
3
4
5
6
7
8
9
10 -
52.78
52.0 0.68 12.3 10.2 0.18 13.0 9.26 1.74 0.71 0.20
54.2 0.47 15.3 7.74 0.13 10.10 9.00 2.93 0.14 0.00
57.0 0.33 14.37 8.17 0.10 7.60 4.55 2.96 4.61 0.10
53.87 0.41 13.16 9.16 0.17 11.83 8.57 2.14 0.68 0.00
56.32 0.33 11.43 9.42 0.18 13.14 6.45 1.75 0.90 0.07
49.94 0.68 16.21 11.96 0.20 6.95 11.61 2.15 0.15 0.14
51.51 0.40 16.32 9.01 0.16 8.37 11.64 2.35 0.20 0.03
47.94 1.24 18.95 11.67 0.14 7.67 8.60 3.21 0.40 0.00
51.33 1.01 13.69 10.16 0.18 8.94 11.60 1.84 0.50 0.10
0.55 11.04 9.35 0.14 15.60 7.60 1.77 0.69 0.11
1: Chill from the East Dyke, representative of magma parental to the Great Dyke (Wilson, 1982). 2 Loljunmaa Dyke associated with the Penikat intrusion (Alapieti et al., 1990). 3: Chilled marginal rock of the Koillismaa intrusion (Alapieti, 1982). 4: Average of 3 subophitic chilled rocks marginal to the Penikat intrusion (Alapieti et al., 1990). 5: Bronzite gabbro dyke, representative of bulk composition of the upper layered series, Jimberlana intrusion (Campbell, 1977). 6: Average composition of orthopyroxenitic sills and marginal rocks, representative of B1 parental magma, Bushveld Complex (Harmer and Sharpe, 1985). 7: Average composition of tholeiitic sills and marginal rocks, representative of B2 parental magma, Bushveld Complex (Harmer and Sharpe, 1985). 8: Average composition of tholeiitic sills, representative of B3 parental magma, Bushveld Complex (Harmer and Sharpe, 1985). 9: Chilled margin of the Hettasch intrusion, representative of parental magma of the Kiglapait intrusion (Morse, 1981). 10: Chilled bronzite gabbro marginal to the Muskox intrusion (Itvine, 1970).
crystallization sequences and compositions in the ultramafic sequence, and hence considered to be the parental magma of the Great Dyke (Wilson, 1982; Wilson and Prendergast, 1989). This is in broad agreement with Hughes (1976) who also postulated the parental magma to be a high-magnesian basalt.
Proterozoic layered intrusions of the Fennoscandian Shield Two belts of more than 20 Early Proterozoic layered intrusions extend eastward from Tornio in northern Finland into the former Soviet Union over a distance of over 400 km (Fig. 2). The layered intrusions were emplaced 2440 Ma ago, soon after cratonization of the Late Archaean crust. The southern belt has been tectonically sliced into several separate blocks during the Svecokarelain orogeny (Alapieti and Piirainen, 1984; GaB1, 1985). The intrusions can be classed into several groups: those of Tornio-Kemi-Penikat; the Portima layered complex, which includes the Suhanko and Kontijarvi bodies, the Koillismaa Complex and the Oulanka layered complex in the former Soviet Union. All of these complexes intrude Archaean granitic basement rocks, and are located at the contact between basement and supracrustal rocks of the Kemi and Koillismaa schist belts. GaAl
Tectonic setting of Proterozoic layered intrusions with reference to the Bushveld Complex
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BARENTS SEA
MURMANSK
Palaeozoic a Ika Ii complexes Svecokarelian granitoids (1900-1750 m.y.) Early Proterozoic sedimentary and volcanic rocks
Archaean rocks
-
0
50
lOOkm
Fig. 2. Distribution of Early Proterozoic layered intrusions in the northeastern part of the Fennoscandian Shield (after Alapieti et al., 1990).
(1985) proposed that these schist belts evolved in the failed arm of a triple junction during the breakup of an Archaean continent and that the layered intrusions were emplaced in response to rifting. Alapieti and Lahtinen (1986) have also suggested that this belt of layered intrusions may be related to an aulacogen.
186
G. von Gruenewaldt and R.E. Harmer
The intrusions of the northern belt are of the same age and considered to belong to the same magmatic event (Alapieti et al., 1990). These include the Koitelainen and some smaller, related layered intrusions in Finnish Lapland, as well as the Monchegorsk intrusion, the Panski Tundra and the Fedorova Tundra intrusions of the Kola Peninsula. The Kola Peninsula intrusions, similar to those of the southern belt, are situated at the junction of Archaean basement and the Early Proterozoic supracrustal rocks of the Pechenga-Verzuga belt. The largest of the 2.44 Ga layered intrusions on the Baltic Shield, the Burakovsky Complex, is situated about 500 km southeast of the Oulanka Complex in Karelia. Although the southern layered intrusions of northern Finland and adjacent former Soviet Union are considered to have been originally a single belt of several intrusive complexes that were emplaced along a deep seated tensional fracture zone (GaAl, 1985), subsequent dislocation into several separate blocks complicates reconstruction of their original form. The general geometry seems to be that of canoe-shaped bodies linked along the length of the entire belt (Ga51, 1985; Alapieti and Lahtinen, 1986). This is demonstrated for the Koillismaa Complex by Alapieti and Piirainen (1984) who postulate that the western part of the complex is a broad synformal feature that is connected to the narrow Narankavaara intrusion by a hidden connecting dyke. This dyke is postulated on the basis of a prominent gravity high that extends for over 80 km from the western part of the complex through the Narankavaara intrusion into the former Soviet Union. It is interpreted to be 3 km wide but to taper gradually with depth and to represent a feeder dyke to the Koillismaa intrusion (Alapieti and Piirainen, 1984). Little is known about the form of the other complexes. The Tornio and Koitelainen intrusions are folded into flat anticlinal structures with basement rocks cropping out in the central areas, whilst Kemi is considered to originally have been funnel-shaped (Alapieti et al, 1990). The Portimo Complex, which comprises seven separate blocks of intrusive rocks, is interpreted as originally consisting of two interconnected magma chambers. In general terms, most of the intrusions are ultramafic near the base, becoming mafic to anorthositic at their tops. The lithology varies significantly along strike and it is not possible to construct a section which is representative of all the intrusions. In the Koillismaa Complex, the Narankavaara intrusion, which is situated within the feeder dyke, is predominantly ultramafic, while the sheet-like intrusions in the western part are mafic, and contain magnetite-rich differentiates which have been mined for vanadium. Chromitite layers are restricted to the ultramafic parts of the Tbrnio, Kemi, Penikat and Burakovsky intrusions (Lahtinen, 1985; Alapieti et al., 1990), while Penikat also contains well-developed megacyclic units with anorthosite layers and thin layers with disseminated sulphides enriched in the platinum-group elements (Alapieti and Lahtinen, 1986). Disseminated Ni-Cu-PGE sulphides are a feature of the gabbroic marginal rocks of the intrusions in the Narkaus area and the western part of the Koillismaa Complex (Vuorelainen et al., 1982; Alapieti and Piirainen, 1984; Lahtinen, 1985) and the Burakovsky intrusion. Disseminated PGE-bearing sulphides are also known from the Kivakka (Alapieti
Tectonic setting of Proterozoic layered intrusions with reference to the Bushveld Complex
187
et al., 1990) and Lukkalaisvaara (TL. Grokhovskaya, personal commununication, 1990) intrusions of the Oulanka Complex. Ultramafic rocks are largely absent in the exposed parts of the Fedorova Tundra and Panski n n d r a intrusions, and apart from feldspathic pyroxenite in the western part of the Panski Tbndra intrusion, marginal gabbroic rocks are overlain in both intrusions by a thick sequence of well-layered noritic, gabbroic and troctolitic rocks, some of which are magnetite-bearing. The Koitelainen layered intrusion seems to be different to all the other intrusions of the Fennoscandian Shield in that Fe-rich chromitite layers occur high in the intrusion, in close proximity to magnetite-bearing gabbroic rocks, and in that PGE-enriched disseminated sulphides also predominate in late differentiates (Mutanen, 1989). The close association of chlorapatite with these platiniferous sulphides suggests to Mutanen et al. (1987) that a high chlorine content of this intrusion caused the platinum-group elements to remain complexed in the magma until an advanced stage in the evolution of the intrusion. Fine-grained, chilled marginal rocks that could be representative of parental magma composition have been encountered in the Koillismaa (Alapieti and Piirainen, 1984) and Penikat (Alapieti and Lahtinen, 1986) intrusions. The marginal rocks from the Koillismaa intrusion have a basaltic composition with a relatively high MgO and S O 2 and low T i 0 2 content (Table 1).Repeated influxes of parental magma are postulated to have resulted in megacycles in all these intrusions (Alapieti and Lahtinen, 1986). There is also some indication that parental magmas of contrasting composition were involved in the evolution of these intrusions (Table 1, Nos. 2 and 4). Several lines of evidence are cited by Alapieti e t al. (1990) to suggest involvement of a high-MgO basalt. Such evidence includes: (1) the recognition of boninitic volcanic rocks within the Pechenga-Vorzuga sequence which may be co-magmatic with the Fedorova and Panski Tundra intrusions; (2) high-MgO mafic dykes below the Penikat intrusion (Tmble l),also considered to be co-magmatic with Penikat; and (3) calculations of bulk composition of the lower megacyclic units in the Penikat and Narkaus intrusions.
The Jinrberlnnn innusion, Western Austualia The Jimberlana intrusion is part of the easterly trending Widgiemooltha dyke suite which was emplaced 2370 5 30 Ma ago into basement rocks of the southern part of the Yilgarn Block of Western Australia (Campbell et al., 1970). The intrusion extends for 180 km and has an average width of 1.5 km, but widens at 7 points along its length into canoe-shaped complexes which are linked by a connecting dyke. The complexes have a very steep, funnel-shaped cross-sections and contain cumulate layers, which are horizontal in the centre and steepen towards their edges (McClay and Campbell, 1976). The Jimberlana intrusion has been divided into: (1) a lower series of five macrorhythmic units, each consisting of olivine cumulates at the base and bronzite cumulates at the top, overlain by a thick layer of plagioclase-augite-hypersthene
188
G. von Gruenewaldt and R.E. Harmer
cumulates; (2) an upper layered series which rests unconformably on the lower series and consists of several macrorhythmic units of olivine and bronzite cumulates overlain by plagioclase-augite-hypersthene cumulates, as well as a granophyric layer at the top; and (3) the steeply dipping, reversed sequence of the marginal layered series below the lower series, consisting of plagioclase-augite-hypersthene cumulates at the base, overlain by bronzite cumulates which in turn are overlain by olivine cumulates. Information on the parental magma composition is limited, but field relations and petrographic evidence suggest that the upper layered series crystallized from a major new pulse of magma which entered the chamber during the final stages of crystallization of the lower layered series. A bronzite-rich gabbro-dyke which cuts the lower layered series has a composition very close to the bulk composition of the upper layered series (Table 1) and is considered by Campbell (1977) to be a possible feeder dyke to this series.
Fox River sill Strongly deformed supracrustal rocks are developed in many areas along the periphery of the stable Archaean crustal block of the Superior Province (Fig. 3). These supracrustal rocks were deposited in narrow annular troughs that originated by continental rifting along the margins of the stable Archaean block (Baragar and Scoates, 1981). Crustal rifting along with consequent subsidence resulted in deposition of miogeosynclinal sediments in ensialic basins. With further rifting, an incipient oceanic rift environment developed and resulted in the invasion and extrusion of komatiitic magma especially along the northern margin of the Superior Province. Subsequent closure of the attenuated areas resulted in compression and deformation of the volcano-sedimentary sequences. The 1720 Ma old Fox River sill is the largest of a number of differentiated syn-volcanic sills emplaced into sedimentary rocks below volcanic rocks (Scoates and Eckstrand, 1986). The sill is between 2 and 2.5 km thick and occurs as a number of compartments that extend for 250 km along strike. More than 75% of the intrusion consists of Mg-rich olivine cumulates that typically comprise the lower part of thick cyclic sequences. A variety of cumulates containing different proportions of olivine, clinopyroxene, plagioclase and orthopyroxene, constitute the remainder of the cyclic sequences. The sill has been subdivided into a marginal zone, a lower central layered zone, an upper central layered zone associated with some sulphide mineralization and a hybrid roof zone of granophyre-bearing gabbronorite (Scoates and Eckstrand, 1986). The volcanic rocks, which range in composition from basaltic komatiite or picrite to tholeiitic basalt, are interpreted by Scoates (1984) to represent lavas expelled from the Fox River sill at different times during its crystallization history. The parental magma of this large layered sill is therefore assumed to have had a composition similar to a basaltic komatiite.
Tectonicsetting of Proterozoic layered intrusions with reference to the Bushveld Complex
189
Proterozoic cover rocks Circum-Superior Belt
Fig. 3. Generalized map of the Canadian Shield showing the major tectonic provinces and the localities of the larger Proterozoic layered intrusions referred to in the text, including the Sudbury irruptive (compiled from diagrams in a.o. Irvine and Baragar, 1972; Baragar and Scoates, 1981).
Kiglapait
The Kiglapait intrusion on the east coast of Labrador (Fig. 3) is a well studied example of a plagioclase-rich layered intrusion (Morse, 1969). It is considered to be one of the younger members of the 1300-1400 Ma anorthositic Nain Complex. This complex covers an area in excess of 10000 km2 and consists of multiple intrusions of anorthositic, gabbroic, troctolitic and adamelitic plutons, emplaced along the boundary between the Archaean Nain structural province in the east and the Proterozoic Churchill Province to the west (Thylor, 1971; Berg, 1977) (Fig. 3). The Kiglapait intrusion, which underlies a n area of about 560 km2, is roughly circular in outline and funnel-shaped in section. Its entire sequence from floor to roof is exposed and is considered to be the product of closed system fractional crystallization of a basaltic parental magma (Morse, 1969). Calculation of the bulk composition of the intrusion led Morse (1979) to conclude that the parental magma was an anhydrous high-alumina, high-FeO, but low-K basalt. This postu-
190
G. von Gruenewaldt and R.E. Harmer
lated parental magma has a close analogue in the chilled magma of the nearby Hettash intrusion (’hble 1). Layered rocks constitute 94% of the volume of the intrusion and occupy the sequence between an Inner Border Zone of plagioclase-olivine orthocumulates and a n Upper Border Zone, which is regarded as an inverted sequence of the upper 20% of the Layered Group (Morse, 1969). The Inner Border Zone grades into the Lower Zone of plagioclase-olivine accumulates, which comprises 78% of the intrusion. Olivine gabbros, mostly titanomagnetite-bearing, arc the most common rock types of the Upper Zone. Such rocks display an extreme iron enrichment stratigraphically upwards and grade into ferrosyenite, the final product of fractional crystallization. Although the tectonic setting of the Kiglapait intrusion is less certain than some Proterozoic layered intrusions, regional gravity data indicate a negative anomaly centered on the Nain Complex. This negative anomaly has been interpreted as an ancient graben and led Berg (1977) to suggest that the Nain Complex, including the Kiglapait intrusion, was emplaced during rifting of a stable continental block. Morse (1981) presented various geochemical arguments that the postulated parental magmas of the Kiglapait and Hettasch intrusions are unlikely to represent primary melts of peridotite mantle. He emphasized that these parental magmas have compositions close to the density minimum for fractionating basaltic liquids and that this would favour their emplacement into the crust. The Kiglapait magmas were argued to be derived from melting of depleted, spinel lherzolite at a depth of about 35 km below a thinned crust in the postulated rift zone (Morse, 1981).
Muskox intrusion The 1200-1250 Ma Muskox intrusion is a large, dyke-like stratiform intrusion of ultramafic, gabbroic and granophyric rocks located in the Bear Province in NW Canada (Fig. 3). The intrusion is 120 km long and funnel-shaped in section. Towards the southern extremity it narrows and grades into a 150-500 m wide vertical “feeder” dyke which extends over 60 km in a south-southeasterly direction (Fig. 4). Northwards the intrusion dips at about 5” beneath its roof rocks with the feeder dyke probably forming a “keel” to the intrusion. A pronounced elongated positive gravity anomaly suggests that the intrusion possibly continues for another 120 km beneath a succession of fairly flat lying, Late Proterozoic sediments and volcanic rocks of the Hornsby Bay, Dismal Lake, Coppermine River and Rae Groups (Fig. 4). The Muskox intrusion has been studied in detail by Smith (1962), Smith and Kapp (1963), Irvine and Smith (1967), and Irvine (1970, 1975) who subdivided the 2000 m thick sequence of layered rocks into 25 cyclic units. Although ultramafic rocks predominate, the sequence grades upward into a thin, discontinuous sheet of granophyre and contact breccia, in which fragments of roof rocks are set in a granophyric matrix. Two stages of fractional crystallization arc noted by Irvine and
Tectonic setting of Proterozoic layered intrusions with reference to the Bushveld Complex
191
Fig. 4. Geological map of the Muskox area. The gravity anomaly northwest of the Muskox intrusion is believed to reflect the extension of the intrusion below younger cover rocks (from Irvine, 1970; Irvine and Baragar, 1972).
Smith (1967). The first, which is responsible for the larger part of the intrusion, gave rise to a sequence comprising cyclic units characterized by the crystallization order olivine, clinopyroxene, plagioclase and intercumulus orthopyroxene. In comparison, the second period of fractional crystallisation, which produced the upper 600 m, has an apparent crystallization order of olivine, orthopyroxene, clinopyroxene and plagioclase. Chromite is an accessory constituent of most olivine bearing rocks, but also is concentrated in two thin layers a t the contact of peridotite and orthopyroxenite layers in cyclic units 21 and 22 (Irvine, 1975) while titanomagnetite and ilmenite are only developed in granophyric gabbro in the uppermost differentiates of the intrusion. Irvine and Smith (1967) proposed that the cyclic units formed as a result of periodic injection of primitive, undifferentiated parental magma into the Muskox magma chamber. The abundance of ultramafic cumulates provides evidence that the residual magma in the chamber after crystallization of a cycle was virtually
192
G. von Gruenewaldt and R.E. Harmer
completely displaced by influxes of fresh primitive magma. Irvine and Smith (1967) suggest that the displaced magma was expelled to surface as a sequence of basaltic flows on top of the sandstones of the Hornsby Group, and that these have been removed by erosion prior to the deposition of the dolomitic sediments of the Dismal Lakes Group and extrusion of the Coppermine River basalts. Irvine (1970) emphasizes that the two crystallization sequences were produced from compositionally different liquids, which can not be related to each other by fractional crystallization. He demonstrated that the change from the one magma type to the other was gradual, but dismissed the idea that this gradual change was brought about a t depth beneath the Muskox intrusion. He considers the fine grained, marginal bronzite gabbro present in the feeder dyke and at the inward dipping, lower walls of the funnel-shaped intrusion to represent the parental liquid of the intrusion, with a composition equivalent to a silica-saturated tholeiitic basalt. This liquid is considered to have changed its crystallization order in response to contamination with sialic material, viz. the granophyric liquid produced by melting of rocks along the roof contact in response to heat released from the basaltic magma of the intrusion. The Muskox intrusion is emplaced into strongly metamorphosed and folded Early Proterozoic rocks of the Epworth Group, a remnant of the Coronation geosyncline (Fig. 4). The sediments of this group grade westward from almost undeformed shelf deposits, which rest unconformably on the 2300-2600 Ma Slave Province craton, to increasingly folded and thrusted miogeosynclinal sediments and a tectonized zone developed in a engeosynclinal assemblage. The western edge of this fold belt is intruded and metamorphosed by granitic rocks of the 1700-1900 Ma old Bear Province. All these basement rocks are overlain by gently inclined and little deformed supracrustal rocks of the Amundsen Basin. The lowest of these, quartz sandstones and minor carbonate rocks of the Hornsby Bay Group, were deposited prior to emplacement of the Muskox intrusion. Fragments of sandstone of this group are common as xenoliths in the granophyre close to the roof contact. The upper contact of the Muskox intrusion is situated close to an unconformity a t the base of the Hornsby Bay Group, but the contact transgresses downward such that granitic basement rocks constitute the roof in places. Irvine and Baragar (1972) conclude from the overall shape and relationship of the intrusion with its country rocks, that its present outcrop occupies the southern extremity of a graben or rift structure. This rift structure developed close to the junction at depth between the Slave Province to the east and the Bear Province to the west. These authors consider this junction to represent a zone of crustal weakness that was susceptible to rifting, and consequently controlled emplacement of the intrusion. The Duluth Complex
The Midcontinent Gravity High of central U.S.A. and southern Canada is closely associated with the Keweenawan Rift, a Late Proterozoic rift that was the
Tectonic setting of Proterozoic layered intrusions with reference to the Bushveld Complex
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site of extensive magmatic activity. This belt of igneous activity owes its origin to vertical fracturing and subsidence of the crust in response to the development of an intra-cratonic rift system which formed between 1200 and 900 Ma. Rifting was, however, arrested at an early stage (King and Zietz, 1971; Craddock, 1972). The magmatic activity includes the extensive Middle Keweenawan plateau basalts as well as associated intrusions of mafic, inter-mediate and felsic composition. The largest and best studied of these is the 1100 Ma Duluth Complex, an arcuate composite intrusion of troctolitic, anorthositic and gabbroic rocks which is exposed over an area of 4700 km2 in northeastern Minnesota (Fig. 5). Four major rock units comprise the complex (Weiblen and Morey, 1980). These are an early layered suite of oxide-rich olivine gabbro in the northern part of the complex known as Nathan’s layered series; an intermediate anorthosite series of laminated but unlayered gabbroic to troctolitic anorthosite; a later troctolite series consisting of sheets of layered troctolite to olivine gabbro, as well as several intrusions with differentiated sequences ranging from picrite and troctolite
194
G. von Gruenewaldt and R.E. Harmer
to granophyre; and lastly, a subordinate felsic series of intermediate to granitic rocks. The plagioclase-enriched nature of the Duluth Complex is related to the overall evolution of the Keweenawan magmas within the intracontinental rift by Miller and Weiblen (1990). They suggest that most Keweenawan basalt developed by fractional crystallization from high-Al olivine tholeiite parental magmas which were generated by partial melting of spinel lherzolite in the upper mantle beneath the rift zone. Early, syn-volcanic intrusions gave rise to the rocks of Nathan’s layered series. However, a large proportion of these primary magmas ponded a t the base of the crust where fractional crystallization resulted in the separation of plagioclase from co-precipitating mafic minerals because of their enhanced buoyancy in basaltic magma at high pressure. They envisage that plagioclaserich mushes were generated in this way. Some of these mushes could have mixed with influxes of hotter primitive melts and the resultant resorption of some of this plagioclase could have given rise to hyperfeldspathic magmas. As rifting progressed, the break-up and heating of the continental crust allowed these viscous, plagioclase-rich mushes and/or plagioclase-phyric hyperfeldspathic magmas to reach the upper crust and intrude beneath the volcanic pile of the North Shore Volcanic Group. The residence time of magmas in the lower crust is considered to have diminished with further rifting and break up of the crust, so that progressively less fractionated and relatively plagioclase crystal-poor, olivine tholeiites were emplaced into the Duluth magma chamber, where they subsequently differentiated to produce the troctolite series rocks.
THE BUSEIVELD COMPLEX AND RELATED MAGMATIC EVENTS
The Bushveld Complex comprises two compositional suites: the Rustenburg Layered Suite of ultramafic to mafic layered cumulates, and the Lebowa Granite Suite, a younger sequence of sheeted intrusive granites (Von Gruenewaldt et al., 1985). In contrast to the layered intrusions discussed above, some authors have proposed that the components of the Bushveld Complex were emplaced under compressional conditions (Hunter, 1974; Sharpe and Snyman, 1980; Hatton and Sharpe, 1988; Hatton and Von Gruenewaldt, 1990). In order to critically examine the arguments on which these proposals are based, the layered cumulates of the Bushveld Complex must be seen in relation to the known geological and geochemical data for the successions hosting the intrusion. Geological setting and age relations The Bushveld Complex lies largely within the confines of the Transvaal Basin (Fig. 6) in which a succession of sedimentary and volcanic rocks, up to 12 km thick was deposited (Fig. 7). Clastic sedimentation and extrusion of the Abel Erasmus basalt of the Wolkberg Group in protobasins of restricted lateral extent was
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Tectonic setting of Proterozoic layered intrusions with reference to the Bushveld Complex
197
followed by the widespread deposition of chemical sediments, mostly dolomitic limestone, chert and banded iron formation of the Chuniespoort Group (Eriksson and Clendenin, 1990). A thick succession of alternating shales and sandstones of the Pretoria Group rests unconformably on these chemical sediments. Volcanic activity occurred intermittently throughout the deposition of these sediments (Fig. 7) and culminated in the voluminous extrusion of andesitic and felsic lavas of the Dullstroom Formation and Rooiberg Group, respectively (Harmer and Von Gruenewaldt, 1991). The ages of the various volcanic components are poorly constrained; those of the Hekpoort Formation (Fig. 7) have been dated a t 2224 f21 Ma (Burger and Coertze, 1975; Armstrong, 1987), whilst the Dullstroom volcanics have yielded ages of 2089 & 26 Ma (Schweitzer, 1986). Deposition of the Pretoria Group sediments occurred in a n intracratonic basin, characterized by alluvial fans, fan-deltas, lacustrine deltas, lake basin and lake margin palaeoenvironments. A half-graben structural setting with a steep footwall to the south can account for all the sedimentological features observed within the Pretoria Group (Eriksson et al., 1991). The extensive Dullstroom-Rooiberg volcanicity was followed by the emplacement of first the Rustenburg Layered Suite and then the Lebowa Granite Suite of the Bushveld Complex (Fig. 7). The age of the Rustenburg Layered Suite is well constrained at 2061 f 27 Ma (Walraven et al., 1990) and, although field relations clearly indicate a younger age of the granites, their age of 2052 48 Ma cannot be resolved from that of the layered mafic rocks (Walraven e t al., 1987; Walraven, 1988). Closely associated with the Rustenburg Layered Suite are a variety of different sills emplaced into the sedimentary succession underlying the Bushveld Complex. Sharpe (1984) estimated that these sills locally represent an aggregate stratigraphic thickness of over 2.5 km and represent a volumetrically significant group of magmatic rocks. They have been subdivided into pre- and syn-Bushveld sills (Sharpe, 1981, 1984); the syn-Bushveld sills form part of the marginal suite of the Rustenburg Layered Suite. Both the intrusive units and their host rocks are little deformed. Rocks generally dip at less than 20" to the center of the basin, except along the ThabazimbiMurchison lineament, where dips steepen to more than 60". To the north of this lineament the Bushveld rocks transgress the enclosing sediments so that the larger part of the exposed northern lobe rests on A ch a e a n basement (Fig. 6). Along the northern margin of this lobe the layered mafic rocks are truncated and intensely deformed by the Palala shear zone which ceased activity before 1700 Ma (Barton and McCourt, 1983). Relationships in the Bushveld Complex correlate with those in the coeval Molopo Farms Complex in southern Botswana (Gould et al., 1987; Von Gruenewaldt et al., 1988) (Fig. l).This complex is also situated in the Kaapvaal Craton and intrudes rocks correlated with the Transvaal Sequence. The entire complex of about 13000 km2 is covered by a thin (up to 250 m) veneer of Phanerozoic sediments so that interpretation of its structural setting and the distribution of
198
G. von Gruenewaldt and R.E. Harmer
rock types is based entirely on limited geophysical and borehole information. Acid volcanic and intrusive rocks are' absent in this area. A prominent linear feature, the Kgomodikae lineament, a western continuation of the Thabazimbi-Murchison lineament, transects the complex. 73 the north ultramafic cumulates predominate, the rocks have steep dips, are intensely faulted and transgress northwards onto basement rocks. ?b the south, a well layered sequence of mafic cumulates overlies the ultramafic rocks, dips are generally shallower and rocks are less deformed. Three additional large mafic to ultramafic complexes, known as the Groenfontein, Tshane and Xade occurrences occur along the western margin of the Kaapvaal Craton (Fig. 1). Very little is known about them and their size is deduced largely from the distribution of positive gravity anomalies.
The Dullstroom Formation The lavas of the Dullstroom Formation were extruded onto sedimentary rocks of the Pretoria Group. Flows are predominantly intermediate in composition (basaltic andesite) with minor basalt, dacite and rhyolite. Impersistent sedimentary units, abundant pyroclastics along with rhyolite and basaltic andesite volcanics characterize the lower part of the formation. The upper stage consists of a thick sequence of mostly porphyritic basic as well as acid flows with only rare sedimentary or pyroclastic intercalations (Schweitzer, 1984). Although basaltic andesite and andesite compositions dominate this volcanic suite, two chemical sub-classes were recognised by Schweitzer (1986) on the basis of the T i 0 2 contents. The low-Ti02 (LTi) volcanics contain less then 0.75% Ti02 and also less than 11% FeO (total Fe expressed as FeO), compared to high-titanium (HTi) volcanics in which the T i 0 2 content is invariably above 1.5% and the FeO content in excess of 11%. LTi volcanics erupted alone in the lower 350 m, but LTi and HTi volcanics are intercalated throughout the greater part of the 1600 m succession. The HTi flows range in composition from basaltic andesite to andesite, while the LTi group have a greater compositional variation and range in composition from basaltic andesite to rhyolite.
The Rooiberg Group Volcanism continued with the eruption of the Rooiberg siliceous volcanics. The contact between the Rooiberg and Dullstroom volcanics is not preserved as the Rustenburg Layered Suite intervenes (the Rooiberg forming the roof, and the Dullstroom locally the floor of the layered mafic rocks). Present Rooiberg volcanic exposures lie entirely within the limits of the Bushveld Complex, except possibly for some highly sheared equivalents in the Koedoesrand Formation at the southern edge of the Palala shear zone (Fig. 5). This suggests that these volcanics must originally have covered an area of at least 55000 km2 and, with a typical stratigraphic thickness of between 3 and 5 km, an eruption volume in the region of 200000 km3 may be derived. As such the
Tectonic setting of Proterozoic layered intrusions with reference to the Bushveld Complex
199
Rooiberg Group ranks amongst the largest accumulations of siliceous volcanics known (Twist and French, 1984). Individual flows range from a few tens of metres to just under 400 m and are laterally extensive - some have been traced along strike for more than 40 km (Twist and Bristow, 1990). These siliceous volcanics show the contrasting features of both lavas (relative scarcity of pyroclastic features, flow contortions) and ash flows or ignimbrites (great lateral extent) and are termed rheoignimbrites by ?hist and Bristow (1990). Two fundamentally different chemical types of felsite occur: a high-magnesian variety (HMF, MgO >1.7%) which is restricted to the lower two units; and a more abundant low-magnesian type (LMF, MgO <1.0%) found throughout the succession. Flows of H M F are intercalated with those of LMF character (Twist, 1985). Besides having contrasting MgO contents, these felsite types differ in their concentrations of Zr, Nb, Y and total REE (all lower in HMF). A difference in ENd is clear evidence that these groups are not related through any closed system differentiation scheme (Twist and Harmer, 1987). In addition to this fundamental grouping, a volumetrically less significant compositional sub-group may be distinguished within the LMF having higher Ti02, Fe203* (>lo%) and P2 0 5 concentrations. Progressive chemical changes are noted through the rhyolite succession: Si02 and Z r in particular increase progressively and Ti02 decreases from base to top. Zircon does not appear to have fractionated, the regular increase of Zr with increasing Si02 in the felsites is in marked contrast to that seen in the Bushveld granites where early zircon precipitation rapidly depletes the magma in Zr (Twist and Harmer, 1987). The compositional relationships between the chemical sub-groups within the Rooiberg and Dullstroom volcanic successions are shown on an Alkalis-FeO*MgO plot in Fig. 8. While the high-Ti02 group of Dullstroom lavas is chemically distinct, the low-Ti02 group is geochemically similar to the Rooiberg Group HMF volcanics as defined from the Loskop Dam section (Schweitzer, 1986). This similarity suggests that volcanism was essentially continuous from Dullstroom through Rooiberg times. ?re-Bushveld sills
The pre-Bushveld sills are characterised by amphibolite assemblages with no unaltered orthopyroxene. Certain of these sills are truncated by the layered suite, clearly indicating the pre-Bushveld timing of their emplacement. They are distributed throughout the Transvaal Sequence, and are particularly abundant within the quartzites of the Daspoort and Lakenvallei Formations (Sharpe, 1984). Both tremolite and hornblende-bearing sills are found and the nature of the amphibole present is ascribed to the relative importance of burial versus thermal metamorphism, with the tremolite sills reflecting the effects of thermal metamorphism (Sharpe, 1984).
G. von Gruenewaldt and R.E. Harmer
200 FeO*
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Fig. 8. Ternary FMA plot to compare compositions of pre-Bushveld sub-suites of the Dullstroom Formation and the Rooiberg Group volcanics (after Schweitzer, 1986).
Syn-Bushveld sills and niarginal rocks The closest available natural analogues to potential parental magmas of the mafic cumulates of the Bushveld Complex are provided by the chilled marginal facies of the complex, as well as sills emanating from the magma chamber. Data on the field relations and geochemistry of these rocks have been provided by Cawthorn and co-workers in the western Bushveld and by Sharpe in the eastern Bushveld (Davies e t al., 1980; Cawthorn et al., 1981; Sharpe, 1981, 1984). The marginal facies components were divided by Sharpe (1981) into three successively intruded suites designated B1, B2 and B3. A further compositional group of ultramafic sills was termed UM. The Bl group consists largely of quenched micro-pyroxenitic and noritic sills. The dominant mineral in these rocks is orthopyroxene and experimental investigations have shown that these sills reproduce the crystallization sequence required to generate the orthopyroxene-dominated cumulates of the Lower and lower Critical Zones (Cawthorn and Davies, 1983). Petrographically, the characteristic features of these rocks are an abundance of orthopyroxene (reversely zoned) in various textural habits, the presence of 1 to 2% biotite, and the occasional presence of interstitial quartz and K-feldspar, frequently in granophyric intergrowth (Hatton and Sharpe, 1988). The B1 suite has an extremely unusual chemistry in being both silica- and magnesium-rich, with anomalously high concentrations of the incompatible elements K20, Rb, Zr and L R EE (Table 2). Initial 87Sr/86Srratios range between 0.7032 and 0.7057 (Harmer and Sharpe, 1985).
Tectonic setting of Proterozoic layered intrusions with reference to the Bushveld Complex
201
TABLE 2 Comparison of characteristic features of parental magma compositions of the Bushveld Complex (from Harmer and Sharpe, 1985) B1
B2 and B3
55 -56 12 -13 0.8- 0.9
50 -51 6 -8 0.1- 0.3
900-1100 270- 300 25- 46 150- 230 40- 60 40
200- 400 100- 150 1- 4 300- 350 20-50 4-12 30 6
10 0.703-0.7057
6 0.7065-0.7077
3 0.7059-0.7072
The essentially gabbroic rocks of the B2 and B3 groups are similar in chemistry and field relations. Plagioclase is present in excess of pyroxenes, and orthopyroxene and clinopyroxene are present in approximately equal amounts. B2 sills are the more abundant and are found in contact with the upper Critical Zone, whereas the B3 sills are confined to an area adjacent to the base of the Main Zone. Sharpe (1981) originally distinguished the B3 group by its coarser grain size, greater abundance of xenoliths and higher amounts of modal clinopyroxene. Subsequent determination of R E E concentrations (Harmer and Sharpe, 1985) confirmed the distinction in that the B3 group has lower LREE contents (i.e. a “flatter” chondrite-normalised pattern) than the B2 group and exhibits a pronounced positive Eu anomaly. Apart from the REE, the B2 and B3 gabbronorites are chemically similar and, in contrast to the B1 group have typical tholeiite basalt chemistry, low K20, Rb and Zr, and high N 2 0 3and Sr (Thble 2). Initial s7Sr/s6Sr ratios for B2 and B3 are indistinguishable, ranging from 0.7059 to 0.7077, and are significantly higher than those of the B1 group. Sills of the harzburgitic UM group are frequently found in close association with the B1 pyroxenite sills. A close relationship between these two types of sills has been inferred since the addition of approximately 70% olivine to the average B1 composition closely reproduces many of the chemical features of the UM group (Davies and Tredoux, 1985; Sharpe and Hulbert, 1985; Harmer and Sharpe, 1985).
Rustenburg Layered Suite Cumulates making up the Rustenburg Layered Suite may be divided into four broad compositional zones: the Lower Zone of orthopyoxenites and harzburgites; a compositionally complex Critical Zone made up of a lower feldspathic pyroxenite sub-zone, and an upper sub-zone where pyroxenite is subordinate to plagioclasebearing cumulates such as norite and anorthosite; a compositionally monotonous
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Main Zone of gabbronorite and anorthosite; and an Upper Zone of magnetitebearing gabbronorite, olivine-gabbro and olivine-diorite (e.g. Von Gruenewaldt et al., 1985). Spectacular ore mineral concentrations are found in the Critical (chrome, platinum metals) and Upper (vanadiferous magnetite) Zones. Extensive Sr isotope data are now available for most portions of the cumulate succession (Hamilton, 1977; Kruger and Marsh, 1982; Sharpe, 1985; Hatton et al., 1986; Kruger et al., 1987) and, used in conjunction with the marginal suite data, suggest that the cumulates developed from mixtures of essentially two compositionally distinct magmas: an ultramafic, “U”, liquid (= Bl); and a tholeiitic, “A’ liquid (B2 and B3) (Irvine and Sharpe, 1982). The Main Zone of the Complex represents a major intrusion of new magma above the Merensky Reef with a very high 87Sr/86Srratio of >0.708 (Kruger and Marsh, 1982; Sharpe, 1985). As no components of the marginal suite have such high ratios, the existence of an additional magma type has been postulated by Hatton (1989) who demonstrated distinct differences in concentrations of the R E E and chondrite normalized multielement patterns for rocks above and below the MerensKy Reef. Approximate mass balance calculations indicate that the composition of this additional magma is akin to that of a high-aluminium basalt (Hatton, 1989).
The Lebowa Granite Suite The granites of the Lebowa Suite comprise the last major component of the Bushveld Complex. Components of the Suite form a sheet of batholithic dimensions in excess of 5 km thick in the central areas of the Complex and cover an area exceeding 30000 km’. The granite sheet intruded above the mafic rocks, and at various localities is seen to cut both the felsites and the mafic rocks. Two main granite types predominate; the Neb0 Granite, a major unit of coarse-grained, hypersolvus, mildly alkalic granite; and a more evolved, sometimes aplitic variety, the Klipkloof Granite. Other varieties, such as the Bobbejaankop, Foothills and Lease Granites, appear to represent localised, hydrothermally altered derivatives of the Nebo-Klipkloof types (Kleeman and Twist, 1989). Several different facies (coarse, fine, porphyritic, albitized) of Klipkloof Granite occur and the volume ratio of Klipkloof to Nebo Granites generally increases upwards through the intrusion. The Neb0 Granite exhibits a well developed and fairly systematic mineralogical and chemical zonation, characterised (from base to top) by a decreasing modal plagioclase concomitant with increasing albite component in the plagioclase; decreasing hornblende and increasing biotite; and increasing quartz. These variations are also reflected in the geochemical trends, i.e. Si, K, R b increase, and Fe, Ti, Ca, P, Ba, Sr and Zr decrease upwards through the sheet (Kleemann and Twist, 1989). The entire granite mass was probably emplaced as a n unusually fluid, very hot (perhaps >900”C), relatively anhydrous (initial water content ~ 2 . 2 % restite-free ) magma. The granites exhibit all the classical features of mildly alkalic A-type magmas (Kleemann and Tnyist, op. cit.), such as the absence of restites, the
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presence of interstitial hornblende and biotite rather than muscovite, hypersolvus character, restricted composition (i.e. high in Si02 and low in CaO), linear trends on bivariate element plots, and enrichment of incompatible trace elements. The tectonic setting of the Bushveld Complex
In view of the high Si02-MgO character of the B1 group of marginal rocks and sills, some authors have noted the similarity of the B1 magmas to “boninites” developed in island arc environments (e.g. Irvine and Sharpe, 1982; Sharpe and Hubert, 1985; Hamlyn and Keays, 1986; Hatton and Sharpe, 1988). Barnes (1989) challenged this view and considers that the B1 pyroxenites are more likely to represent extensively contaminated komatiitic parental magmas than being primary mantle melts. Experimental data on the Bushveld B1 samples indicate that these magmas would have had olivine on the liquidus only a t pressures less than 4 kbars (Cawthorn and Davies, 1983) and hence cannot represent primary melts from peridotitic mantle. Cawthorn and Davies (1983) consider this as evidence that the magmas had assimilated siliceous crust. The Bl’s do have elevated initial s7Sr/s6Sr ratios relative to a primitive mantle reference value at 2060 Ma of ca. 0.702, which, considered along with the high incompatible element contents, is consistent with assimilation of crustal materials. However, the B2/B3 sills have even higher s7Sr/86Srratios (>0.706), coupled with depleted incompatible element concentrations, which cannot easily be explained by simple crustal contamination. Hatton and Sharpe (1988), however, reject simple contamination of the Bushveld magmas en route to the surface, and argue that the crustal component was added to the mantle source region through the subduction of sediments below the Kaapvaal Craton. This model requires that the Bushveld Complex was generated close to an active subduction zone. Eriksson et al. (1991) argue, however, that the ’Ikansvaal sedimentary basin developed in a n intracratonic halfgraben setting and there is no compelling evidence from the geological record for active subduction below the central Kaapvaal Craton a t 2.06 Ga as most crust forming processes had ceased by about 2.6 Ga (e.g. Hunter, 1974, 1991; De Beer and Eglington, 1991). If the Bushveld magmas are considered in the wider context of Proterozoic magmatism on the Kaapvaal Craton, it is found that apparent volcanic arc geochemical signatures are common (Crow and Condie, 1988; Crow and Condie, 1990; Harmer and Von Gruenewaldt, 1991). It is interesting to note that subduction processes have also been invoked for the Late Ac h a e a n (2.8 Ga.) Ventersdorp lavas (Crow and Condie, 1988). The trace element distributions in the volcanics of the Transvaal Sequence are summarised on Pearce (1983)-style MORB-normalised concentration diagrams in Fig. 9 (after Harmer and Von Gruenewaldt, 1991). All these diagrams reflect relative enrichments in the elements K to Th. Patterns for the Hekpoort and Dullstroom Formations are “spiky” through the element range ?a to Yb due to marked relative depletions at Nb, P and Ti. The consistency of the patterns for the Hekpoort volcanics from east and west exposures (separated
G. lion Gruenewaldt and R.E. Harmer
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MACHADODORP VOLCANICS (Silverton Formation)
100
m
8 10 ? Y
g 1 0.1
100
m
10
t0
1
8 Y
HEKPOORT ( W . N L )
ABEL ERASMUS
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by over 450 km) is significant. Such patterns have been interpreted to reflect the presence of a "subduction zone components" in magma source regions (e.g. Pearce, 1983; Thompson et al., 1984). Twist and Harmer (1987) demonstrate that the high-magnesian felsites (HMF) of the Rooiberg Group have the trace element characteristics of volcanic arc granites (using diagrams of Pearce et al., 1984), whereas the low-magnesian felsites (LMF) and the Bushveld granites have compositions typical of within-plate acid magmas (Fig. 10). Apparent "subduction zone" geochemical signatures thus appear to be a characteristic of many Kaapvaal Craton magmas, both basic and silicic, and apparently are not unique only to the Bushveld Complex, Several authors have noted that apparent subduction zone geochemical patterns - particularly the characteristic fractionation of Nb, Tm and Ti from the rare earth elements - may be imprinted to mantle melts through interaction
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Y+Nb Fig. 10. Plot of R b versus Y + Nb for acid magmatic units spatially related to the mafic components of the Bushveld Complex: basal rhyolites of the Dullstroom Formation (fine stipple); high MgO (heavier stipple) and low MgO (diagonal ruling inclined to the right) felsites of the Rooiberg Group; and granites of the Lebowa Granite Suite (diagonal ruling inclined to the left). Compositional fields for granites from syn-collision (SUN-COLG),volcanic arc (VAG) and within-plate (WPG) tectonic settings are from Pearce et al. (1984).
with siliceous crustal materials (Dupuy and Dostal, 1984; Thompson et al., 1984; Arculus, 1987). As a result, the trace element patterns of volcanics sampled in continental areas are likely to yield conflicting “tectonic signals” with respect to the discrimination of within-plate versus island arc settings (e.g. Arculus, 1987; Duncan, 1987). The dilliculty of distinguishing the effects of contamination by continental crust from primary subduction zone signatures is illustrated by the fact that the same geochemical data set for Kaapvaal Craton volcanics has been interpreted as reflecting both subduction related processes and crustal contamination (Crow and Condie, 1988, 1990; Condie and Crow, 1990). The voluminous magmatic package of the Dullstroom Formation, Rooiberg Group and Bushveld Complex comprises a compositionally bimodal assemblage of igneous rocks. Siliceous units followed by basaltic andesitic lavas in the Dullstroom are in turn followed by the siliceous Rooiberg volcanics. These are intruded by ultrabasic to basic magmas of the Rustenburg Layered Suite, then by granites of the Lebowa Suite. Bimodal volcanic successions are typically found in continental rift settings (Marsh, 1987). The huge volume of siliceous magma in this succession, some 350000 km3, can neither represent differentiates of the mafic magmas nor be direct derivatives from peridotitic mantle and must conceivably contain a large proportion of melted crustal material. Harmer and Von Gruenewaldt (1991) note that siliceous crustal melts compositionally equivalent to the rhyolitic magmas in this magmatic sequence (see Fig. lo), particularly those with volcanic arc trace element signatures (Dullstroom basal rhyolites; Rooiberg HMF), if mixed with primary mantle melts would produce the “subduction zone” trace element patterns observed in the ’li-ansvaal Basin volcanics and Bushveld B1 magmas.
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The consistency in the “subduction zone” trace element signature of volcanic successions preceding the Bushveld event, particularly in the Hekpoort lavas, which are distributed over a larger surface area than the Bushveld Complex, is considered evidence against magma contamination during passage through the continental crust (Harmer and Von Gruenewaldt, 1991). It is thus concluded that the subduction zone component was an integral part of the sub-Kaapvaal lithosphere (i.e. lower crust and upper mantle) during much of the Proterozoic. An Archaean T C H age ~ ~reported for the Rooiberg H M F (ca. 3.4 Ga; Twist and Harmer, 1987) may suggest that this component was introduced even earlier, possibly during generation of the Kaapvaal Craton. A “hybrid” lower crust/mantle zone as envisaged by Cox (1980), Hildreth and Moorbath (1988) or Arndt and Goldstein (1989) is capable of repeatedly producing “subduction-like” trace element fractionations in magmas widely distributed in space and time. To conclude, we do not regard the evidence of geochemical “subduction zone” signatures as persuasive as the known stratigraphic and geochronological evidence that the Transvaal Basin developed in an intracratonic setting, possibly in a half-graben, and conclude that the Bushveld Complex was emplaced in a rifted continental environment.
DISCUSSIONS AND CONCLUSIONS
Rifting within a continental environment is considered by many authors to be the dominant tectonic setting of Proterozoic layered intrusions. Hatton and Von Gruenewaldt (1990) attempt to attribute the formation of Archaean and Early Proterozoic layered intrusions to long term cyclicity of geological processes, such as the extended Wilson cycle, in which a repetitive cycle of continental rifting and continental collision is proposed. According to this model, large sheet-like layered intrusions such as the 2700 Ma old Stillwater and the 2050 Ma old Bushveld Complex formed under conditions of compression in the crust, immediately prior to collision events. In contrast, dyke-like layered intrusions such as the Great Dyke and Jimberlana formed during widespread rifting events of continental crust. The evidence presented in this chapter favours emplacement of the Bushveld Complex into an intracratonic environment in which rifting had occurred during deposition of the sediments of the Transvaal Sequence (Eriksson et al., 1991). Furthermore, age relations of the large Proterozoic layered intrusions do not display a systematic pattern and cannot be related to any longterm cyclicity of geological processes. A possible exception is the widespread occurrence of -2400 Ma old layered intrusions, which could have formed in response to the rifting ol a Late Archaean super-continent as suggested by Gail (1985). If layered intrusions form under conditions of rifting in the earth’s crust, as the available evidence seems to indicate, then the question arises of how the sheet-like nature of the Rustenburg Layered Suite of the Bushveld Complex and some other Proterozoic intrusions can be reconciled with a regional tensional
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environment? The answer to this question may be found in the mechanism of sill emplacement suggested by Roberts (1970). His mathematical presentation indicates that depression of the earth’s crust, due to subsidence of a sedimentary basin, will give rise to a zone of horizontal compression at depth. The depth of maximum horizontal compression depends on the size of the basin and will tend to coincide with the chord to the arc of the earth’s surface occupied by the basin. This zone of horizontal compression is considered by Roberts to be conducive to sill formation. In addition, he demonstrates that compressional stress conditions below the chord favours low-angle sheet intrusion. We envisage a situation in the evolution of the Transvaal Basin whereby basin development was initiated by continental rifting. Deposition of a n 8 km thick sedimentary sequence is required for the floor of a depository the size of the Transvaal Basin to reach the chord position (as depicted by Sharpe and Snyman, 1980, figs. 7 and lo), i.e. the position where maximum horizontal compression is to be expected. The tensional regime which prevails at the time is evidently dominant so that sedimentation continues with periodic extrusion of basaltic to andesitic magma. Through continued subsidence lower crustal regimes were depressed into high-temperature environments which eventually led to partial melting and the extrusion of the voluminous Dullstroom and Rooiberg volcanic sequences. This rapid addition of volcanics onto the existing basin material is thought to have expedited subsidence of the basin to such an extent that the original tensional regime, which largely dictated the extent of rifting and the speed of subsidence within the basin, was replaced by horizontal compressional forces near the chord position within the basin. The chord position was situated some distance below the interface of the clastic sediments of the Pretoria Group and the overlying volcanic sequence and therefore controlled both the level of maximum sill formation within the Transvaal Sequence and the level of emplacement of the Rustenburg Layered Suite (Sharpe and Snyman, 1980). The shallow, funnel- to canoe-shaped nature of certain large layered intrusions could be governed by shallow-angled shearing regimes beneath the level of maximum horizontal compression within or below large sedimentary basins. Dykelike feeders to large sheeted layered intrusion, e.g. as observed in the Muskox and Koillismaa intrusions, suggests that tensional conditions prevailed in the crust beneath the postulated chord position of the basin during emplacement of the magma. It is therefore concluded that flat, funnel to canoe-shaped and sheet-like layered intrusions form in shallow crustal environmenls during o r after rift-induced subsidence of the crust. Depending on the thickness of the volcano-sedimentary pile that accumulated, the size of the rift-induced basin, and the prevailing chord position, such intrusions may be emplaced within the sedimentary pile or in basement rocks beneath such basins. The composition of the postulated parental magmas of layered intrusions varies considerably from one intrusion to the next, and perusal of Table 1 shows a spectrum of compositions from high-MgO basalts to aluminous tholeiites. High-MgO
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basalts and “boninite-like” parental magmas seem restricted to Early Proterozoic intrusions, e.g. the 2400 Ma dyke-like intrusions and the earlier B1-type parental magmas of the Bushveld Complex. Such magmas, which gave rise to the so-called “noritic suites” of Hall and Hughes (1990), were apparently widespread in the Early Proterozoic. They (ibid, 1990) debated the origin of these noritic suites at length and concluded that their large volume and widespread occurrence in the Early Proterozoic, as well as their distinctive chemical composition and tectonic setting does not support derivation through crustal contamination of komatiitic magma. Instead, temporal restriction of these noritic suites to the Early Proterozoic is thought to be related to a major Late Archaean crust-forming event and thickening of the crust during the Early Proterozoic facilitated the generation of high-Mg magma by melting of metasomatized harzburgite source material. Evidence for concomitant tholeiitic magmatism is especially well documented for the Bushveld Complex, but there are also indications that two magma types were involved in the formation of the Penikat and related layered intrusions of northern Finland and the adjoining Kola Peninsula. This is especially significant, as mixing of two different types of magma within layered intrusions is now widely accepted as an important mechanism for the formation of mineralized layers in such intrusions (Irvine et al., 1983; Sharpe and Irvine, 1983; Campbell et al., 1983; Hatton et al., 1986; Naldrett and Von Gruenewaldt, 1989). From the above discussion and the description of the various layered intrusions, it is evident that the composition of their parental magmas is largely dictated by the nature of the crust-mantle transition beneath the intrusion and long term processes that have modified the upper mantle prior to generation of these magmas. Virtually all types of basaltic magma can give rise to layered intrusions within shallow crustal environments. Layering can be enhanced by processes such as periodic influxes of undifferentiated magma and magma mixing, particularly between influxes of parental magmas of contrasting composition. Under such situations a density stratification of liquids within the magma chamber can develop, which will further enhance the development of layering in the intrusions. ACKNOWLEDGEMENTS
We are indebted to M. Knoper for constructive criticism of the manuscript, as well as to M. van Leeuwen and R. Kuschke for drafting of the diagrams and typing of the manuscript respectively. Constructive reviews by Kent Condie, Norman Page and Don Hunter improved the final form of the manuscript and are greatly appreciated. REFERENCES Alapieti, TT, 1982. The Koillismaa layered igneous complex, Finland - its structure, mineralogy and geochemistry, with emphasis on the distribution of chromium. Geol. Sum. Fin]., Bull., 319, 116 pp.
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Alapieti, T?: and Piirainen, T, 1984. Ni-Cu-PGE mineralization in the marginal series of the early Proterozoic Koillismaa layered igenous complex, northeast Finland. In: D.L. Buchanan and M.J. Jones (Editors), Sulphide Deposits in Mafic and Ultramfic Rocks. Inst. Min. Metall., London: 123131. Alapieti, TT and Lahtinen, J.J., 1986. Stratigraphy, petrology and platinum-group element mineralization of the early Proterozoic Penikat layered intrusion, northern Finland. Econ. Geol., 81, 11261136. Alapieti, TT, Filen, B.A., Lahtinen, J.J., Lavrov, M.M., Smolkin, V.E and Voitsekhovsky, S.N., 1990. Early Proterozoic layered intrusions in the northeastern part of the Fennoscandian Shield. Min. Petrol., 42, 1-22. Arculus, R.J., 1987. The significance of source versus process in the tectonic controls of magmas genesis. J. Volcanol. Geotherm. Res., 32, 1-12. Armstrong, R.A., 1987. Geochronological studies on Archaean and Proterozoic formations of the foreland of' the Namaqualand front and possible correlations on the Kaapvaal Craton. Ph.D. Thesis, University of t h e Witwatersrand, 274 pp. (unpublished). Arndt, N.T and Goldstein, S.L., 1989. An open boundary between lower continental crust and mantle: its role in crust formation and crustal recycling. Tectonophysics, 161: 201-212. Baragar W.R.A. and Scoates, R.F.J., 1981. The circum-superior Belt: a Proterozoic plate margin? In: A. Kroner (Editor), Precambrian Plate Tectonics. Elsevier, Amsterdam, pp. 297-330. Barnes, S.J., 1989. Are Bushveld U-type parent magmas boninites or contaminated komatiites? Contrib. Mineral. Petrol., 101: 447-457. Barton, J.M. and McCourt, S., 1983. Rb-Sr age for the Palala granite, Limpopo mobile belt. Geol. SOC. S. Afr., Spec., Publ., 8 45-46. Berg, A.H., 1977. Regional geobarometry in the contact aureoles of the anorthositic Nain Complex, Labrador. J. Petrol., 18: 399-430. Burger, A.J. and Coertze, EJ., 1975. Age determinations - April 1972 to March 1974. Ann. Geol. Sum. S. Afr., 1 0 135-141. Campbell, I.H., 1977. A study of macro-rhythmic layering and cumulate processes in the Jimberlana intrusion. J. Petrol., 18: 183-215. Campbell, I.H., McCall, G.J.H. and Tyrwhitt, D.S., 1970. The Jimberlana norite, Western Australia - a smaller analogue of the Great Dyke OP Rhodesia. Geol. Mag., 107: 1-11. Campbell, I.H., Naldrett, A.J. and Barnes, S.J., 1983. A model for the origin of platinum-rich sulfide horizons in the Bushveld and Stillwater complexes. J. Petrol., 24: 133-165. Cawthorn, R.G. and Davies, G., 1983. Experimental data at 3 kbars pressure on parental magma to the Bushveld Complex. Contrib. Mineral. Petrol., 8 3 128-135. Cawthorn, R.G., Davies, G., Clubley-Armstrong, A.R. and McCarthy, TS., 1981. Sills associated with the Bushveld Complex, South Africa: an estimate of parental magma composition. Lithos, 14: 1-15. Condie, K.C. and Crow, C., 1990. Early Precambrian within-plate basalts from the Kaapvaal Craton in southern Africa: a case for crustally contaminated komatiites. J. Petrol., 98: 100-107. Cox, K.G., 1980. A model for flood basalt volcanism. J. Petrol., 21: 629-650. Craddock, C., 1972. The regional geologic setting of the Late Precambrian. In: P.K. Sims and G.B. Morey (Editors), Geology of Minnesota: A Centennial Volume. Minn. Geological Survey, pp. 281291. Crow, C. and Condie, K.C., 1988. Geochemistry and origin of late Archaean Volcanics from the Ventersdorp Supergroup, South Africa. Precambrian Res., 4 2 19-42. Crow, C. and Condie, K.C., 1990. Geochemistry and origin of early Proterozoic volcanic rocks from the Transvaal and Soutpansberg successions, South Africa. Precambrian Res., 47: 17-26.
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Davies, G. and Tredoux, M., 1985. The platinum-group element and gold contents of the marginal rocks and sills of the Bushveld Complex. Econ. Geol., 80: 838-848. Davies, G., Cawthorn, R.G., Barton, J.M. and Morton, M., 1980. Parental magma to the Bushveld Complex. Nature, 287 33-35. De Beer, J.H. and Eglington, B.M., 1991. Archaean sedimentation on the Kaapvaal craton in relation to tectonbm in the granite-greenstone terrains: geophysical and geochronological constraints. J. Afr. Earth Sci., 13: 27-44. Duncan, A.R., 1987. The Karoo igneous province - a problem area for inferring tectonic setting from basalt geochemistry. J. Volcanol. Geotherm. Res., 3 2 13-34. Dupuy, C. and Dostal, J., 1984. %ace element geochemistry of some continental tholeiites. Earth Planet. Sci. Lett., 67: 61-69. Eriksson, P.G. and Clendenin, C.W., 1990. A review of the Transvaal Sequence, South Africa. J. Afr. Earth Sci., 10: 101-116. Eriksson, P.G., Schreiber, U.M. and Van der Neut, M., 1991. A review of the sedimentology of the early Proterozoic Pretoria Group, South Africa: Implication for tectonic setting. J. Afr. Earth Sci., 13: 107-119. Gail, G., 1985. Nickel metallogeny related to tectonics. Geol. Surv. Finl., Bull., 333: 143-155. Gould, D., Rathbone, P.A. and Kimbell, G.S., 1987. The geology of the Molopo Farms and adjacent areas, southern Botswana. Botswana Geol. Surv. Dept., Bull., 23, 178 pp. Hall, R.P. and Hughes, D.J., 1990. Noritic magmatism. In: R.P. Hall and D.J. Hughes (Editors), Early Precambrian Basic Magmatism. Blackie and Son, London, pp. 83-110. Hamilton, P.J., 1977. Sr isotopic and trace element studies of the Great Dyke and Bushveld mafic phase and their relation to early Proterozoic magma genesis in southern Africa. J. Petrol., 18: 24-52. Hamlyn, P.R. and Keays, R.R., 1986. Sulphur saturation and second stage melts: application to the Bushveld platinum metal deposits. Econ. Geol., 81: 1431-1445. Harmer, R.E. and Sharpe, M.R., 1985. Field relations and strontium isotope systematics of the marginal rocks of the eastern Bushveld Complex. Econ. Geol., 8 0 813-837. Harmer, R.E. and Von Gruenewaldt, G., 1991. A review of magmatism associated with the Transvaal basin - implications for its tectonic setting. J. Geol. S. Afr., 93: 104-122. Hatton, C.J., 1989. Densities and liquidus temperatures of Bushveld parental magmas as constraints on the formation of the Merenshy Reef in the Bushveld Complex, South Africa. In: M.D. Prendergast and M.J. Jones (Editors), Magmatic Sulphides - The Zimbabwe Volume. Inst. Min. Metall., London, pp. 87-94. Hatton, C.J. and Sharpe, M.R., 1988. Significance and origin of boninite-like rocks associated with the Bushveld Complex. In: A.J. Crawford (Editor), Boninites. Unwin Hyman, London, pp. 174-207. Hatton, C.J. and Von Gruenewaldt, G., 1990. Early Precambrian layered intrusions. In: R.P. Hall and D.J. Hughes (Editors), Early Precambrian Basic Magmatism. Blackie and Son, London, pp. 56-82. Hatton, C.J., Harmer, R.E. and Sharpe, M.R., 1986. Petrogenesis of the middle group of chromitite layers: Doornvlei, eastern Bushveld Complex. In: M.J. Gallagher, R.A. Ixer, C.R. Neary and H.M. Prichard (Editors), Metallogeny of Basic and Ultrabasic Rocks. Inst. Min. Metall., London, pp. 241247. Hildreth, W. and Moorbath, S., 1988. Crustal contributions to arc magmatism in the Andes of Central Chile. Contrib. Mineral. Petrol., 98: 455-489. Hughes, C.J., 1976. Parental magma of the Great Dyke of Southern Rhodesia - voluminous late Archaean high magnesium basalt. Trans. Geol. SOC.S. Afr., 79: 171-182. Hunter, D.R., 1974. Crustal development in the Kaapvaal Craton, 11. The Proterozoic. Precambrian Res., 1:295-326.
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Hunter, D.R., 1991. Crustal processes during Archaean evolution of the southeastern Kaapvaal province. J. Afr. Earth Sci., 13: 13-25. Irvine, TN., 1970. Crystallization sequences in the Muskox intrusion and other layered intrusions, I. Olivine-pyroxene-plagioclase relations. Geol. SOC.S . Afr., Spec. Publ., 1: 441-476. Irvine, TN., 1975. Crystallization sequences in the Muskox intrusion and other layered intrusions, 11. Origin of chromitite layers and similar deposits of the other magmatic ores. Geochim. Cosmochim. Acta, 39: 991-1020. Irvine, TN. and Smith, C.H., 1967. The ultramafic rocks of the Muskox intrusion, Northwest Territories, Canada. In: P.J. Wyllie (Editor), Ultramafic and Related Rocks. John Wiley and Sons, New York, N.Y., pp. 38-49. Irvine, TN. and Baragar, W.R.A., 1972. Muskox intrusion and Coppermine River Lavas, Northwestern Territories, Canada. 24th Int. Geol. Congr., Montreal, Excursion A29, Guidebook, 70 pp. Irvine, T.N. and Sharpe, M.R., 1982. Source rock compositions and depths of origin of Bushveld and Stillwater magmas. Carnegie Inst. Washington, Yearb., 81: 294-303. Irvine, T.N., Keith, D.W. and Todd, S.G., 1983. The J-M platinum-palladium reef of the Stillwater Complex, Montana, 11. Origin by double-diffusive convective magma mixing and implications for the Bushveld Complex. Econ. Geol., 7 8 1287-1334. King, E.R. and Zietz, J., 1971. Aeromagmatic study of the Midcontinent Gravity High of central United States. Geol. SOC.Am., Bull., 82: 2187-2208. Kleemann, G. and 'hist, D., 1989. The compositionally-zoned sheet-like granite pluton of the Bushveld Complex: Evidence bearing on the nature of A-type magmatism. J. Petrol., 30: 1383-1414. Kruger, EJ. and Marsh, J.S., 1982. Significance of "SrP'Sr ratios in the Merensky cyclic unit of the Bushveld Complex. Nature, 298: 53-55. Kruger, EJ., Cawthorn, R.G. and Walsh, K.L., 1987. Sr-isotopic evidence against magma addition in the Upper Zone of the Bushveld Complex. Earth Planet. Sci. Lett., 84: 51-58. Lahtinen, J., 1985. PGE-bearing copper-nickel occurrences in t h e Marginal Series of the Early Proterozoic Koillismaa layered intrusion, northern Finland. Geol. Surv. Finl.., Bull., 333 161-178. Marsh, J.S., 1987. Basalt geochemistry and tectonic discrimination within continental flood basalt provinces. In: S.D. Weaver and R.W. Johnson (Editors), Tectonic Controls on Magma Chemistry. J. Volcanol. Res., 3 2 35-50. McClay, K.R. and Campbell, I.H., 1976. The structure and shape of the Jimberlana intrusion, Western Australia as indicated by a combined geological and geophysical investigation of the Bronzite Complex. Geol. Mag., 9 6 75-80. Meixner, H.M. and Peart, R.J., 1984. The Kalahari Drilling Project. Geol. Surv. Botswana, Bull., 27, 224 PP. Miller, J.D. and Weiblen, P.W., 1990. Anorthositic rocks of the Duluth Complex: Examples of rocks formed from plagioclase crystal mush. J. Petrol., 31: 295-339. Morse, S.A., 1969. Geology of the Kiglapait layered intrusion, Labrador. Geol. SOC.Am., Mem., 112. Morse, S.A., 1979. Kiglapait geochen~istry,I. Systematics, sampling and density. J. Petrol., 20: 555-590. Morse, S.A., 1981. Kiglapait geochemistry, IV. The major elements. Geochim. Cosmochim. Acta, 45: 461-479. Mutanen, T, 1989. Koitelainen intrusion and Keivitsa-Satovaara Complex. Geol. Surv. Finl., Guide 28: 49 PP. Mutanen, T, Tornroos, R. and Johanson, B., 1987. The significance of cumulus chlorapatite and high temperature dashkesanite to the genesis of PGE mineralization in the Koitelainen and KeivitsaSatovaara complexes, northern Finland. In: H.M. Prichard, P.J. Potts, J.EW. Bowles and S.J. Cribb (Editors), Geo-Platinum 87. Elsevier, London, pp. 159-160 (abstract).
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Naldrett, A.J., 1989. Ores associated with flood basalt. In: J.A. Whitney and A.J. Naldrett (Editors), Ore Deposits Associated with Magmas. Rev. Econ. Geol., 4 103-118. Naldrett, A.J. and Von Gruenewaldt, G., 1989. Association of platinum-group elements with chromitite in layered intrusions and ophiolite complexes. Econ. Geol., 84 180-187. Parsons, I. (Editor), 1987. Origins of Igneous Layering. NATO Adv. Sci. Inst., Ser. C, 196, D. Reidel, Dordrecht, 666 pp. Pearce, J.A, 1983. Role of the sub-continental lithosphere in magma genesis at active continental margins. In: C.J. Hawkesworth and M.J. Norry (Editors), Continental Basalts and Mantle Xenoliths. Shiva, Nantwich, pp. 230-249. Pearce, J.A., Harris, N.B.W. and Tindle, A.G., 1984. Trace element discrimination for the tectonic interpretation of granitic rocks. J. Petrol., 25: 956-983. Podmore, E and Wilson, A.H., 1987. A reappraisal of t h e structure, geology and emplacement of the Great Dyke, Zimbabwe. Geol. Assoc. Can., Spec. Pap., 34: 317-330. Prendergast, M.D., 1987. The chromitite ore field of the Great Dyke, Zimbabwe. In: C.W. Stowe (Editor), Evolution of Chromium Ore Fields. Van Nostrand Reinhold, New York, N.Y., pp. 89-108. Pye, E.G., Naldrett, A.J. and Giblin, P.E. (Editors), 1984. Geology and Ore Deposits of the Sudbury Structure. Ont. Geol. Surv., Spec. Vol., 1, 603 pp. Roberts, J.L., 1970. The intrusion of magma into brittle rocks. In: G. Newall and N. Rast (Editors), Mechanism of Igneous Intrusion. Liverpool Manchester J. Geol., Spec. Iss., 2 287-338. Schweitzer, J., 1984. The Dullstroom volcanics and their relations to the Rooiberg Felsite. Inst. Geol. Res. Bushveld Complex, University of Pretoria, Annu. Rep., 1983: 52-58. Schweitzer, J., 1986. The geochemical transition from the Dullstroom Basalt Formation to the Rooiberg Felsite Group. Inst. Geol. Res. Bushveld Complex, Annu. Rep., 1985: 72-81. Scoates, R.EJ., 1984. The Fox River Sill, northeastern Manitoba - a subvolcanic intrusion. Geol. Assoc. Can., Mineral. Assoc. Can., Prog. Abstr., 9: 103. Scoates, R.EJ. and Eckstrand, O.R., 1986. Platinum-group elements in the upper central layered zone of the Fox River Sill, Northeastern Manitoba. Econ. Geol., 81: 1137-1146. Sharpe, M.R., 1981. The chronology of magma influxes to the eastern compartment of the Bushveld Complex as exemplified by its marginal border groups. J. Geol. Soc., London, 138: 307-326. Sharpe, M.R., 1984. Petrography, classification and chronology of mafic sill intrusions beneath the eastern Bushveld Complex. Geol. Surv. S. Afr., Bull., 77, 40 pp. Sharpe, M.R., 1985. Strontium isotopic evidence for preserved density stratification in the main zone of the Bushveld Complex, South Africa. Nature, 316: 119-126. Sharpe, M.R. and Snyman, J.A. 1980. A model for the eniplacement of the eastern compartment of the Bushveld Complex. Tectonophysics, 6 5 85-110. Sharpe, M.R. and Irvine, TN., 1983. Melting relations ol two Bushveld chilled margin rocks and implications for the origin of chromitite. Carnegie Inst. Washington, Yearb., 82: 295-300. Sharpe, M.R. and Hulbert, L.J., 1985. Ultramafic sills beneath the eastern Bushveld Complex: Mobilized suspensions of early lower zone cumulates in a parental magma with boninitic affinities. Econ. Geol., 80: 849-871. Smith, C.H., 1962. Notes on the Muskox intrusion, Coppermine River area, Northwest Territories, Canada. Geol. Surv. Can., Pap., 61-25, 16 pp. Smith, C.H. and Kapp, H.E., 1963. The Muskox intrusion, a recently discovered layered intrusion i n the Coppermine River area, Northwest Territories, Canada. Mineral. SOC.Am., Spec. Pap., 1: 30-35. Taylor, EC., 1971. A revision of Precambrian structural province in northeastern Quebec and northern Labrador. Can. J. Earth Sci., 8: 579-584. Thompson, R.N., Morrison, M.A., Hendry, G.L. and Parry, S.J., 1984. An assessment of the relative roles of crust and mantle in magma genesis: an elemental approach. Philos. Trans. R. Soc. London,
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Ser. A, 310: 549-590. Twist, D., 1985. Geochemical evolution of the Rooiberg silicic lavas in the Loskop Dam area, southeastem Bushveld. Econ. Geol., 80: 1153-1165. Twist, D. and French, B.M., 1984. Voluminous acid volcanism in the Bushveld Complex: a review of the Rooiberg Felsite. Bull. Volcanol., 46: 225-242. Twist, D. and Harmer, R.E., 1987. Geochemistry of contrasting siliceous magma suites in the Bushveld Complex: Genetic aspects and implications for tectonic discrimination diagrams. J. Volcanol. Geotherm. Res., 3 2 83-98. Twist, D. and Bristow, J.W, 1990. Extensive lava-like siliceous Rows in Southern Africa: A review of occurrences. Inst. Geol. Res. Bushveld Complex, Univ. Pretoria, Res. Rep., 82, 35 pp. Van Biljon, W.J. and Legg, J.H. (Editors), 1983. The Lirnpopo Belt. Geol. Soc. S. Afr., Spec. Publ., 8, 203 pp. Von Gruenewaldt, G., Behr, S.H. and Wilheim, H.J., 1988. Some preliminary petrological investigations of the Molopo Farms Complex, Botswana, and its Ni-Cu sulphide mineralization. In: M.D. Prendergast and M.J. Jones (Editors), Magmatic Sulphides - The Zimbabwe Volume. Inst. Min. Metall., London, pp. 95-105. Von Gruenewaldt, G., Sharpe, M.R. and Hatton, C.J., 1985. The Bushveld Complex: Introduction and review. Econ. Geol., 80: 803-812. Vuorelainen, Y., Hakli, TA., Hanninen, E., Papunen, H., Reino, J. and Tornroos, R., 1982. Isomertieite and other platinum-group minerals from the Konttijarvi layered mafic intrusion, northern Finland. Econ. Geol., 77: 1511-1518. Wager, L.R. and Brown, G.M. 1968. Layered Igneous Rocks. Oliver and Boyd, Edinburgh, 588 pp. Weiblen, P.W. and Morey, G.B., 1980. A summary of the stratigraphy, petrology and structure of the Duluth Complex. Am. J. Sci., 280-A: 88-133. Walraven, E, 1988. Notes on the age and genetic rclationships of the Makhutso Granite, Bushveld Complex, South Africa. Chem. Geol. (Isot. Geosci. Sect.), 7 2 17-28. Walraven, E, Retief, E.A., Burger, A.J. and Swanepoel, D.J., 1987. Implications of new U-Pb zircon age dating on the Neb0 Granite of the Bushveld Complex. J. S. Afr. Geol., 90: 344-351. Walraven, E, Armstrong, R.A. and Kruger, EJ., 1990. A chronostratigraphic framework for the northcentral Kaapvaal Craton, the Bushveld Complex and the Vredefort structure. Tectonophysics, 71: 23-48. Wiles, J.W., 1968. Some aspects of the metamorphism of the Basement Complex in the Sipolilo district. Trans. Geol. Soc. S. Afr., 71 (Annexure): 71-88. Wilson, A.H., 1982. The geology of the Great Dyke, Zimbabwe: The ultramafic rocks. J. Petrol., 23: 240-292. Wilson, A.H. and Prendergast, M.D., 1989. The Great Dyke of Zimbabwe, I. Tectonic setting, stratigraphy, petrology, structure, emplacement and crystallization. In: M.D. Prendergast and M.J. Jones (Editors), Magmatic Sulphides - The Zimbabwe Volume. Inst. Min. Metall., London, pp. 1-20. Worst, B.G., 1960. The Great Dyke of Southern Rhodesia. S. Rhod. Geol. Surv., Bull., 47, 234 pp.
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Chapter 6
PROTEROZOIC ANORTHOSITE COMPLEXES R.A. WIEBE
INTRODUCTION
Anorthosite massifs (up to 1000’s of km2 in area) appear to represent a unique magmatic episode in the mid-Proterozoic. They have many characteristics that distinguish them from intrusive complexes emplaced either in the Archean or Phanerozoic. Of primary importance is the dominance of massive to weakly layered anorthositic plutons containing between 75% and 95% plagioclase of intermediate composition (typically An60 to A n 4 0 ) . Within these plutons orthopyroxene or olivine with intermediate compositions [Mg# = 100Mg/(Mg + Fe) = 70-401 are the most common mafic minerals. The anorthositic rocks are characteristically very coarse-grained with unzoned or weakly zoned subhedral plagioclase commonly between 1 and 10 cm and rarely up to 1 m in length. Although many complexes contain large volumes of associated granitic rocks, mafic rocks (including layered gabbroic intrusions) rarely comprise more than 10%. Some of the associated mafic rocks in most massif anorthosite complexes are Fe- and Ti-rich diorites or jotunites; Fe-Ti oxide deposits are common in many complexes. Ultramafic rocks are essentially absent except for small isolated bodies of oxide-rich cumulates associated with mafic rocks. The distinctive compositions of anorthosite massifs and their restricted ages (typically between 1000 and 1700 Ma) have long suggested that they should provide important clues to the evolution of the crust and mantle during the Proterozoic. Although there appears to be a growing consensus about some aspects of anorthosite genesis (e.g., the involvement of basalt, underplating of basaltic magma beneath stable continental crust, and concentration of plagioclase by fractional crystallization), many questions and controversies remain about the nature of the mantle sources, the involvement of crust, the physical state of the magmas during emplacement, and the roles that Proterozoic mantle processes and lithospheric structure played in the generation of the anorthosite massifs. A major focus of much current research is on isotopic studies directed toward understanding the sources of the magmas that produced these rocks. Although there have been many excellent reviews of anorthosites in the last several years (Emslie, 1978, 1980,1985; Morse, 1982a; Duchesne, 1984; Leelanandam, 1987; Ashwal, in prep.), no volume attempting to characterize the Proterozoic and its evolution would be complete without a consideration of anorthosite massifs.
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This review focuses on field relations within anorthosite complexes - partly because of my experience and partly to redress a common tendency to gloss over field relations in favor of other essential, more readily quantifiable, data (e.g., geochemistry, isotopes). I hope to provide enough detail about the internal structure and composition of these complexes so that a non-petrologist has some basis for comparing them with other, better known, plutonic masses (e.g., calc-alkaline granitic batholiths). Another focus of this review is on individual plutons (magma batches) as basic units within anorthosite complexes. It is important to understand the internal structure and compositions of individual plutons within complexes and their relations to one another before attempting to explain the broader significance of the anorthosite complexes as a whole to magmatic processes in the Proterozoic.
COMPOSITION AND ROCK NOMENCLATURE
The dominant plagioclase-rich plutons found in anorthosite massifs typically have 75-95% plagioclase with varying proportions of olivine, orthopyroxene (including inverted pigeonite) and augite. Other minerals (e.g., ilmenite, magnetite, hornblende, quartz) generally only occur as accessory phases. I will try to follow closely the rock classification of Streckeisen (1976) in which anorthosite is restricted to rocks with more than 90% plagioclase (color index, CI = 0-10). Although rocks with CI = 10-35 are broadly anorthositic, they are more precisely termed leuconorite, leucogabbro or leucotroctolite depending whether the dominant mafic mineral is, respectively, orthopyroxene, augite or olivine. Some mafic rocks (CI = 35-65) with similar mineralogy occur in many anorthosite complexes; norites, troctolites and gabbronorites all may be present. Many small bodies of fine- to medium-grained, Fe-rich, mafic rocks occur in most Proterozoic anorthosite complexes. These generally contain plagioclase in the andesine range with Fe-rich pyroxenes and olivine dominant over hornblende or biotite. They also commonly contain abundant ilmenite and/or magnetite. These rocks are variably termed diorites (Wiebe, 1990a), ferrodiorites (Emslie, 1978) or jotunites (De Waard, 1970). Where subordinate alkali-feldspar is also present, the rocks have been termed monzonorites (Duchesne, 1990). For the sake of simplicity, all of these Fe-rich rocks wiII generally be referred to as diorites. Major bodies of K-rich granitic rocks are very commonly associated with the anorthosite massifs. These may include granites, monzonites and syenites in the sense of Streckeisen (1976). Because many of these rocks are relatively anhydrous and bear orthopyroxene, granitoid rocks in some complexes have been termed charnockite and mangerite (Streckeisen, 1974).
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SIZE AND SHAPE OF ANORTHOSITE MASSIFS
Anorthosite massifs vary widely in size - from complexes larger than 5000 km2 such as Lac Saint-Jean, Harp Lake, and Nain in Canada (Fig. 1)to occurrences as small as a few km2. More than 35 complexes larger than 500 km2 occur throughout the world (Ashwal, in prep.). Where the complexes are not highly deformated and where exposures permit, the larger massifs can generally consist of many separate plutons (e.g. Emslie, 1980; Duchesne et al., 1985; Wiebe, 1990b). In this respect they resemble the more familiar calc-alkaline granitic batholiths. Although the massifs are often roughly equant in shape, no regular form should be expected since most larger ones are made up of many separate plutons. Many massifs such as the Morin (Martignole and Schrijver, 1970) have shapes that were probably modified by subsequent deformation. Gravity studies of most massifs suggest that they are thin, plate-like bodies. A study by Smithson and Ramberg (1979) of the Egersund massif (South RogaIand, Norway) suggests that this roughly 1000 km2 complex is only about 4 km thick. These thicknesses, however, are strongly dependent upon the assumed average color index (and hence, density) of the anorthosite (Morse, 1982a) and the assumed densities of the crustal envelope.
Fig 1. Distribution of Proterozoic anorthosite massifs and associated granitic rocks in northeastern Canada and the United States. Names of some major complexes include: 1 = Nain; 2 = Harp Lake; 3 = Michikamau; 4 = Mealy Mountains; 5 = Lac Fournier; 6 = Lac Allard; 7 = Lac Saint-Jean; 8 = Morin; 9 = Adirondacks. Modified from Emslie and Hunt (1990).
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Nonetheless, both gravity (Hodge et al., 1973) and seismic reflection profiling (Smithson et al., 1977) suggest that the Laramie anorthosite complex in Wyoming (roughly 1000 km2 in extent) is no more than 6 km thick. Similar slab-like forms satisfy both gravity and seismic data obtained at other massifs. These same gravity studies also indicate a scarcity (relative to basalt and cogenetic cumulates) of mafic and ultramafic rocks near the present level of exposure. They do permit, however, the existence of substantial mafic to ultramafic material at depth in the lower crust or uppermost mantle (Morse, 1982a). Positive gravity anomalies associated with some complexes in eastern Canada may indicate the presence of subjacent mafic cumulates ('Einner, 1969), and mafic intrusions occur at the present level of exposure of some anorthosite massifs (e.g., the Kiglapait intrusion in the Nain complex - Morse, 1969). COMPARISON WITH ARCHEAN AND OTHER ANORTHOSITES
The Proterozoic massif-type anorthosites are distinct in many ways from the Archean anorthosites. The latter typically have highly calcic plagioclase associated with magnesian augite or hornblende. Archean anorthosites commonly occur as stratiform bodies within layered mafic intrusions and may be associated with chromite-rich cumulates (Subramanian, 1956; Windley, 1973). In association with greenstone belts, Archean anorthosites commonly consist of dense concentrations of well-formed, equant megacrysts of calcic plagioclase in a finer-grained mafic matrix (Phinney et al., 1988). K-rich granitic rocks are not associated with Archean anorthosites. Anorthositic rocks that occur within some of the high-level Paleozoic ring complexes of Nigeria have some compositional and petrographic similarities to anorthosite within typical Proterozoic massifs (Brown et al., 1989). Even though these bodies are of small size and restricted occurrence, they eventually may prove of some help in understanding Proterozoic anorthosites. They will not be considered in this review. They do serve, however, as a reminder that the processes that created anorthosites of intermediate composition were not strictly limited to the mid-Proterozoic. DISTRIBUTION
Although the greatest concentrations of Proterozoic massif-type anorthosites occur in northeastern North America (Fig. 1) and northern Europe extending into the former Soviet Union, they are known on all continents but Australia. Major anorthosite complexes occur in Africa (Vermaak, 1981), India (Leelanandam and Reddy, 1988), and Madagascar (Boulanger, 1959). Herz (1969) first drew attention to the fact that, when the continents are returned to their Pangaeic pre-drift positions, most anorthosite massifs appear to occur in two broad belts: one in the northern hemisphere (Laurasia) that trends from the Ukraine through
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Fennoscandia and southern Greenland into eastern North America and one in the southern hemisphere (Gondwana) largely defined by bodies in India, Madagascar, and Africa. Subsequent workers (Bridgwater and Windley, 1973; Emslie, 1978; Anderson, 1983) noted that Proterozoic anorogenic granites closely followed the Laurasian belt and reinforced its extension to anorthositic rocks of the southwestern United States. Recent studies of plutonic rocks that occur beneath North American mid-continent Paleozoic sedimentary rocks have strengthened the evidence for a major magmatic belt on the order of 500 km wide (Bickford, 1988). The significance of these apparent belts remains controversial.
GEOLOGIC SETTING
Anorthosite massifs appear to have been emplaced into stable cratonic terranes of either Proterozoic or Archean age. Some were emplaced at the boundaries between distinct terranes (e.g. the Laramie complex, Wyoming - Geist et al., 1990; the Nain complex - at and near the boundary of the Nain and Churchill Provinces); some also appear to have been emplaced near unconformities between Archean gneisses and overlying Proterozoic supercrustal rocks (e.g. the Nain complex, Labrador - Speer, 1975). The common occurrence of anorthosite massifs within granulite-facies terrains of the Grenville Province neither provides an indication of the emplacement depth nor establishes a genetic link between granulite metamorphism and generation of anorthosites because many of these massifs were metamorphosed after emplacement (Morse, 1982a; Emslie, 1985). Thermobarometric studies of some unmetamorphosed massifs (Berg, 1979; Fuhrman et al., 1988) indicate they were emplaced at shallow to intermediate crustals levels (e.g. 6 to 16 km). Contact aureoles indicate that ambient temperatures of the country rock were as low as greenschist facies (Berg, 1977). Emplacement depths of up to 25 km or more have also been inferred from mineral assemblages in contact aureoles of other complexes (Anderson, 1980; Emslie, 1981). If massif-type anorthosites were emplaced in a single tectonic environment, then the bulk of available evidence suggests that they were emplaced anorogenically. Although the strong deformation shown by Grenville anorthosite massifs was often thought to result from syntectonic emplacement (Martignole and Schrijver, 1970), recent age and oxygen isotope studies of the anorthosites within the Grenville Province (Valley and O'Neil, 1982; McLelland and Chiarenzelli, 1990; Emslie and Hunt, 1990) suggest that the anorthosites were emplaced before the Grenville metamorphic event and that much of the deformation of these anorthosites was superimposed long after their emplacement. In areas where there is no convincing evidence of a later compressive tectonic event, the intense deformation displayed by the margins of some plutons may be attributed to deformation during emplacement of a crystal-rich mush (Duchesne et al., 1985). To date, no convincing evidence has been found to show that any massif was emplaced during a regional compressive tectonic event.
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In recent years, a rifting environment has been suggested for the emplacement of anorthosites (Bridgwater and Windley, 1973; Berg, 1977; Emslie, 1978). The Proterozoic Duluth Complex (Miller and Weiblen, 1990), commonly considered transitional between typical massif anorthosites and layered mafic intrusions, occurs within the Keweenawan Mid-continent Rift and lends support to the notion that at least some anorthosites were emplaced in rift environments. Nonetheless, no evidence of active rifting during emplacement has been established for typical massif anorthosites (Emslie, 1985).
ISOTOPIC AGES
Important progress in the precise dating of anorthosites has been made recently by obtaining U-Pb ages from zircon and baddelyite in anorthositic, gabbroic and
granitic rocks associated with the massifs (McLelland and Chiarenzelli, 1990; Emslie and Hunt, 1990). Even though the emplacement ages of many anorthosites are not well known or remain controversial because of uncertainties regarding later metamorphic events, reliably dated anorthosite massifs range in age from about 0.9 to 1.7 Ga. This age range dispels the idea that massif anorthosites represent a short-lived catastrophic event in the Proterozoic. Currently available ages of anorthosite massifs in eastern North America range between 1646 Ma (Mealy Mtns. - Emslie, 1976) and 1113 Ma (Marcy-McLelland and Chiarenzelli, 1990), and there is no apparent systematic regional variation in ages. The growing number of precise isotopic ages of separate plutons within single anorthosite complexes offers hope that it will soon be possible to have a better understanding of the range of intrusive ages within a single complex. Recent age data from the Nain complex (Krogh and Davis, 1973; DePaolo, 1985; Simmons et al., 1986; Simmons and Simmons, 1987; Carlson et al., 1992) suggest that magmatism was restricted to a relatively short time period of roughly 20 to 30 Ma. In contrast, the available age data for the Rogaland (Norway) anorthosite complex may permit either a long and complex emplacement history or reactivation of the complex in a time span of from 200 to 500 Ma (Duchesne et al., 1985; Weis, 1986).
INTERNAL CONSTITUTION OF MASSIF ANORTHOSITE COMPLEXES
The composite nature of the larger massif anorthosite complexes can be appreciated by examining the internal structure of two massifs: the Rogaland complex of southern Norway and the Nain complex of Labrador. The Rogaland complex has been thoroughly mapped and intensively studied over the last 50 years, and the main outlines and character of individual plutons were first well established by Michot (1960). A recent summary of this complex (Duchesne et al., 1985) provides both an historical view of the extensive earlier work and a clear statement of some current research problems. The Rogaland massif displays
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DlORlTE (JOTUNITE) ANORTHOSITE COUNTRY ROCKS 0
10
20
Fig. 2. Simplified geologic map of the Rogaland anorthosite complex (southern Nonvay). Modified after Duchesne and Demaiffe (1978). Anorthositic plutons: E - 0 = Egersund-Ogna; H-H = HalandHelleren; A-S = Ana-Sira; B-S = Bjerkreim-Sokndal lopolith (including a central area of mangerite and quartz mangerite); H = Hidra; G = Garsaknatt.
most of the distinctive features of Proterozoic anorthosite complexes. It consists of several separate anorthositic plutons of different character (Fig. 2), some of which may be composite in nature. One of them, the Bjerkreim-Sokndal lopolith is well-layered, shows clear geochemical evidence of magma replenishment and grades upward to noritic and mangeritic rocks. Three large granitoid plutons occur to the southeast of the anorthosites. Dioritic rocks (locally termed jotunites and monzonorites) occur as smaller irregular bodies and as dikes, and important concentrations of Fe-Ti ores occur a t several locations. The earlier anorthositic bodies show intense deformation especially along their margins and have been interpreted as diapiric intrusions of crystal-rich magmas. These bodies also carry high-Al orthopyroxene megacrysts (HAOM) that are thought to have crystallized at depth (Duchesne et al., 1985). Two apparently younger anorthositic bodies (Hidra and Garsaknatt) appear to be undeformed. The Nain complex of Labrador (Fig. 3) is much larger than the Rogaland complex and is composed of many more and a greater variety of plutons. The size of the complex, logistical difficulties, and the gaps in outcrop (between islands) have so far deterred efforts to complete a detailed map of individual plutons throughout the entire complex. Many large areas of Nain anorthositic rocks have not yet been adequately subdivided into separate plutons. Nonetheless, the excellent coastal exposures and the lack of any later tectonic or metamorphic overprint have made the Nain complex an ideal place in which to examine primary igneous contacts and internal features of anorthositic plutons. E.P. Wheeler undertook singlehandedly for many years a pioneering effort to map the Nain complex (Wheeler, 1942, 1960, 1968). His efforts to inspire work in the Nain area ultimately led S.A. Morse to establish the “Nain Anorthosite Project” (NAP) which operated from 1971 to 1981 and supported detailed studies by many workers. A complete bibliography of work supported by the NAP and earlier Field Reports of the NAP can be found in Morse (1983a). The anorthositic rocks of the Nain complex (Fig. 3) probably consist of more than 20 different plutons. Detailed studies of some areas (Wiebe, 1978, 1990b;
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Fig. 3. Simplified geologic map of the Nain anorthosite complex. Modified from Hill (1982) and Ryan (1990). A = South Aulatsivik Island (see Fig. 4b); P = Paul Island (see Fig. 4a); T = Tunungayualok Island (see Fig. 4c). The mafic Kiglapait layered intrusion (Morse, 1969) is north ofA.
Proterozoic anorthosite complexes
223
Ranson, 1981; Hill, 1982, 1988) suggest that this number may be a conservative estimate (see Fig. 4). Most plutons have a n average color index (CI) between 5 and 15. Some leucotroctolite bodies (e.g. the Hettasch intrusion - Berg, 1980) are somewhat more mafic on average and can be considered intermediate between typical anorthosite and mafic layered intrusions. Anorthositic rocks in the southeastern part of the Nain complex tend to be more evolved with lower An and Mg# and locally have cumulus clinopyroxene, ilmenite, and apatite in addition to Fe-rich olivine (Hill, 1988). Individual anorthositic plutons with plagioclase lamination and subtle layering include both diapirs and bodies that appear to have solidified by deposition of cumulates on magma chamber floors. Some batholith-sized areas of massive, undeformed, hypidiomorphic leuconorite to anorthosite may represent stagnant accumulations of plagioclase in the interior or beneath the roofs of magma chambers (Wiebe, 1990b). On the basis of cross-cutting relations between plutons in the Nain complex, Morse (1983b) recognized that plutons with steep foliations (possible diapirs) tend to be emplaced early and layered troctolitic anorthosites tend to be late. Nonetheless, some massive anorthosite plutons are younger than some layered troctolites (Wiebe, 1988). It appears that in both the Rogaland and Nain complexes, the earlier anorthositic plutons commonly have deformed margins or steep folations, while later bodies typically show less deformation and may be either massive or layered. In Fig. 3, all mafic intrusions that essentially lack an anorthositic component are shown in black. Mappable bodies of layered and massive diorite are widespread, and diorite dikes are locally abundant within anorthosite. Many diorites are thin, subhorizontal sheets commonly located structurally above anorthositic rocks and below granitic or hybrid mixtures of dioritic and granitic rocks. The Nain complex also includes three mafic (troctolitic) layered intrusions. The largest and most thoroughly studied is the Kiglapait intrusion (Morse, 1969, 1981) (Fig. 4b). The other two, the Barth Island structure (De Waard, 1976) and the Newark Island intrusion (Wiebe, 1988) are hybrid intrusions involving multiple replenishments of both troctolitic and granitic magmas. Large bodies of granitic rock are associated with the Nain anorthosites (Fig. 3). Very few have been thoroughly studied. Exceptions include peralkaline and subalkaline granites located in the southeastern Nain complex (Collerson, 1982; Hill, 1982). Most granitic rocks appear to have very roughly equal amounts of quartz, alkali-feldspar and plagioclase. Some syenitic and monzonitic cumulate rocks are associated with diorites or occur along the margins of massive granitic bodies. Granites were emplaced a t different times, overlapping with the emplacement of the anorthosite plutons, and there is clear evidence of anorthosite locally cutting granite (Wiebe, 1988). The granitic rocks commonly show evidence of commingling with dioritic magma (Wiebe, 1980a). Many granitic plutons in the Nain complex appear be thin, subhorizontal sheets (Wheeler, 1968; Hill, 1982), so that the volume percent of granite may be much less than that suggested by its areal extent.
R.A. W e b e
224
Both the Nain and Rogaland as well as many other complexes are cut by basaltic dikes that may be only slightly younger than the anorthosites. The compositions of the Nain dikes closely resemble the compositions of some Fe-rich troctolitic
“]LOWER
LEUCONORITE
Fig. 4. Detailed geologic maps of selected areas in the Nain complex. For locations within the Nain complex, see Fig. 3. a. Paul Island (Wiebe, 1990b). b. South Aulatsivik Island and adjacent areas. Contacts are taken mainly from Field Reports of the Nain Anorthosite Project (see Morse, 1983a; Wiebe and Wild, 1983), or based on more recent field work. Symbols for anorthositic plutons:A = undifferentiated anorthosite; S = Slambang leuconorite; P = Port Manvers Run anorthosite (LZ = olivine-bearing lower zone; UZ = leuconoritic upper zone); H = Hettasch layered intrusion (leucotroctolitic). Mafic plutons: KIG = Kiglapait layered intrusion (troctolitic); TIG = Tigalak composite layered intrusion (dioritic); NILI = Newark Island composite layered intrusion (troctolitic).
Proterozoic anorthosite complexes
225
Fig. 4. (continued) c. mnungayualok Island area (Wiebe, 1978).
liquids that were emplaced contemporaneously with the anorthosites (Carlson et a]., 1992). It is possible that these dikes represent the waning stages of magmatism that generated the anorthosites.
CHARACTERISTICS OF ANORTHOSITIC PLUTONS AND ASSOCIATED DIKES
Types of plutons The internal structure, composition, and contact relations of anorthositic plutons can be described by considering three main end-member types: (1) diupirs, (2) Zyered intrusions developed by bottom accumulation, and (3) massive bodies apparently devoid of internal or marginal deformation. Even though these three types provide a useful context for considering the principle field relations, internal structures and petrographic character of anorthosite, it is possible that they may in some places represent different phases or levels of the same pluton. Although examples of individual plutons will be drawn mainly from the Rogaland and Nain anorthosite complexes, these three types appear to display the range of pluton
226
R.A. Wiebe
characteristics found in most unmetamorphosed anorthosite complexes. It should be emphasized, however, that the three types described here are end-members: as more detailed field work is done within anorthosite complexes, plutons of intermediate character may prove to be common. For example, Scoates (1990) has recently suggested that part of the Laramie anorthosite developed by diapiric ascent of a partially crystallized layered anorthosite - with the base of the layered body (now the core of a diapir) showing solid state deformation, and the top of the layered section (now the outer portions of the diapir) showing slumping and compaction due to the presence of liquid during diapiric uprise. Recent field work in the Nain complex also suggests that some massive bodies may represent the roof zones of layered bodies. Pefrographic characteristics of anorthosite plutons Anorthositic rocks that lack cumulus mafic phases are dominated by tabular to equant plagioclase with sizes typically in the range of 1 to 10 cm. Some rocks have plagioclase of nearly uniform size while others show a wide and characteristically continuous range in sizes. Weak alignment of tabular plagioclase (lamination) is quite common but often difficult to detect in rocks with low
Fig. 5. Etched slab of leuconorite from the lhnungayualok Island leuconorite. Width of photo is about 8 cm. Dark areas are compositionally zoned with orthopyroxene interiors and rims including clinopyroxene, hornblende and oxides. White areas in feldspar are weakly granulated.
Proterozoic anorthosite complaes
227
color index. Mafic minerals, usually either olivine or orthopyroxene, fill interstices between plagioclase crystals creating subophitic to poikilitic textures that permit clear recognition of the habit and form of the plagioclase (Fig. 5). Very commonly, mafic minerals occur as evenly spaced poikilitic areas that give the anorthosite a spotted appearance in the field (Fig. 6). These “spots” commonly range from 2 to 20 cm in size and may be either equant or lensoid. As the abundance of the spots increases they gradually coalesce and lead to a rock best described as a leuconorite or leucotroctolite with irregular patches of anorthosite. The maficfree areas of some spotted leuconorites contain other minerals such as quartz and alkali-feldspar as interstitial phases (Wiebe, 1978). These textures formed by early crystallization (and probably accumulation) of plagioclase followed by widely scattered nucleation and growth of a mafic phase. The size and spacing of spots decreases rapidly toward the intrusive margin of one pluton in the Nain complex (Wiebe, 1980b), suggesting that the rate of cooling is a controlling factor in the development of this spotted texture. Nearly pure anorthosites in diapiric intrusions commonly have bent plagioclase with strongly sutured contacts between adjacent grains. Calcic myrmekite (Dymek and Schiffries, 1987) may also be abundant in these rocks. Interstitial phases (augite, opaques, etc.) typically occur as thin curved lenses along the
Fig. 6. Leuconorite dikes (with poikilitic spots comparable to those illustrated in Fig. 5) cut an older anorthosite on Tunugayualok Island (Wiebe, 1978).
RA.Wiebe
228 TABLE 1 Average modes of three types of anorthositic plutons
Quartz K-feldspar Plagioclase Olivine OPX Inv. pigeonite CPX Hornblende Biotite Opaques Apatite Color Index Number of samples Size of grid Total points counted a
Diapir
Layered
Massive
Tun-An a
Harp Lake
Tun-Ln
92.4
-
5.7 1.2 0.6
0.1
1.6 14 2mm 2433
-
16.6
2.0 2.1 83.6 2.2 6.0 1.9 0.9 0.2 1.04 0.1 12.3
22 0.4-0.5 mm 33,493
14 2 mm 2424
83.4 10.8 2.5
-
2.1
-
0.3 0.9
-
Thn-An: nnungayualok Island older anorthosite; Fig. 4c. Average of specimens, 1-22, from table 8 of Emslie (1985). Tun-Ln: Tunungayualok Island younger leuconorite; Fig. 4c.
sutured boundaries. These textures may have formed during slow compaction of plagioclase-rich mushes during diapiric emplacement and probably indicate that filter-pressing was the dominant mechanism for expelling interstitial liquid within these massive and diapiric bodies. These same rocks have commonly undergone varying degrees of recrystallization that has produced finer-grained areas with mosaic textures. Modal analyses of rocks from examples of the three types of plutons are given in Table 1. Rocks from diapirs commonly lack olivine and have the lowest average CI. Layered bodies are typically leucotroctolitic. Because of the sporadic occurrence of cumulus olivine, modes of individual rocks vary widely, and the average CI may be higher than in the other types of bodies. Massive plutons (type 3) are typically leuconoritic and, in the Nain complex, commonly have small amounts of interstitial quartz and alkali-feldspar, the presence of which may indicate either fractional crystallization to oversaturated liquids or contamination of the parental melt by crustally derived granitic liquids. The range and variability of isotopic compositions in these massive leuconorites (see below) tend to support the latter process.
Diapirs Some anorthositic bodies have variably deformed marginal zones that display a strong foliation conformable with their steep intrusive contacts. Although these
Proterozoic anorthosite complexes
229
Fig. 7. Strongly deformed leuconorite in the marginal zone of the Egersund-Ogna body (Rogaland complex - see Fig. 2). Large orthopyroxenes are variably stretched into lenses.
structural features may also develop where anorthosite has been subjected to later deformation, there are plutons within the Rogaland and Nain complexes where such features are best explained by diapiric emplacement of a plagioclase-rich crystal mush lubricated by interstitial noritic to leuconoritic liquid (Longhi and Ashwal, 1985). The Egersund-Ogna intrusion in the Rogaland complex (Fig. 2) is a classic example of a dome-shaped body with a n intensely deformed margin 1 to 3 km wide (Duchesne e t al., 1985) (Fig. 7). The Lower Leuconorite (LLN) of Paul Island (Fig. 4a) is a diapiric body with a less strongly deformed marginal zone (Wiebe, 1990b). These bodies have many features in common. Both are essentially leuconoritic in composition and contain within their interiors large
230
R A . Wiebe
blocks of gneissic anorthosite and leuconorite that resemble their gneissic margins in texture and mineral chemistry. High-Al orthopyroxene megacrysts (HAOM) with exsolved plagioclase lamellae occur within these blocks associated with megaplagioclase. The reintegrated compositions of the HAOM are much richer in A1203and Cr than orthopyroxenes from the surrounding matrix anorthosite. Within these plutons, the HAOM are thought to have crystallized at depth and been carried up with the rising crystal mush (Duchesne et al., 1985; Wiebe, 1990b). The marginal zone of the Egersund-Ogna body locally shows intense granulation, while the marginal zone of the LLN on Paul Island only shows moderate granulation and stretching of primary ophitic textures. Leuconorite veins cut the marginal foliation of both bodies a t large angles and probably represent residual liquid that was filter-pressed near the final stages of diapiric uprise. Although much of the LLN shows a moderate to strong lamination of plagioclase (1 to 5 cm in length), compositional layering is essentially absent, and no structures suggestive of crystal accumulation on a chamber floor were found. The lamination probably developed during movement of the crystal mush during emplacement.
Layered an orthositic intrusions Anorthositic plutons with gently dipping layering, good to fair lamination of plagioclase, and sporadic occurrences of cumulus olivine or, less commonly, orthopyroxene occur in many complexes. Layers generally are defined by varying proportions and distributions (e.g., different size poikilitic spots) of interstitial mafic minerals. Layers are typically discontinuous and vary in thickness between centimeters and meters. Many subunits (separate plutons?) within the Harp Lake complex (Emslie, 1980) show these features as well as a variety of structures that lend strong support to their interpretation as bottom cumulates. The marginal phases of these plutons are typically finer-grained and more mafic (gabbroic or troctolitic) in composition (Emslie, 1980), and the floors of these layered bodies may be gently dipping or funnel-like as suggested for the Michikamau intrusion (Emslie, 1970). No roofs against overlying country rock have been identified. Two layered bodies in the Nain complex (the LTR unit on Paul Island, Fig. 4a; the Port Manvers Run intrusion, Fig. 4b) clearly grade upward into overlying massive leuconorite, and layered portions of the Harp Lake complex may show similar gradations (Emslie, 1980). In the Nain complex several different layered plutons display evidence of bottom accumulation (e.g. Berg and Briegel, 1983; Wiebe, 1990b). Typically, these bodies do not appear to fractionate upward to mafic cotectic cumulates as might be expected for closed system fractional crystallization. However, minor intrusive bodies of Fe-rich dioritic rocks that might represent residual liquids commonly occur along the margins and structurally above anorthosite in many complexes. An unusual body in the Rogaland complex, the Bjerkreim-Sokndal lopolith,
Proterozoic anorthosite complexes
23 1
records several cycles of fractionation from anorthositic cumulates to cotectic norites (Duchesne, et al., 1985). These cycles apparently resulted from periodic replenishment by feldspathic magma. The LTR unit on Paul Island (Wiebe, 1990b) has features that are characteristic of the most common layered anorthositic bodies (Fig. 4a). This unit consists dominantly of olivine-bearing anorthosite and leucotroctolite and has an average CI of 10-15. Plagioclase is the only cumulus phase, except in the upper half of the unit where many lenses and layers of unimodal and modally graded olivine-plagioclase cumulates up to 10 m thick and up to a few kilometers long are intercalated with massive to laminated anorthosite. Cumulus olivine always occurs in less than cotectic proportions. The occurrence of more primitive plagioclase compositions in layers with cumulus olivine indicates that the layers represent replenishment episodes (Wiebe, 1990b). Replenishments also occur in the Hettasch intrusion of the Nain complex (Berg, 1980). The impersistence (upward disappearance) of cumulus olivine is characteristic of layered anorthositic plutons (Emslie, 1980). Density-graded layers with cumulus olivine provide convincing evidence for deposition of floor cumulates. There are also many features that suggest soft-sediment deformation, slumping and redeposition of incompletely solidified plagioclase cumulate material (Fig. S), and some layers with abundant unsorted cognate anorthosite blocks are intercalated between coherent layers of well-laminated anorthosite. In the Nain complex, xenocrystic HAOM appear to be absent from olivine-bearing layered bodies; they have been reported, however, in similar layered rocks within the Harp lake complex (Emslie, 1980). The Port Manvers Run intrusion (average CI = 15) has an exposed section about 5 km thick and has a basal zone of olivine-bearing anorthosite that grades upward to leuconorite and at the highest exposed levels to massive leucogabbronorite with prominant interstitial inverted pigeonite, augite, oxides and quartz (Snyder, 1984) (Fig. 4b). Plagioclase composition varies irregularly upward (from about an average of h 6 0 a t the base to A ~ ~ + . at s othe top). The surprising aspect of this intrusion is the fact that the mafic minerals show evidence of effective Fe-Mg fractionation upward - even though, on the basis of textural criteria (Irvine, 1982), cumulus mafic minerals are absent except for a few meters of section with cumulus olivine. This fractionation in Fe-Mg might be explained by a heteradcumulate process in which nucleation and partial growth of a mafic phase occurred interstitially within the cumulate pile of plagioclase crystals, while chemical exchange between the magma reservoir and intercumulus liquid was still possible. It is also possible, however, that this upward variation in Fe/Mg (and Si02) could be a record of magma stratification - perhaps generated in part by the collection and mixing-in of Si02-rich, crustally derived melts beneath the roof of the magma chamber.
232
R A . Wiebe
Fig. 8. Leucotroctolitic rocks from the layered LTR unit on Paul Island (Fig. 4a) with prominent slump structures.
Massive plutons Massive plutons that are characterized by a massive interior (the general absence of any measurable planar or linear fabric) and the absence of deformed marginal rocks. As noted above, it is possible that some of these plutons represent stagnant accumulations of plagioclase near the roofs of chambers that produced layered leucotroctolitic cumulates on their floors. The margins of these plutons are typically more fine-grained within distances of meters to 10’s of meters from the country rock. Marginal rocks may range from leuconorite (CI = 10-15) to a wide variety of massive and layered mafic rocks, and the character of these rocks may vary widely along the contact of a single pluton. Large composite blocks of earlier-formed margin attached to country rock occur in the interior of some bodies. In the Nain complex, examples of these bodies include the Rnungayualok Island leuconorite (Fig. 4c) and the ULN unit on Paul Island (Fig. 4a). In the Rogaland complex, the Hidra body (Fig. 2) consists of a massive interior of leuconorite (with interstitial quartz and alkali-feldspar) that grades outward to a fine-grained dioritic margin (Demaiffe and Hertogen, 1981). Anorthosite and mafic dikes commonly extend as apophyses of the main anorthositic plutons into the adjacent country rock. Partial melting is common
Proterozoic anorthosite complexes
233
in quartzo-feldspathic gneisses that lie along the contact and locally leads to the generation of incoherent migmatites. Granitic melt generated in the adjacent country rock has mixed into and contaminated both anorthositic dikes and marginal phases of these plutons (Wiebe, 1978). The homogeneous interiors of massive anorthosite bodies typically consist of leuconorite with a CI of about 15. The leuconorite is characterized by a seriate porphyritic texture with maximum plagioclase size up to 10-15 cm. Plagioclase with delicate and complex compositional zoning is widespread in some plutons. Interstitial oxides as well as quartz and alkali-feldspar are common in many massive plutons; xenocrystic HAOM has not been found in this type of pluton. Large unsorted anorthosite blocks (up to 100’s of meters in length) are common and locally abundant over large areas. The ULN unit of Paul Island (Fig. 4a) consists of massive, coarse-grained leuconoritic rocks with about 15% interstitial orthopyroxene, augite, ilmenite, magnetite, apatite, and scarce alkali-feIdspar and quartz. The texture is dominated by 1 5 2 5 % complexly zoned, seriate plagioclase phenocrysts up to 20 cm in length. These phenocrysts are iridescent, and variably colored zones commonly exhibit up to three major cycles of normal zoning separated by abrupt reversals of 10-15% An that may truncate inner zones and appear to record episodes of partial resorption of the plagioclase (Fig. 9). A few percent of similar crystals occur sporadically within the upper portions of the underlying olivine-bearing LTR unit (Fig. 4a); their presence strongly suggests that both the LTR and the ULN units formed within the same magma chamber. The zoned iridescent phenocrysts were probably suspended within the chamber while cumulates within the LTR were being deposited; the resorption episodes recorded in their compositional zoning
Fig. 9. Oscillatory zoned iridescent plagioclase from the massive ULN unit on Paul Island (Fig. 4a).
234
R A . Wiebe
may correspond to the replenishment episodes that are apparent within the LTR cumulates (Wiebe, 1990b). The eastern contact of the ULN body dips moderately outward. A finer-grained leuconorite occurs along much of it, and thin irregular anorthositic veins locally produce migmatic areas within basic granulite country rock. Within the ULN large angular inclusions occur near the highest exposures up to several km in from the eastern contact. The inclusions consist dominantly of fine-grained iridescent anorthosite and leuconorite (CI = 5-20); some are composite blocks consisting of similar anorthosite and basic granulite in sharp contact with each other. Because the anorthosite in the blocks resembles the marginal fine-grained phase of the intrusion, and the granulites are similar to the adjacent country rock, the blocks are likely to have spalled from a relatively flat roof to the intrusion. These relations suggest that the ULN unit formed as a stagnant accumulation of plagioclase crystals below the roof of a magma chamber. Dikes Anorthositic dikes varying in thickness from a few centimeters to a meter have been reported in several anorthosite complexes (Emslie, 1975a; Wiebe, 1979a, 1990b; Dymek, 1980; Leelanandam and Jyothender Reddy, 1990). Some dikes consist of nearly pure anorthosite (CI = 5 ) with unzoned plagioclase and are interpreted as having primary igneous textures (Dymek, 1980). Most of the anorthositic dikes within the Nain complex have a CI of about 15 and are more precisely termed leuconorites or leucogabbros. A few fine-grained dikes have strongly zoned plagioclase (An56-37)with calcic patchy cores and more sodic normally zoned rims (Wiebe, 1979a). They lack any signs of granulation that might be expected if they were injected as a largely crystalline mush (Van der Molen and Paterson, 1979). Some of these dikes appear to have been injected largely as liquids, with perhaps as little as 20-30% suspended plagioclase crystals (Wiebe, 1990b). The high liquidus temperatures of one of these dikes (1365" to 1420°C from 1bar to 20 kbar) strongly suggest that it could not have been injected entirely as a liquid (Fram and Longhi, 1991a). Nonetheless, the liquid carrying plagioclase crystals is likely to have been highly oversaturated in plagioclase components and could be appropriately termed hyperfeldspathic (Wiebe, 1990b). The more than 20 leuconoritic dikes that have been studied within the Nain complex are widespread, but are most commonly found either in country rock along the margins of the massive undeformed plutons (type 3) or within diapiric plutons (type 1) (Wiebe, 1990b). Dikes into country rock occur mostly within 50 m of the contact of a massive pluton and within refractory rocks like basic granulite. Where a massive anorthosite pluton is in contact with quartzofeldspathic gneisses, locally derived granitic melts from the country rock tend to fill fractures before anorthositic magmas can penetrate far. Along one such contact, anorthositic dikes could be traced only a few meters from the contact to a point where mixing with granitic melt occurred (Wiebe, 1978). Fine-grained leuconoritic veins (CI = 15)
Proterozoic anorthosite complexes
235
within the diapiric LLN body on Paul Island (Fig. 4a) are typically just a few centimeters thick and occur at high angles to the foliation within about 100 m of the contact. These appear to represent liquid that was filter-pressed into cracks that formed during the final consolidation of the diapir carapace. The geologic settings of these dikes and their wide variation in minor-, traceelement and isotopic composition (Simmons et al., 1985; Wiebe, 1990b) suggest that hyperfeldspathic liquids may have been generated only in response to specific conditions of diapiric emplacement and magma chamber dynamics. Rather than representing parental liquids to the anorthosite plutons, their compositions may strongly reflect local processes (e.g., remelting of suspended plagioclase and contamination by silicic crustal liquids) that operated during the emplacement of individual plutons.
MINERALOGY
PIagioclase Perhaps the most distinctive aspects of plagioclase in massif-type anorthosites are its large size and the common occurrence of iridescence ( A n 4 8 - 5 6 ) due to “Boggild” exsolution (Ribbe, 1975). The sharply defined colors (red, orange, yellow, green and blue) may be caused by an even spacing of lamellae that is dependent upon An-content, with red more calcic and blue more sodic (Ribbe, 1975). In many plutons plagioclase appears nearly black in the field due to the presence of thin plates of exsolved oxides (Anderson, 1966). Excluding minor thin calcic rims (Morse and Nolan, 1984), plagioclase compositions generally fall in the range k h 6 5 - 3 5 . The more calcic plagioclase typically occurs with olivine. Sodic plagioclase may be strongly antiperthitic (e.g. Herz, 1968). Plagioclase from anorthosite massifs is significantly higher in Sr and lower in Ca compared with plagioclase from other basic intrusions (Emslie, 1985). In most plutons, individual plagioclase crystals show little or no primary igneous zoning; however, in many massive (type 3) bodies, normal and oscillatory zoning (up to 10-15% An) is common (Wiebe, 1990b). Exceptionally, compositional zoning shown by concentric non-iridescent and iridescent rings of different colors can be seen in the field (Fig. 9). In massive plutons, the range of zoning in a single crystal is commonly much greater than the range in average plagioclase composition of hand-samples throughout the pluton (Wiebe, 1978). Exturally, these rocks resemble orthocumulates, but the uniformity of average plagioclase compositions over large areas provides little evidence of large-scale fractional crystallization, a t least at the level of exposure. The zoning and overall homogeneity of these bodies are consistent with crystallization of a stagnant mush. In the diapiric and layered plutons, plagioclase crystals typically show little zoning except for thin calcic rims. In layered plutons, areal variations in plagioclase composition probably records fractional crystallization and magma chamber
R.A. Wiebe
236
900T
Sr 70,
-T-LN
5095
%A"
Fig. 10. Sr versus An in plagioclase separates from different anorthosite plutons. Mineral chemistry determined by X-ray fluorescence. L L N , ULN, and LTR: see Fig. 4a. T-AN, T-LN and K refer respectively to nnungayualok Island anorthosite and leuconorites and Kheovik leuconorites in Fig. 4c.
replenishment as suggested for the LTR unit on Paul Island (Wiebe, 1990b). In diapirs, plagioclase from marginal and interior portions may have different major and trace-element compositions - e.g., plagioclase in the interior of the Egersund-Ogna massif is more sodic and higher in Sr than plagioclase from the margins (Duchesne et al., 1985). An inverse relationship between Sr and An-content of plagioclase is a common characteristic of plagioclase from anorthosite complexes (Emslie, 1985). Plagioclases from different plutons typically have different Sr/Ca ratios that probably reflect variations in source and/or fractionation of different batches of magma (Fig. 10). Because the partition coefficient of Sr in plagioclase is greater than one, the increase in Sr implies that the crystallization of a Ca-rich, Sr-poor phase like clinopyroxene is necessary to explain compositional differences between plutons (Emslie, 1985). Olivine and pyroxenes Olivine (F070-50) or orthopyroxene (En70-40) is generally the second most abundant mineral in anorthosite. Zoning in both is typically absent even in large poikilitic areas - probably due to Mg-Fe exchange during slow cooling. Orthopyroxene typically has thin exsolution lamellae of augite. Coarsely exsolved inverted pigeonite (En50-40) is prominent in many plutons. Original zoning of poikilitic pyroxene areas can be seen where cores are orthopyroxene and rims of variable thickness consist of inverted pigeonite and augite. Augite with complex exsolution is present in many plutons and is the dominant pyroxene in some complexes (e.g. Laramie complex - Scoates, 1990). Figure 11 shows the range of pyroxene compositions from some plutons in the Nain complex. Plagioclase coexisting with olivine and orthopyroxene of given compositions tends to be more sodic in anorthosite plutons than in mafic layered intrusions like the Bushveld and Stillwater (fig. 2 in Emslie, 1985). Within individual hand-
Proterozoic anorthosite complexes
237
Fig. 11. Compositional range of pyroxenes in some Nain anorthosites (an) from Tunugayualok and Kheovik islands (Fig. 4c) and associated diorites from the Goodnews complex (Wiebe, 1979b). Also shown is the typical compositional range of high-A1 orthpytoxene megacrysts (HAOM) that occur in the Nain and other anorthosite complexes.
specimens, covariation of An in plagioclase and En in orthopyroxene (or Fo in olivine), which is generally well displayed in layered mafic intrusions, is not often clearly defined - probably because plagioclase is the only cumulus phase.
High-A1 orthopyroxene megacrysts High-Al orthopyroxene megacrysts (HAOM) with exsolved lamellae of plagioclase occur either as isolated xenocrysts or as subophitic to ophitic intergrowths with plagioclase within many anorthosite complexes (Emslie, 1975b; Morse, 1975; Dymek and Gromet, 1984; Duchesne, 1984; Wiebe, 1986). Individual crystals as large as 10 to 20 cm are common. Compared with typical interstitial orthopyroxene in anorthosites, these crystals tend to be Mg-rich (En75--68) and to have much higher Al, Cr and Ni (Emslie, 1975b) (Fig. 12). Unexsolved pyroxenes that are compositionally similar to the HAOM occur in nodules with euhedral plagioclase and as large xenocrysts within late basaltic dikes that cut the Nain complex (Wiebe, 1986) (Fig. 12). Emslie (1975b) suggested that these orthopyroxenes crystallized at depth and were carried up within rising plagioclase-rich magmas. Emslie (1975b) and Maquil (1978) conducted high-pressure experiments that tended to support a high pressure origin for these megacrysts. Because the HAOM commonly occur in ophitic intergrowths with plagioclase, Morse (1975) and Dymek and Gromet (1984) suggested that they grew rapidly in place at the final level of emplacement and are chemically distinct for kinetic reasons from other orthopyroxenes within the same outcrop. Although the ophitic texture is evidence that the HAOM crystallized contemporaneously with adjacent plagioclase, the likelihood that many plutons rose to shallow levels as crystal-rich and inclusion-rich diapirs suggests that both the intergrown plagioclase and the HAOM may have crystallized a t depth. Within the Nain complex, which was emplaced at shallow depths in the crust (Berg, 1979), HAOM are unknown in plutons that appear to have cqstallized largely at the final level of emplacement (i.e., the layered olivine-bearing bodies and the massive leuconoritic bodies with zoned plagioclase). HAOM occur
R.A. Wiebe
238
5 3-
U
0 2-
* .
1-
0 7
70
60
50
40
100 Mg/(Mg+ FeT)
Fig. 12. A comparison of A1203 versus En in interstitial orthopyroxenes from the lhnugayualok and Kheovik bodies (small solid circles) with bulk analyses of high-Al orthopyroxene megacrysts (HAOM) containing exsolved plagioclase lamellae (open squares and diamonds) from different complexes (Emdie, 1978, 1980; Dymek and Gromet, 1984). The compositional range of unexsolved HAOM from nodules in Nain dikes (Wiebe, 1986) are shown by the horizontally ruled areas.
only in undoubted diapirs or plutons of probable diapiric origin (i.e., those with internal deformation and steep foliation). In the Rogaland complex these megacrysts occur prominently in plutons interpreted as diapirs but are apparently absent from layered and massive bodies like the Berkreim-Sokndal lopolith and the Hidra body. These field relations suggest that the HAOM crystallized at relatively high pressure and were brought to upper crustal levels within crystal-rich magmas.
Other minerals Hornblende in various shades of brown to green occurs commonly as late magmatic rims on pyroxenes. In the Nain complex; it is particularly abundant in leuconorites that contain interstitial quartz and alkali-feldspar. Where anorthosites occur in the vicinity of hydrous granitic rocks, subsolidus replacement of pyroxenes by actinolitic hornblende is also widespread. Biotite is a common accessory mineral in many anorthositic rocks and is nearly ubiquitous in very small amounts within olivine-bearing anorthosites. The volatile content of these minerals is not well known, but they are generally assumed to contain some water and to indicate that the last residual liquids were enriched in H 2 0 (Dymek and Schiffries, 1987).
Proterozoic anorthosite complexes
239
Fe-Ti oxide minerals (dominantly ilmenite) are widespread in anorthositic rocks and occur up to several percent in many leuconoritic rocks and at greater abundances in associated Fe-rich mafic rocks and ore deposits. They typically display complex oxy-exsolution that has been described by Buddington and Lindsley (1964) and Haggerty (1976). Exsolution and re-equilibration to low temperatures hinders their use as a geothermometer. Quartz and alkali-feldspar occur as scarce interstitial phases within many leuconorites. Interstitial micrographic intergrowths of quartz and alkali-feldspar occur in the massive Hidra leuconorite intrusion (Fig. 2) (Demaiffe and Hertogen, 1981). Other common accessory minerals include apatite and sulfides.
GEOCHEMISTRY
Chemical compositions of anorthosites (Table 2) largely reflect the dominance of plagioclase, that is, they are characteristically high in Na20, CaO and A 1 2 0 3 and low in major and trace elements that tend to be excluded from plagioclase. The sampling of many anorthosites for chemical analysis is difficult, not so much for the large size of the plagioclase (it tends to be relatively unzoned and similar in composition throughout a large area), but for the very uneven distribution of interstitial phases on the scale of centimeters to meters. The wide spacing of poikilitic mafic areas and their tendency to be mineralogically zoned indicate that chemical components within the interstitial liquid diffused over long distances and crystallized over a considerable temperature range. Because individual samples collected in the field rarely can be of adequate size to be representative of minor components, trends of groups of individual samples on variation diagrams are likely to be an artifact of small sample size. Such an effect was noted in the anorthosite of the Stillwater layered intrusion (Salpas et al., 1983). Oxide-oxide plots of individual analyses from most anorthositic plutons show considerable scatter and generally do not show clear trends on plots of normative An versus Mg#. Representative chemical analyses and norms of hand-specimens from three Nain plutons located on Paul Island (Fig. 4a) are given in Table 2. %ace elements that tend to be excluded from plagioclase are generally of very low abundance. Those compatible with pyroxenes (e.g., Cr, Co, Sc) commonly show expected positive correlations with FeO and MgO (Haskin and Salpas, 1986). The abundances of trace elements largely excluded from both plagioclase and pyroxene (e.g., P, Y, REE, Zr, Rb) can be expected to provide a measure of the amount of trapped liquid in the anorthosite sample. Since textures suggest that plagioclase is the only cumulus phase, one might expect that positive correlations would exist between these incompatible trace elements and those compatible with pyroxene. Instead, it has commonly been found that samples with higher Cr, Co and Sc tend to have relatively low abundances of light R E E (Haskin and Seifert, 1981; Haskin et al., 1981). This behavior can most simply be explained by the nucleation and growth of poikilitic pyroxene (up to several centimeters
R.A. Wiebe
240
TABLE 2 Representative chemical analyses of three different types of anorthositic plutons from Paul Island, Nain complex, Labrador (Fig. 4a). Oxides in weight %; trace elements in ppm. LLN = lower leuconorite; LTR = leucotroctolite; ULN = upper leuconorite. %An = 100An/(An+Ab+Or); Mg# = 100 Mg/(Mg FeT) [cation]
+
~
~~
~
LLN - diapir Soecirnen: SiOz Ti02 A1203 Fez 0 3 FeO MnO MgO CaO Naz 0 K2 0 pz 0 s LO I Total Sr Ba Y
Zr V Ni Cr sc
CIP w noms AP I1 Mt Or Ab An
B-53
8-7B
53.85 53.88 0.07 0.13 21.54 25.68 0.38 0.86 0.66 2.13 0.01 0.05 0.62 2.09 10.63 9.79 4.75 4.73 0.35 0.40 0.01 0.08 0.82 0.73 99.69 100.55 853 274 3 12 8 7 7 2
760
288 4 15 24 15 54 6 0.18 0.25 1.29 2.43 41.22 49.08 0.34 3.91 1.11
Q
0.02 0.13 0.56 2.09 40.59 53.20 1.33 0.09 0.85
%An Mg # An+Ab+ Or
55.5 52.5 95.88
52.9 56.2 92.73
Di HY 01
Ne C
-
-
-
LTR - layered B-7
51.88 0.66
23.55 1.90 3.56 0.06 2.45 10.34 3.90 0.33 0.02 0.59 99.29 804 256 5 18 128 26 70 10
B-123 52.02 0.11 28.43 0.67 0.70 0.02 1.03 12.38 3.77 0.26 0.03 0.27 99.69 617 174 4 12 17 14 8 3
B-148 50.52 0.08 27.03 0.95 1.62 0.03 3.00 11.59 3.66 0.20 0.03 1.19 99.90 597 126 4 13 10 84 7 2
3.37
0.07 0.21 0.98 1.55 32.20 60.45 1.18 1.62 1.15
0.07 0.15 1.39 1.19 31.08 56.95 0.50 2.24 5.15 -
56.7 46.1 81.39
64.2 58.5 94.20
63.8 68.3 89.22
0.04 1.27 4.45 1.97 33.27 46.15 4.85 2.76
-
-
ULN - massive B-126
B-172
B-102
8-104
48.87 53.69 52.67 52.61 0.14 0.15 0.16 0.09 27.12 24.91 25.09 24.59 1.06 0.88 1.49 0.60 4.38 2.50 1.86 1.17 0.03 0.05 0.05 0.06 1.97 3.09 3.44 5.80 11.42 9.99 10.49 10.81 4.23 3.88 4.15 3.15 0.30 0.37 0.29 0.24 0.03 0.03 0.02 0.02 0.54 0.52 0.69 - - _ _ -0.46 100.54 100.41 99.90 99.30 533 146 4 14 15 118 5 3
582 181 5 14 30 21 82 6
0.04 0.17 0.86 1.40 26.25 51.46 1.06 1.26 16.11 -
0.07 0.27 1.57 1.74 35.27 55.53 1.73 3.17 0.14 -
65.0 67.7 79.11
59.7 62.4 93.04
-
656 221 4 16 27 29 72 6
626 190 4 15 33 37 104 7
0.07 0.30 1.34 2.29 37.52 50.21 1.33 6.44 0.39
0.05 0.32 2.25 1.78 34.13 52.17 1.57 6.24 0.93
55.8 62.6 90.02
59.2 65.7 88.08
in diameter) within a framework of cumulus plagioclase crystals; incompatible trace elements will necessarily diffuse away from the growing poikilitic pyroxenes and eventually be concentrated in interstitial liquid away from pyroxene. Samples
241
Proterozoic anorthosite complexes
I
,
La Ce
Nd
S b Eu
Th
I
Y h Lu
Fig. 13. Chondrite-normalized REEs in orthopyroxene-bearing anorthositic rocks and associated chilled dioritic rocks from the Nain complex. The size of the E u anomaly in the anorthosites does not correlate well with the modal abundance of plagioclase. See text for discussion. The relative abundances and patterns of the dioritic liquids are, in general, appropriate for residual liquids from the anorthosites.
which have relatively high proportions of poikilitic pyroxene will therefore be low in incompatible elements like P, Zr, and REE. Some trends on variation diagrams may therefore be the result of inadequate sample size rather than any geologic process. Nain anorthosites are typically light REE-enriched (Fig. 13) and show a positive Eu-anomaly, the size of which does not correlate simply with the total percent of plagioclase in a sample. In many plutons, large grain-size and irregular distribution of interstitial material contribute to much of the variation in the pattern and the size of the Eu-anomaly. Some fine-grained dikes from the Nain complex show little variation while nearby coarse-grained leuconorites with nearly identical color index show a wide variation (Wiebe, 1980b). Averages of many analyses from individual homogeneous plutons give a better estimate of the pluton composition, particularly with respect to elements excluded from plagioclase. The average values of CaO, ST, (Sr x 103)/Ca, Mg#, and normative An are listed for several Nain plutons in Table 3. Layered plutons tend to have higher Mg# and normative An compared with the other types. The relatively low Mg# of all anorthositic rocks indicates that they are not simply accumulations of plagioclase in a mantle-derived melt that retains its primary Mg#. Although a strong positive correlation between Mg# and An is not apparent among individual samples from a single pluton (Table 2), it is strongly apparent when the averages of several different Nain plutons (Bble 3) are plotted (Fig. 14).
R A . Wiebe
242 TABLE 3 Some compositional characteristics of different plutons in the Nain complex, Labrador Pluton 1 2 3 4 5 6 7 8 9 10 11
Type layered layered massive diapir diapir diapir massive massive diapir massive massive
%An
64.1 60.6 57.2 57.6 55.6 54.4 53.5 52.3 50.7 49.2 38.1
n
Mg# 62.6 56.2 56.6 55.8 54.3 52.4 52.2 47.8 37.5 40.0 18.3
7.2 8.0 8.5 8.0 9.9 10.9 9.1 9.6 9.2 8.9 9.5
39 24 50 28 17 18 12 20 18 15 6
Notes: %An and Mg# defined as in Table 2. n = the number of samples averaged.
1: 2 3 4:
Leucotroctolite, Paul Island, Fig. 4c. Port Manvers Run - lower zone, Fig. 4b. Upper leuconorite, Paul Island, Fig. 4c. Northern leuconorite, Paul Island, Fig. 4c. 5: Kheovik leuconorite, Fig. 4b. 6: Lower leuconorite, Paul Island, Fig. 4c.
7: Port Manvers Run - upper zone, Fig. 4b. 8: Slambang leuconorite, Fig. 4b.
9: Tunungayualok Island older anorthosite, Fig. 4a. 10: Rnungayualok Island leuconorite, Fig. 4a. 11: Uivakh leuconorite, Fig. 4a.
For these same plutons, the Sr/Ca ratio increases as An decreases (Fig. 15), and there is considerable scatter in Sr/Ca at low An. For a given An-content, the diapiric plutons typically have higher Sr/Ca: the three plutons with the highest Sr/ Ca (6, 12, and 13) all appear to be relatively early diapiric bodies with significant internal deformation and the widespread occurrence of HAOM. The most sodic pluton with relatively low Sr/Ca (#11) is a younger massive leuconorite body with a chilled leuconoritic margin and a CI of 20-25. If, as seems likely, these plutons have ultimately been generated from similar parental magmas that have accumulated plagioclase to varying degrees, then this covariation of Mg# and An in different magma batches (plutons) appears to require the cotectic separation of mafic minerals (olivine and pyroxenes) and plagioclase (or at high pressures, clinopyroxene + spinel). The strong increase in Sr/Ca as An decreases within each magma batch strongly favors early removal of the high pressure assemblage, clinopyroxene + spinel, rather than plagioclase. The scatter in the trend of Sr/Ca (Fig. 15) could reflect some separation of plagioclase as well as clinopyroxene in the fractionation that produced each batch of magma: at a given An, magma batches with lower Sr/Ca ratios may have been affected more by prior plagioclase fractionation. It is also possible some of the scatter a t low An is due to mixing-in of crustal melts. In the Nain complex, coarse-grained inclusions up to 50 m in length with relict high-pressure mineral assemblages occur in a leucotroctolite body (Huntington, 1980). The dominant high pressure cumulus mineral assemblages are:
Proterozoic anorthosite complexes
243
i
/
%An Fig. 14. Normative (CIPW) An versus Mg# [= cation lOOMg/(Mg + F ~ T )for ] the average compositions of several Nain plutons. The data and key to the numbers are listed in Table 3. Also shown are troctolitic cumulates at the 15 and 80 percent-solidified levels in the Kiglapait layered intrusion (KI-15 and KI-80, respectively) (Morse, 1981). The dashed line is the trend of cumulates from the Skaergaard layered intrusion (Wager and Brown, 1968). If anorthositic plutons have been generated from a similar parental magma, the compositional variation between different plutons requires crystallization of a mafic phase at depth prior to emplacement as plagioclase-rich magmas.
;I 10
11.
05 8.
. 9 7. 10.
.3 40.2
6
630
Fig. 15. Average weight % normative An versus (Sr x lo3)/& for the same plutons listed in Table 3 and shown in Fig. 14. Also shown are the compositions of individual samples from two other anorthositic plutons: the Susie Brook Slab (12) and the Bird Lake Massif (13) (Simmons and Hanson, 1978). The strong increase in Sr/Ca as An declines can best be explained by substantial crystallization of Ca-rich clinopyroxene (probably along with olivine and spinel) prior to the emplacement of the individual batches of magma that created the separate plutons. The scatter in the trend may reflect the variable role of plagioclase fractionation at depth-samples that define the steepest increase having been least affected by plagioclase fractionation.
244
R.A. %ebe
(1) pleonaste spinel + augite + high-Al orthopyroxene with exsolved plagioclase lamellae (HAOM) and (2) plagioclase + augite + pleonaste spinel (Huntington, 1980; Wiebe, unpublished data). Well developed coronas of secondary olivine and plagioclase occur between spinel and augite primocrysts. These inclusions with coexisting augite and pleonaste spinel may represent early high-pressure cumulates from the troctolitic magma that were brought from depths of a t least 35 km (Huntington, 1980). Their occurrence lends further support to the role of augite + spinel fractionation as an explanation for the compositional variation between different anorthositic plutons in the Nain complex. It also suggests that cumulates with appropriate ultramafic assemblages do occur at great depth below the Nain complex. Cumulate troctolites from the Kiglapait intrusion (Morse, 1981, 1982b) have distinctly lower Sr/Ca than anorthositic rocks at similar normative An compositions (Fig. 15). If the parental magma for the troctolitic Kiglapait intrusion is similar to those that produced the troctolitic anorthosites, an explanation is needed to explain this difference in Sr. Emslie (1985) has suggested that high pressures might increase the partition coefficient (plagioclase/melt) of Sr. If true, it suggests that varying proportions of the plagioclase within the anorthosites were crystallized at depth and carried upward, whereas all (or most) of the Kiglapait plagioclase crystallized within the high-level Kiglapait magma chamber. The chemical compositions of anorthositic and mafic rocks found along the margins of anorthositic plutons may provide information concerning the nature of parental magmas to the anorthosites, and Emslie (1980) has suggested that some of the fine-grained gabbroic rocks along the margin of the Harp Lake complex could be parental to the anorthosites. In the Nain complex, rocks along the margins of anorthositic plutons show a very wide range in composition (ranging between basaltic, anorthositic and dioritic), and many marginal rocks are clearly affected by variable amounts of fractionation, crystal accumulation and contamination from the adjacent country rock. To date, none of the finer-grained mafic rocks found along the margins of the anorthosites have been identified unambiguously as appropriate parental magmas; most of them are likely, however, to have been comagmatic with the anorthosites. Some of the closely associated Feand Ti-rich dioritic rocks have compositions which suggest they crystallized from residual liquids from the anorthosites (see section on diorites).
Nd, Sr, AND Pb ISOTOPIC COMPOSITIONS
The isotopic compositions of most massif anorthosites suggest that they were generated by mantle-derived magmas. One of the earliest studies of Sr isotopes in anorthositic rocks from several complexes (Heath and Fairbairn, 1968) found a range in Sri (s7Sr/s6Sr initial isotope ratio) from 0.703 to 0.706, comparable to continental basaltic rocks. More recent studies of Sr isotopes in anorthositic rocks have established an even wider range in Sri that suggests some involvement of continental crust in addition to a dominant mantle source (e.g. Simmons et al.,
Proterozoic anortlzosite complexes
245
1985). Many anorthositic bodies have positive 6Nd values (Ashwal and Wooden, 1983; Demaiffe et al., 1986), suggesting derivation from a source with long-term depletion in light R E E comparable to sources of recent mid-ocean ridge basalts (Ashwal and Wooden, 1985) (Fig. 16). As more detailed isotopic studies of anorthositic rocks have been undertaken, there has been growing awareness of variation in initial isotopic ratios on many scales: between different plutons, within a single pluton, and between coexisting minerals in a single hand-specimen. For example, initial Pb isotope ratios in different bodies of the Rogaland complex (Norway) plot along an array with a slope corresponding to an age of 1470 f 250 Ma, considerably older than the currently accepted emplacement ages of the complex (Weis, 1986). One possible explanation of this array is mixing between a mantle source and a much less radiogenic crustal contaminant (Weis, 1986). Emslie (1985) was among the first to note that olivine-bearing anorthositic rocks characteristically have more primitive isotopic compositions (lower Sri and higher ENd) than orthopyroxenebearing anorthosites and attributed this to crustal contamination by felsic material. Isotopic disequilibrium (in both Sm-Nd and Rb-Sr systematics) between coexisting plagioclase and pyroxenes in single hand-specimens has been interpreted as due to progressive crustal contamination of the anorthositic magma during crystallization and emplacement (Ashwal and Wiebe, 1989). Progressive contamination is also suggested by the fact that high-Al orthopyroxene megacrysts and associated mega-plagioclase, thought to have crystallized early at depth, typically have more primitive isotopic compositions than the anorthositic rocks that contain them (Weis, 1986; Menuge, 1988).
‘lo/.EALY
-o l[;
- ”7 0 0
MTNS
L1(,:jl)] I 705
I
T Sri
710
715
Fig. 16. The ranges of Sri and €Ndi for some unmetarnorphosed anorthosites from Labrador and south Norway. Sources of data: Nain - Simmons et al. (1985); Harp Lake and Mealy Mountains - Ashwal et al. (1986); Rogaland - Dernaiffe et al. (1986) and Menuge (1988). The wide range of some complexes may reflect crustal contamination. Mantle sources probably range from depleted to enriched.
246
R.A. Wiebe
Regional isotopic studies of massif anorthosites in eastern Canada have led to the recognition of a major isotopic discontinuity located along the Grenville Front (Ashwal and Wooden, 1985; Ashwal et al., 1986). Anorthositic rocks in the Grenville Province characteristically have positive ENd and Sri less than 0.703, whereas northwest of the Grenville Front they typically have negative ENd and Sri greater than 0.704. Ashwal and Wooden (1985) attribute this contrast largely to the effect of contamination by much older continental crust north of the Grenville Front in the Churchill and Nain provinces. Although the overall variability of isotopic compositions strongly suggests some crustal contamination, it also has been argued that the isotopic discontinuity mainly records differences in the isotopic character of the mantle on either side of the Grenville Front (Hamilton and Morse, 1988).
OXYGEN ISOTOPES
Bylor (1968) established that most massif-type anorthosites have 1sO/160ratios in the range, d " 0 = 5 to 7%0-values that are similar to those of mantlederived mafic rocks. Isotopic compositions in this range have more recently been reported for both the Rogaland and Nain anorthosites (Demaiffe and Javoy, 1980; Simmons et al., 1985). Nonetheless, a few anorthosite massifs (e.g., the Adirondack anorthosites) have somewhat higher values (Thylor, 1968) that might have resulted from magmatic assimilation of lower crustal material (Morrison and Valley, 198s) or isotopic exchange during later metamorphism that affected the Adirondack anorthosites (Thylor, 1968). In the Rogaland and Nain complexes, the SlsO of associated granitic rocks are generally higher than the anorthosites in the range 7 to 9%0; these values are likely to reflect either crustal contamination or a crustal origin for these rocks (Demaiffe and Javoy, 1980; Snyder et al., 1988).
DIORITES
Although they represent only a small proportion of any complex, Fe-rich dioritic rocks (including rocks termed jotunites and monzonorites in some complexes) occur widely in close association with anorthositic plutons - either as sharply cross-cutting dikes, as small massive bodies along the margins of the anorthosites, or as separate, commonly layered, Felrich plutons (Fig. 17a). Some massive leuconoritic plutons (e.g., the Hidra body in the Rogaland complex Fig. 2) grade to margins of finer-grained dioritic rocks (Demaiffe and Hertogen, 1981). Dioritic rocks, including chilled pillows and cumulates, also occur along the margins or beneath some granitic bodies; commingling and hybridization between dioritic and granitic magmas are typical in this setting (Wiebe, 1980a) (Figs. 17b and 17c). In the Nain complex, the Tigalak layered intrusion (Wiebe and Wild, 1983) and the Goodnews complex (Wiebe, 1979b) both display exten-
247
Fig. 17. Field relations of dioritic rocks in the Nain complex. a. Typical fine-scale layering in dioritic cumulate bodies. A small inclusion of leuconorite lies below the hammer (Wiebe, 1978). b. The hammer rests on a strongly chilled pillow of Fe-rich diorite contained within texturally heterogeneous granitic rocks with variably digested inclusions of diorite. c. A small diorite inclusion (lacking a chilled margins and containing scattered xenocrysts of alkali-feldspar) occurs within a coarser-grained hybrid dioritic rock with abundant xenocrysts of alkali-feldspar and scarce quartz (Wiebe, 1990a).
248
R.A. Wiebe
sive interactions between diorite and granite and both have basal dioritic rocks that locally grade downward into underlying massive leuconoritic anorthosite. In the Laramie anorthosite complex, the Maloin Ranch pluton, a composite layered intrusion containing Fe-rich diorites, monzonites and syenites, displays abundant evidence for commingling between mafic and silicic magmas (Kolker and Lindsley, 1989). The textures of the dioritic rocks vary widely depending on whether they are chilled rocks (dikes and pillows in granite) or cumulates (Fig. IS), or partially hybridized with granitic material (Wiebe, 1990a). The dominant minerals in the most common Nain diorites are plagioclase, inverted pigeonite, augite and Fe-Ti oxides (ilmenite and magnetite). Comparable Fe-rich rocks within the Newark Island layered intrusion are troctolitic and have olivine rather than a low-Ca pyroxene (Wiebe, 19SS). Fe-rich olivine, ilmenite, and apatite are prominent in the most evolved dioritic rocks. In some complexes similar Fe-rich dioritic rocks also contain alkali-feldspar and are termed monzonorites (Duchesne, 1990). Hornblende and biotite are generally scarce except in diorites that are closely associated with hydrous granitic rocks. Plagioclase is typically in the range An45-30 and normally zoned. Fe-rich pyroxenes (mainly inverted pigeonite and ferroaugite) plot on a continuation of the trends of pyroxenes from anorthositic rocks (Fig. l l ) , generally display coarse exsolution, and lack zoning in Fe/Mg - probably as a result of subsolidus re-equilibration. Dioritic rocks have chemical compositions that are typically high in Fe, Ti and P (Bble 4). The valid interpretation of diorite composition o r trends of diorite compositions on variation diagrams requires great care in distinguishing between cumulates, chilled liquids and hybrids. On the whole, diorites have lower Mg# and normative A n than most anorthosites and plot on an extension of the trend of Mg# vs. An that is shown in Fig. 13. The chilled diorites (liquids) are higher in R E E than the anorthosites and have similar LREE enrichments with or without a negative Eu-anomaly (Wiebe, 1984; Mitchell et al., 1991). Although many dioritic rocks (possibly cumulates?) have high Sr (Emslie, 1978), dioritic liquids associated with specific anorthositic bodies in the Nain complex are typically lower in Sr (Wiebe, 1990a). In the Nain complex there is no evidence that the dioritic liquids have fractionated to granite. In any event, the amount of granite that could be produced would be very small compared with the total volume of granitic rocks typically associated with the anorthosite complexes. Nonetheless, monzonoritic rocks in the Rogaland complex (Duchesne, 1990) are thought to show a continuous compositional trend extending to quartz mangerites (granites). The Sr- and Nd-isotopic compositions of the dioritic members of massif anorthosites show wide variations in different complexes. In many complexes, Sri is higher and initial ENd is lower than associated anorthosites, suggesting some crustal involvement in their origin. Based in part on Sr isotopic differences between anorthosite and dioritic rocks, Duchesne et al. (1985) have proposed that the Rogaland dioritic rocks originated by anatectic melting of Fe-rich sources in
Proterozoic anorthosite complexes
249
Fig. 18. Photomicrographs of dioritic rocks. The length of each is 9 mm. a. p p i c a l texture of a chilled dioritic pillow with moderately tabular plagioclase, two pyroxenes and disseminated ilmenite, magnetite and apatite. b. Cuniulate diorite from the Tigalak layered intrusion with cumulus plagioclase, inverted pigeonite, augite, and ilmenite.
R A . Wiebe
250
TABLE 4 Representative chemical analyses of dioritic rocks from the Nain complex Chilled pillows
Layered rocks
Spec
P67B
P102
P63
SiOz
51.86 2.88 13.75 1.69 13.30 0.22 3.65 7.04 3.44 1.49 0.62 1.25 101.19
47.39 3.19 13.13 2.82 13.53 0.27 4.27 8.01 3.57 0.89 0.83 1.21 99.11
46.88 45.58 3.55 3.84 12.80 13.41 5.03 7.36 12.53 13.62 0.28 0.27 4.74 3.76 8.84 8.46 2.86 3.02 0.52 0.71 0.87 0.75 0 . 9 6 1.07 - 101.44 100.27
11 366 1162 370 172 22
6 450 1285 480 181 22
12 392 626 140 226 32
Ti02
Alz0 Fez 0
3 3
FeO MnO MgO G O Na2 0
Kz0 PZOS LO I Total Rb Sr
Ba Zr V Ni
P32
16 337 636 265 175 18
P66 50.37 2.39 15.52 4.99 9.48 0.20 4.77 7.31 3.57 0.12 0.52 0.77 100.61 7 537 771 223 186 22
P18 49.18 2.90 14.01 3.44 12.12 0.21 4.85 7.09 2.82 1.14 0.28
1.71 8.81 29.12 17.69 11.23 22.09 2.45 5.47 1.44
5.37 30.85 17.54 14.78 12.05 7.07 4.18 6.19 1.97
0.90 3.06 24.08 22.01 13.95 20.21
2.65 4.23 25.74 19.43 14.32 13.48
2.99 4.26 30.25 24.23 7.16 18.10
7.26 6.71 1.73
10.76 7.35 2.04
Mg/(Mg+FeT)
0.31
0.32
0.32
An/(An+Ab)
0.38
0.36
0.48
hY 01
mag ilm aP
-
46.31 3.61 12.73 5.73 13.28 0.27 6.16 8.25 2.82 0.47 0.44
1.05 1.41 101.18 99.45
30 393 610 215 264 56
CIPW norms
qtz or ab an di
P2A
7 335 387 82 318 44
-
P3E 42.16 4.51 13.48 5.48 13.40 0.27 5.05
10.24 2.23 0.31 2.23 0.99
100.35 2 405
307 60 221 18 1.06 1.84 18.98 26.01 8.77 21.51 7.99 8.62
7.25 4.55 1.21
2.39 6.87 24.34 22.65 9.49 22.90 5.09 5.62 0.66
2.77 23.83 20.66 14.33 21.61 0.51 8.30 6.96 1.02
0.26
0.38
0.36
0.37
0.33
0.43
0.44
0.48
0.46
0.58
-
-
5.21
the lower crust. The Sr isotopic differences could also be explained by variable crustal contamination. Although the origin of the dioritic rocks remains controversial, recent studies of different complexes have provided increased support for their comagmatic relationship to the anorthosites (e.g. Nain: Hill, 1982; Wiebe, 1990a; Adirondacks: Whitney, 1989; Laramie: Kolker et al., 1990). The common absence of negative Eu-anomalies in chilled dioritic liquids has led some workers to reject them as possible residual liquids from the anorthosites. However, if extensive crystallization
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of clinopyroxene occurred before and concurrently with plagioclase fractionation, residual liquids need not have negative Eu-anomalies (Emslie, 1985). The isotopic variability in both diorites and anorthosites seems best explained by crustal contamination both at depth and at the final level of emplacement (Simmons et al., 1985). If this is true, isotopic evidence alone cannot readily preclude or demonstrate a comagmatic relationship between anorthosite and diorite. Although the proposed fractionation of monzonoritic magma to granite in the Rogaland complex is permitted by available data (Duchesne et al., 1985), the isotopic differences between other monzonoritic and granitic rocks in that complex (Demaiffe et al., 1986) and the widespread occurrence of commingling between diorite and granite (Wiebe, 1984) suggest that fractionation from monzonorite to granite is not a dominant process.
GRANITIC ROCKS
Granitic plutons are closely associated with many anorthosite complexes, and, in North America, large areas are underlain by similar granitic plutons occurring independently of anorthosite. These granites and anorthosites define a broad transcontinental belt of 1.4 to 1.5 Ga anorogenic complexes that trends from Labrador into the southwestern U.S. (Anderson and Bender, 1989). The granitic rocks appear to be broadly contemporaneous with the anorthosites. Although, granite typically intrudes associated anorthosite, it also occurs commingled with diorites that grade to anorthosite. In the Nain complex, some anorthositic plutons also intrude composite mixtures of granite and diorite (Wiebe, 1988). Precise U-Pb zircon ages confirm the general contemporaneity of granitic, dioritic and anorthositic rocks in the Nain complex (Simmons et al., 1986; Simmons and Simmons, 1987). Granitic and charnockitic plutons associated with the Nain, Rogaland and Harp Lake (Emslie, 1980) complexes tend to be massive and commonly contain roughly equal proportions of alkali-feldspar, plagioclase, and quartz. The three-dimensional shape of most plutons are not well known, but some bodies in the Nain complex appear to be relatively thin curved sheets that occur structurally above the anorthosites (Wheeler, 1968; Wiebe, 1980; Hill, 1982). In other complexes (e.g., the Adirondacks, Laramie), well layered monzonitic and syenitic plutons appear to be of a cumulate origin (Hargraves, 1968; Fuhrman et al., 1988; Kolker and Lindsley, 1989). The granitic rocks may be either hypersolvus or subsolvus with perthitic or mesoperthitic alkali-feldspar (Collerson, 1982). Plagioclase compositions are mainly in the range A n 3 ~ - ~and ~ , rapakivi textures are common (Emslie, 1978). Fayalitic olivine and Fe-rich pyroxenes are common in many granites, but many granites have hornblende and biotite either exclusively or as rims to olivine and pyroxenes. In the Nain complex, many plutons have accessory apatite, allanite, fluorite and zircon. Mineral compositions of granitic rocks commonly produce
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continuous trends with minerals in associated dioritic and anorthositic rocks, and this continuity is commonly cited as evidence for a comagmatic relationship between all of these rocks (Fuhrman et al., 1988). Hypersolvus feldspars, relatively high mafic content, and common occurrence of augite and fayalite suggests that parental magmas to the granites were relatively hot and dry compared with magmas that produced calc-alkaline granites. The chemical compositions of anorogenic granitic rocks are typically high in K20 and Fe/Mg and low in Ca, Mg, and Sr relative to calc-alkaline orogenic granitic rocks at similar S i 0 2 levels. They can generally be classified as A-type granites (Whalen e t al., 1987). R E E are extremely variable depending particularly on whether the granitic rocks represent dominately cumulates or liquids. Patterns typically show light R E E enrichment with La between 100 and 1000 x chondrites and may have either positive or negative E u anomalies (Anderson, 1983; Anderson and Bender, 1989). In the Newark Island layered intrusion, most granitic cumulates and dikes (liquids) show negative E u anomalies (Wiebe, 1988). Available isotopic data (Sr, Nd, and Pb) from many complexes indicate that granitic rocks are crustal melts rather than differentiates from magmas that produced the anorthosites and diorites (Emslie, 1978; Demaiffe et al., 1986; Weis, 1986; Geist e t al., 1990).
Fe-Ti OXIDE DEPOSITS
Fe-Ti oxides form distinctive and often economically important bodies in many massif anorthosites. Major deposits occur within the Rogaland complex (Tellnes - Wilmart et al., 1989), the Adirondacks (Sanford Hill - Gross, 1968) and in Quebec (Lac Tio - Hammond, 1952). The deposits typically occur either as conformable layers within Fe-rich gabbroic or dioritic rocks or as massive bodies that cut host rock anorthosite sharply or act as a matrix to blocks of anorthositic rocks. No large bodies have been located within the Nain complex, though many small concentrations occur within layered Fe-rich dioritic and troctolitic rocks. Most ore bodies are dominated by ilmenite, but complexly exsolved titanomagnetite is dominant in some (Anderson, 1968). Silicate minerals (plagioclase, Fe-rich pyroxenes and olivine) occur in varying proportions up to 20-30%. Apatite is an important phase in many small bodies termed nelsonites (Kolker, 1982). The occurrences and mineralogy of Fe-Ti deposits suggest they are comagmatic with the anorthositic rocks. The two models most commonly proposed for their origin are crystal accumulation from residual liquids and liquid immiscibility. Layered conformable bodies can readily be explained as cumulates from residual Fe-rich liquids, and although the massive, cross-cutting bodies have been interpreted as remobilized oxide-rich cumulates (Emslie, 1975a), in recent years there has been increasing support for the role of liquid immiscibility. Melting experiments on an Fe-rich pegmatitic troctolite associated with Fe-Ti oxide deposits in the Laramie complex have produced immiscible silicate and oxide liquids (Lindsley and Frost, 1990). The high temperature necessary to produce an oxide
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melt can be lowered by the presence of a flux like apatite (Philpotts, 1967) or carbon (Weidner, 1982). Although apatite is not present in many massive oxide bodies, graphite has been found in Fe-Ti oxide bodies in the Laramie anorthosite (Bolsover and Lindsley, 1983).
LATE BASALTIC DIKES
Alkali to transitional tholeiitic basaltic dikes cut many anorthosite complexes, including the Harp Lake, Nain, and Rogaland complexes. They commonly appear to be only slightly younger than the anorthosites they cut, and their chemical compositions bear some similarities to Fe-Ti-rich diorites and troctolites emplaced contemporaneously with the anorthosites (Meyers and Emslie, 1977; Wiebe, 1985; Carlson et al., 1992). The Nain dikes, with relatively high normative olivine and low clinopyroxene, have the same troctolitic character found in liquids that were intruded at earlier stages during the emplacement of the anorthosites (Berg, 1980; Morse, 1981; Wiebe, 1988). Because of this similarity, these dikes may represent the waning stages of the magmatism that produced the anorthosites. Some compositional variation in the dikes suggests that they may also have been affected by fractional crystallization of clinopyroxene at depth (Wiebe, 1985).
THE GENERATION OF MASSIF-TYPE ANORTHOSITES
Anorthosite massifs were probably generated as a result of the ponding and fractional crystallization of basaltic or ultramafic magmas near the base of a stable crust (Emslie, 1985; Hoffman, 1989; Olson and Morse, 1990). The blanketing effect of a Late Proterozoic supercontinent on the underlying mantle may have reduced heat loss and caused broad convective upwelling in the underlying mantle beneath areas that were thousands of kilometers in diameter (Sawkins, 1976; Hoffman, 1989). The high temperatures of the anorthositic plutons and the widespread occurrence of orthopyroxene megacrysts with high Cr and Ni suggest that a basaltic magma was involved at some stage of their generation. A mantle origin for the anorthosites is also supported by their isotopic compositions and the common presence of large volumes of isotopically distinct granitic rocks that must have been derived from crustal melting. If anorthosites were derived from MORBtype mantle, extensive fractionation of ultramafic cumulates is required to explain the relatively low Mg# of nearly all rocks associated with the anorthosites. In addition, the positive correlation of An and Mg# and inverse correlation with Sr/ Ca in different anorthositic plutons require continuing fractional crystallization of clinopyroxene and pleonaste spinel to explain compositional differences between individual plutons. Ultramafic cumulates are absent a t or near the final level of emplacement and may exist near the base of the crust. In the Nain complex, appropriate high-pressure ultramafic cumulates (augite spinel HAOM) occur
+
+
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as inclusions in olivine-bearing anorthosites. The low content of augite in many anorthosite complexes probably results from fractional crystallization of augite at high pressures near the base of the crust where the augite field is expanded relative to its low pressure volume (Thy, 1983; Emslie, 1985). Although it is possible that the source of the anorthositic rocks could have been an unusually Fe-rich mantle low in clinopyroxene (Morse, 1982a; Olsen and Morse, 1990), the compositional variation between individual plutons in the Nain complex suggests that fractional crystallization of augite and pleonaste spinel at depth was important. It is also useful to note that in the Newark Island intrusion (Wiebe, 1988) liquids with the highest Mg# have much higher normative clinopyroxene than the more evolved, lower Mg#, Fe-Ti-rich liquids. If the mantle had been unusually low in clinopyroxene, the more primitive liquids should have been lower in clinopyroxene. Fractional crystallization of basaltic magma, ponded a t the base of the crust, should produce ultramafic cumulates on a chamber floor and eventually lead to saturation in plagioclase. Just prior to saturation, the liquid should be somewhat richer in plagioclase than at low pressure (Emslie, 1971), and the effect of high pressure and earlier crystallization of clinopyroxene should shift the equilibrium composition of plagioclase toward albite (Green, 1969; Fram and Longhi, 1991b). The resulting plagioclase of intermediate composition should float readily in basaltic magma at the base of the crust (Kushiro, 1980), and accumulate near the tops of the chambers. If, as seems likely, the chambers were periodically replenished by primitive magma, then some suspended plagioclase probably would be remelted. Through time, there could have evolved a stratified magma chamber in which the uppermost liquids became enriched in plagioclase components and still contained high proportions of suspended plagioclase crystals (Wiebe, 1990b). Similar resorption of plagioclase on a much smaller scale has been suggested as an explanation for the increasing Sr contents of evolved ocean-ridge basalts (Flower, 1984; Elthon, 1984). Resorption is likely to be even more effective in deep chambers beneath a stable craton because heat should b e lost from the chamber much more slowly. Because magmas related to the anorthosites appear to be low in H20, a decrease in pressure should lower their liquidus temperatures. When relatively low-density, hyperfeldspathic liquids with suspended plagioclase move upward from lower crustal chambers, they may initially lose heat relatively rapidly, heating the crust they pass through and crystallizing more plagioclase. Because of extensive crystallization, these early batches of magma may rise mainly as crystal-rich diapiric mushes (Longhi and Ashwal, 1985). Later batches of magma will probably follow previously heated paths, rise more rapidly, and may more nearly approach an adiabatic path. If very little heat is lost to the crust from the rising anorthositic magmas, some suspended plagioclase could partially melt during upward movement and cause the liquid to become even more hyperfeldspathic. While early anorthositic plutons may reach their final level of emplacement by diapiric movement of crystal-rich mushes, later batches of magma may rise with more liquid and
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establish magma chambers in the upper crust that are capable of internal convection. These later batches of magma probably produced the olivine-bearing layered anorthositic bodies and the massive leuconoritic bodies that are generally younger than the diapiric plutons. The younger massive leuconoritic plutons commonly have interstitial quartz and alkali-feldspar and isotopic compositions that indicate substantial pervasive crustal contamination. These characteristics may indicate that the upper crustal leuconoritic magma chambers acted as traps for crustally derived granitic melts.
ACKNOWLEDGEMENTS
I am greatly indebted to S.A. Morse for providing my initial opportunity to work in the Nain complex and for providing years of logistical support, counsel, and constructive criticism both in and out of the field. Although I remain responsible for the views expressed here, much of my sense of anorthosites has undoubtedly been influenced by extensive discussions both in and out of the field with S.A. Morse, J.H. Berg, R.E Emslie, L.D Ashwal, and J.C. Duchesne. J.H. Berg, ICC. Condie, R.E Emslie, and W.C. Phinney provided helpful reviews that materially improved the final version of this paper. Support was provided by NSF Grant EAR-8916045.
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Longhi, J. and Ashwal, L.D., 1985. n o - s t a g e models for lunar and terrestrial anorthosites: Petrogenesis without a magma ocean. Proc. 15th Lunar Planet. Sci. Conf., Part 2; J. Geophys. Res., 9 0 C571(2584. Maquil, R., 1978. Preliminary investigation on giant orthopyroxenes with plagioclase exsolution lamellae form the Egersund-Ogna anorthosite massif (S. Norway). Progress in Experimental Petrology, Fourth Progress Report. National Environment Research Council, Publ., Ser. D, 11: 144-146. Martignole, J. and Schrijver, K., 1970. Tectonic setting and evolution of the Morin anorthosite, Grenville Province, Quebec. Bull. Geol. SOC.Finl., 4 2 165-209. McLelland, J.M. and Chiarenzelli, J., 1990. Isotopic constraints on emplacement age of anorthositic rocks of the Marcy massif, Adirondack Mts., N.Y. J. Geol., 98: 19-41. Menuge, J.E, 1988. The petrogenesis of massif anorthosites: a Nd and Sr isotopic investigation of the Proterozoic of Rogalandwest-Agder, SW Norway. Contrib. Mineral. Petrol., 9 8 363-373. Meyers, R.E. and Ernslie, R.E, 1977. The Harp dikes and their relationship to the Helikian geological record in central Labrador. Can. J. Earth Sci., 14: 2683-2696. Michot, P., 1960. La gCologie de la catazone: le problerne des anorthosites, la palingtntse basique et la tectonique catazonale dans le Rogaland mkridonale (Norvkge mkridonale). Nor. Geol. Unders., 212: 1-54. Miller, J.D. and Weiblen, P.W., 1990. Anorthositic rocks of the Duluth Complex: examples of rocks formgd from plagioclase crystal mush. J. Petrol., 31: 295-339. Mitchell, J.N., Scoates, J.S., Kolker, A. and Ghazi, A.M., 1991. REE geochemistry of ferrodiorites and ferrogabbros in the Laramie anorthosite complex. EOS, %am. Am. Geophys. Union, 72: 305. Morrison, J. and Valley, J.W, 1988. Contamination of the Marcy anorthosite massif, Adirondack Mountains, N.Y.:petrologic and isotopic evidence. Contrib. Mineral. Petrol., 98: 97-108. Morse, S.A., 1969. The Kiglapait Layered Intrusion, Labrador. Geol. SOC.Am., Mem., 112, 204 pp. Morse, S.A., 1975. Plagioclase lamellae in hypersthene, Tikkoatokhakh Bay, Labrador. Earth Planet. Sci. Lett., 26: 331-336. Morse, S.A., 1981. Kiglapait geochemistry, IV The major elements. Geochim. Cosmochim. Acta, 45: 461-479. Morse, S.A., 1982a. A partisan review of Proterozoic anorthosites. Am. Mineral., 67: 1087-1100. Morse, S.A., 1982b. Kiglapait geochemistry, V Strontium. Geochim. Cosmochim. A d a, 46: 223-234. Morse, S.A. (Editor), 1983a. The Nain Anorthosite Project, Labrador: Field Report 1981. Univ. Mass., Dep. Geol., Geogr. Contrib., No. 40,153 pp. Morse, S.A., 1983b. Emplacement history of the Nain complex. In: S.A. Morse (Editor), The Nain Anorthosite Project, Labrador: Field Report 1981. Univ. Mass., Dept. Geol., Geogr. Contrib., 40: 9-15. Morse, S.A. and Nolan, K., 1984. Origin of stongly reversed rims on plagioclase in cumulates. Earth Planet. Sci. Lett., 68: 485-498. Olson, K.E. and Morse, S.A., 1990. Regional Al-Fe mafic magmas associated with anorthosite-bearing terranes. Nature, 344: 760-762. Philpotts, A.R., 1967. Origin of certain iron-titanium oxide and apatite rocks. Econ. Geol., 6 2 303-315. Phinney, WC., Morrison, D.A. and Maczuga, D.E., 1988. Anorthosites and related megacrystic units in the evolution of the Archean crust. J. Petrol., 2 9 1283-1323. Ranson, W.A., 1981. Anorthosites of diverse magma types in the Puttuaaluk Lake area, Nain complex, Labrador. Can. J. Earth Sci., 18: 26-41. Ribbe, P.H., 1975. Exsolution textures and interference colors in feldspars. In: P.H. Ribbe (Editor), Feldspar Mineralogy. Reviews in Mineralogy, Mineral. SOC.Am., 2: R73-R96. Ryan, B., 1990. Preliminary Geological Map of the Nain Plutonic Suite and Surrounding Rocks (NainNutak, NTS 14 S.W). Newfoundland Department of Mines and Energy, Geol. Sum. Branch, Map
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90-44, scale 1 :500 000. Salpas, P.A., Haskin, L.A. and McCallum, I.S., 1983. Stillwater anorthosites: a lunar analog? Proc. 14th Lunar Planet. Sci. Conf., Part 1;J. Geophys. Res., Suppl., 88: B27-B29. Sawkins, EJ., 1976. Widespread continental rifting: some considerations of timing and mechanism. Geology, 4: 427-430. Scoates, J., 1990. Syn-magmatic deformation of anorthosite: an assessment of the subsolidus evolution of anorthositic rocks in the 1.4 Ga Laramie anorthosite complex, Wyoming. Geol. SOC.Am., Abstr. Prog., 2 2 A300. Simmons, K.R. and Simmons, E.C., 1987. Petrogenetic inplications of Pb- and Sr-isotopic compositions for rocks from the Nain anorthosite complex, Labrador. Geol. SOC.Am., Abstr. Prog., 19: 845. Simmons, E.C. and Hanson, G.N., 1978. Geochemistry and origin of massif-type anorthosites. Contrib. Mineral. Petrol., 66: 119-135. Simmons, E.C., Snyder, G.A., Kalamarides, R.I. and Wiebe, R.A., 1985. Origins of massif-type anorthosites and related rocks - isotopic evidence from the southern Nain complex, Labrador. Geol. SOC. Am., Abstr. Prog., 17: 717. Simmons, K.R., Wiebe, R.A., Snyder, G.A. and Simmons, E.C., 1986. U-Pb zircon age for the Newark Island layered intrusion, Nain anorthosite complex, Labrador. Geol. SOC.Am., Abstr. Progr., 18: 751. Smithson, S.B. and Ramberg, I.B., 1979. Gravity interpretation of the Egersund anorthosite complex, Norway: its petrological and geothermal significance. Geol. SOC.Am. Bull., 9 0 199-204. Smithson, S.B., Shive, P.N. and Brown, S.K., 1977. Seismic reflections from Precambrian crust. Earth Planet. Sci. Lett., 37: 333-338. Snyder, D., 1984. Fractional Crystallization and Cumulate Processes in the Port Manvers Run Intrusion, Nain, Labrador. B.A. Thesis, Franklin and Marshall College, 41 pp. Snyder, G.A., Simmons, E.C., Kalamarides, R.I., Simmons, K.R. and Wiebe, R.A., 1988. Pb, Sr and 0 isotopic data for rocks from the Nain anorthosite complex, Labrador. Geol. SOC.Am., Abstr. Progr., 20: A118. Speer, J.A., 1975. The contact aureole of the Kiglapait intrusion. In: S.A. Morse (Editor), The Nain Anorthosite Project, Labrador: Field Report 1974. Univ. Mass., Amherst Contrib., 17: 17-26. Streckeisen, A., 1974. How should charnockitic rocks be named? In: J.C. Duchesne and J. Belliere (Editors), GCologie des Domaines Crystallins. SOC.Geol. Belg., Liege, pp. 349-360. Streckeisen, A, 1976. To each plutonic rock its proper name. Earth Sci. Rev., 1 2 1-33. Subramanian, A.P., 1956. Mineralogy and petrology of the Sittampundi complex, Salem District, Madras State, India. Geol. SOC.Am. Bull., 67: 327-379. Tanner, J.G., 1969. A Geophysical Interpretation of Structural Boundaries in the Eastern Canadian Shield. Ph.D. Thesis, University of Durham, 194 pp. Taylor, H.P., 1968. Oxygen isotope studies of anorthosites, with particular reference to the origin of bodies in the Adirondack Mountains, New York. In: Y.W. Isachsen (Editor), Origin of Anorthosite and Related Rocks. N.Y. State Mus. Sci. Serv., Mem., 18: 111-134. Thy, P., 1983. Phase relations in transitional and alkali basaltic glasses from Iceland. Contrib. Mineral. Petrol., 82: 232-251. Valley, J.W. and O’Neil, J.R., 1982. Oxygen isotope evidence for shallow emplacement of Adirondack anorthosite. Nature, 300: 497-500. Van der Molen, I. and Paterson, M.S., 1979. Experimental deformation of partially-melted granite. Contrib. Mineral. Petrol., 7 0 299-318. Vermaak, C.E, 1981. Kunene Anorthosite Complex. In: D.R. Hunter (Editor), Precambrian of the Southern Hemisphere. Developments in Precambrian Geology, 2. Elsevier, Amsterdam, pp. 578599.
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Wager, R.L. and Brown, G.M., 1968. Layered Igneous Rocks. Oliver and Boyd, Edinburgh, 588 pp. Weidner, J.R., 1982. Iron-oxide magmas in the system Fe-C-0. Can. Mineral., 2 0 555-566. Weis, D., 1986. Genetic implications of Pb isotopic geochemistry in the Rogaland anorthositic complex (southwest Norway). Chem. Geol., 57: 181-199. Whalen, J.B., Currie, K.L. and Chappell, B.W., 1987. A-type granites: geochemical characteristics, discrimination and petrogenesis. Contrib. Mineral. Petrol., 95: 407-419. Wheeler, E.P., 1942. Anorthosite and associated rocks about Nain, Labrador. J. Geol., 5 0 611-642. Wheeler, E.P., 1960. Anorthosite-adamellite complex of Nain, Labrador. Geol. SOC.Am. Bull., 71: 1755-1762. Wheeler, E.P., 1968. Minor intrusives associated with the Nain anorthosite. In: I.W. Isachsen (Editor), Origin of Anorthosite and Related Rocks. N.Y. State Mus. Sci. Sew., Mem., 18: 189-206. Whitney, P.R., 1989. Distribution, geochemistry, and origin of Adirondack jotunites. Geol. SOC.Am., Abstr. Prog., 21: A107. Wiebe, R.A., 1978. Anorthosite and related plutons, southern Nain complex, Labrador. Can. J. Earth Sci., 15: 1326-1340. Wiebe, R.A., 1979a. Anorthositic dikes, southern Nain complex, Labrador. Am. J. Sci., 279: 394-410. Wiebe, R.A., 1979b. Fractionation and liquid immiscibility in an anorthositic pluton of the Nain complex, Labrador. J. Petrol., 20: 239-269. Wiebe, R.A., 1980a. Commingling of contrasted magmas in the plutonic environment: examples from the Nain anorthosite complex. J. Geol., 88: 197-208. Wiebe, R.A., 1980b. Anorthositic magmas and the origin of Proterozoic anorthosite massifs. Nature, 286: 564-567. Wiebe, R.A., 1984. Commingling of magmas in the Bjerkreim-Sogndal lopolith (S.W. Noway): evidence for the compositions of residual liquids. Lithos, 17: 171-188. Wiebe, R.A., 1985. Proterozoic basalt dikes in the Nain anorthosite complex, Labrador. Can. J. Earth Sci., 2 2 1149-1157. Wiebe, R.A., 1986. Lower crustal cumulate nodules in Proterozoic dikes of the Nain complex: evidence for the origin of Proterozoic anorthosites. J. Petrol., 27: 1253-1275. Wiebe, R.A., 1988. Structural and magmatic evolution of a magma chamber: The Newark Island layered intrusion, Nain, Labrador. J. Petrol., 29: 383-411. Wiebe, R.A., 1990a. Dioritic rocks in the Nain complex, Labrador. Schweiz. Mineral. Petrogr. Mitt., 70: 199-208. Wiebe, R.A., 1990b. Evidence for unusually feldspathic liquids in the Nain complex, Labrador. Am. Mineral., 75: 1-12. Wiebe, R.A. and Wild, T, 1983. Fractional crystallization and magma mixing in the Tigalak layered intrusion, the Nain anorthosite complex, Labrador. Contrib. Mineral. Petrol., 84: 327-344. Wilmart, E., Demaiffe, D. and Duchesne, J.C., 1989. Geochemical constraints on the genesis of the Tellnes ilmenite deposit, southwest Norway. Econ. Geol., 84: 1047-1056. Windley, B.F., 1973. Archean anorthosites: a review with the Fiskenaesset Complex, West Greenland as a model for interpretation. Spec. Pub. Geol. SOC.S. Afr., 3 319-332.
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Chapter 7
THE ROLE OF ANOROGENIC GRANITES IN THE PROTEROZOIC CRUSTAL DEVELOPMENT OF NORTH AMERICA J.L. ANDERSON AND J. MORRISON
INTRODUCTION
Between 1.6 and 1.9 Ga, the Early Proterozoic North American craton underwent considerable crustal growth through punctuated orogenic episodes at approximately 1.82 to 1.90 Ga (Penokean and Hudsonian orogenies), 1.70 to 1.78 Ga (Yavapai and Ivanpah orogenies) and 1.61 to 1.68 (Labradorian and Mazatzal orogenies). Much of the new crust generated was juvenile, mantle-derived material with minimal reworking of the Archean craton (DePaolo, 1981; Nelson and DePaolo, 1985; Bennett and DePaolo, 1987; Anderson and Cullers, 1987). The only orogenic episode between 1.6 and 1.1 Ga was the 1.21 to 1.33 Ga Elzevirian orogeny of the southern Adirondack Mountains (Daly and McLelland, 1991). Throughout much of the continent, the 1.1 to 1.6 Ga age range is characterized by a wide spectrum of anorogenic igneous activity including emplacement of anorthosite massifs, charnockite intrusions, diabase dike swarms, and batholiths of potassic rapakivi granite. Sections of coeval rhyolite make up much of the basement of the middle and southern midcontinent (Bickford et al., 1981; Thomas et al., 1984; Bickford et al., 1986; Van Schmus et al., 1987) and portions of the Grenville Province (McLelland, 1986). By far, most of the magmatism occurred between 1.4 and 1.5 Ga, but repeated activity occurred at 1.34 to 1.39 Ga and 1.0 to 1.2 Ga (Anderson, 1983). While anorthosite and charnockitic rocks comprise much of the northeast portion of the continent (Labrador and portions of the Grenville Province), granite and lesser amounts of rhyolite are the most common expression of the magmatism in the central (midcontinent) and western regions. In volume and tectonic setting, the anorogenic magmatism is an expression of a Middle Proterozoic global event that has no clear analogue in younger geologic time. The objective of this paper is to review the compositional diversity of the granitic suites, their conditions of emplacement, and constraints on their origin, emphasizing their role in the large-scale differentiation of continental crust. An emerging view of the Proterozoic is that orogenic growth was globally episodic. This conclusion was originally proposed by Gastil (1960) and Sutton (1963), and has only recently gained renewed attention (Anderson, 1987a; Condie, 1989; Hoffman, 1989; Anderson and Bender, 1989). Major 1.6 to 2.1 Ga orogenies have been recognized on nearly every continent, including the Svecofennian orogen of northern Europe, the Pans-Amazonian orogen of South America,
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the Eburnean orogen of central Africa, and the Eastern Ghats orogen of India (Condie, 1989). Coeval with the 1.02 to 1.07 Ga Grenville orogeny of North America are the Dalslandian of the Baltic, Kibaran of South Africa, and the Satpura of India. Late Proterozoic orogenic growth is absent in North America but is well represented in Africa and South America as the 700 to 900 Ma Pan-African and Brazilian0 orogenies, respectively (see review in Condie, 1989 and Windley, 1984). Anorogenic igneous activity occurred in many areas and lasted for hundreds of millions of years. For example, the span of time between 2.0 and 2.6 Ga has yet, on a global scale, to reveal but a few regions undergoing orogenic crustal development and, instead, is marked by large mafic intrusions and diabase dikes, some of which coincide with intracontinental rifts (Windley, 1984). Exceptions to this generalization are striking, including 2.1 to 2.2 Ga orogenies in the Guiana and west Africa shields and 2.4 to 2.6 orogenies in India and China (Condie, 1989, 1990). Most of the world’s anorthosite, charnockite, and rapakivi granite (the “anorogenic trinity” of Anderson, 1983) was emplaced between 1.65 and 1.1 Ga, prior to the onset of the Grenville-aged orogenies, in a belt that trends across North America and southern Greenland into the Baltic region of northern Europe to as far east as the Ural Mountains and the Ukraine (Herz, 1969; Bridgwater and Windley, 1973; Emslie, 1978; Anderson, 1983; Aberg, 1988). Others occur in Peninsular India (Leelanandam and Reddy, 1988), Africa (Conradie and Schoch, 1986), Brazil (Dall’agnol et al., 1991; Lafon et al., 1991), China (Jianhua et al., 1991). Younger anorogenic plutonism in North America includes the 1.1 Ga Keweenawan intrusions and related volcanics of the Lake Superior region (Green, 1977) and some well known granitic intrusions such as the 1.03 Ga Pikes Peak batholith of Colorado (Barker et al., 1975). Similar-aged charnockite and rapakivi granite are post-tectonic to deformation in the Natal Belt of South Africa (Kerr, 1985; Eglington and Kerr, 1989). Anorthosite and charnockite ranging in age from 910 to 955 Ma are post-tectonic to the Sveconorwegian orogeny of Norway (Pasteels et al., 1979; Duchesne et al., 1985). The last significant production of anorogenic Proterozoic plutons occurs in the Middle East subsequent to the Pan-African orogeny (Jackson et al., 1984; Kuster and Harms, 1991). One striking aspect of the anorogenic magmatism is its near exclusive geographic restriction to crust formed in an earlier Proterozoic orogenic cycle. Phanerozoic analogues (Windley, 1989) are far less widespread but include anorogenic intrusions in Nigeria and rhyolites in Argentina, both of which are temporally related to opening of the south Atlantic.
ANOROGENIC MAGMATISM OF NORTH AMERICA
Anorogenic igneous rocks ranging in age between 1.40 and 1.49 Ga are widespread in North America and comprise a remarkable 4800 km long belt that extends from Labrador to California. Hosted primarily in Early Proterozoic orogenic crust, the plutons are often of batholithic size and most are undeformed
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excluding effects of intrusion and deformation of a much later age. Several occurrences in Labrador and elsewhere in northeast Canada have ages near 1.45 Ga and are comprised of anorthosite and younger charnockitic and granitic intrusives (Emslie, 1978; Emslie, 1980; Ranson, 1981; Ashwal and Wooden, 1983a; Wooden et al., 1987). Intrusions in Ontario and the northern midcontinent (including Wisconsin and basement drill holes in northern Illinois, Michigan, and Indiana) have U-Pb (zircon) ages ranging from 1.43 to 1.49 Ga (Van Schmus et al., 1975; Hoppe et al., 1983; Van Schmus et al., 1987). The 1.49 Ga Wolf River batholith of Wisconsin contains minor anorthosite and charnockitic rocks engulfed by multiple plutons of rapakivi granite (Anderson and Cullers, 1978; Anderson, 1980). Radiometric ages of rocks in the central midcontinent are based largely on samples retrieved from basement drillholes but include the St. Francois complex of southeast Missouri (Bickford et al., 1981). Over fourteen 1.44 to 1.46 Ga (ages from U-Pb, zircon) localities have been identified from Kentucky through southern Illinois, Missouri, and Kansas to southern Nebraska (Bickford et al., 1981; Bickford et al., 1986; Van Schmus et al., 1987). Granite and rhyolite are the most common rock types but anorthosite has been obtained from the subsurface near the Red Willow batholith of Nebraska (Lidiak, 1972). Over 40 mid-Proterozoic plutons occur in the western United States (southern Wyoming, Colorado, New Mexico, Arizona, southern Nevada, and southern California). Porphyritic monzogranite and syenogranite are by far the most common lithologies, but several of the intrusions have low silica portions (at 58 to 65 wt.% S O z ) that include quartz monzodiorite, quartz monzonite, and rare granodiorite; U-Pb ages cluster near 1.44 Ga but range from 1.40 to 1.46 Ga (see references in Anderson, 1983; Anderson and Bender, 1989). Gabbro, anorthosite, and charnockite are conspicuous by their absence in most regions of the central and western United States, exceptions being the Laramie anorthosite of Wyoming (Fountain et al., 1981; Goldberg, 1984; Kolker et al., 1990) and anorthosite (inclusions in Proterozoic granite) exposed in southern New Mexico (Hedlund, 1980). Alkalic rocks are also rare with the exception of two examples in southeastern California: the Mountain Pass carbonatite-shonkinite intrusion of the Clark Mountains (Anderson, 1983; E. DeWitt, personal communication, 1987) and the Barrel Springs intrusion of the Piute Mountains (Gleason et al., 1988). In Wisconsin, the Wausau syenite (with nepheline syenite) forms a peripheral intrusion to the Wolf River batholith (Van Schmus et al., 1975). Continuation of anorogenic magmatism at 1.34 to 1.39 Ga (mostly 1.37 to 1.39 Ga) formed an extensive terrain of rhyolite and shallow granite in the southern midcontinent. Bickford et al. (1981), Thomas et al. (1984), and Bickford et al. (1986) have reported U-Pb ages for over 25 eruptive centers and related intrusions, most of which occur in the subsurface of Missouri, Kansas, Oklahoma, southern New Mexico, and northern Texas. Surface exposures of this event include the Munger granite of the St. Francois Mountains, Missouri, the nay, Tishomingo, and Spavinaw granites of the Arbuckle Mountains, Oklahoma, and the San Isabel batholith of the Wet Mountains, Colorado.
Fig. 1. Distribution of Middle Proterozoic plutons and batholiths in North America (modified from Anderson, 1983). Numbered localities refer to names and ages of localities discussed in the text.
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Up to 75% of the world's massif anorthosite occurs in a restricted region from the Adirondack Mountains of New York northward through the Grenville Province to Labrador. As summarized by Ashwal and Wooden (1983b), some anorthosite was emplaced at ~ 1 . 4 Ga, 4 but most of the anorthosite is younger. McLelland et al. (1988a) and Daly and McLelland (1991) have demonstrated that much of the anorthosite, and related intrusions of charnockite and granite, have crystallization ages between the Elzevirian (1.21 to 1.33 Ga) and Ottawan (1.03 to 1.08 Ga) phases of the Grenville orogeny. The principal age range in the Adirondack Mountains (from U-Pb, zircon) is from 1.13 to 1.16 Ga. Though affected by post-intrusion granulite-facies metamorphism of the Grenviile orogeny, there is considerable agreement that these massive intrusions are anorogenic (McLelland, 1986; Windley, 1989; Emslie and Hunt, 1990). Anorthosite and charnockite of essentially the same age occur in Grenville rocks of the Blue Ridge, Virginia (Pettingill et al., 1984). A few anorogenic intrusions of similar and younger age occur elsewhere within the continent well away from the affects of the Grenville orogeny. The 1.22 Ga San Gabriel anorthosite and related jotunite and syenite (Carter and Silver, 1972; Ekstrom et al., 1991) occupy a considerable portion of the Transverse Ranges of southern California. Swarms of diabase in Arizona and California have approximately the same age based on U-Pb dating of granophyre (Silver, 1978; Hammond, 1986; Hammond and Wooden, 1990). The Keweenawan midcontinent rift is comprised a bimodal igneous suite with ages ranging from 1.09 to 1.12 Ga (Green, 1977). Between 1.03 to 1.05 Ga, a number of dispersed but equally impressive intrusions were emplaced, including the Pikes Peak batholith (Barker et al., 1975) of Colorado and the Enchanted Rock, Lone Grove, and Franklin Mountain batholiths of central and west Texas (Garrison et al., 1979; Muehlberger et al., 1966).
COMPOSITIONAL VARIATIONS AMONGST ANOROGENIC GRANITES
Mid-Proterozoic anorogenic granites are compositionally similar in many respects, and are often referred to as A-type, following the original definition of Loiselle and Wones (1979). Striking enrichment exists for several large ion lithophile elements (LILE) such as K, Rb, Ba, Th, and REE, high field strength elements (HFSE, including Ti, Ga, Nb, Y, and Zr), and fluorine. Relative to calcalkaline granites, much of this enrichment is notable at intermediate silica levels (64 to 71 wt.% Si02) and the contrast lessens for more siliceous members. Most are subalkaline to marginally alkaline (based on total alkalies; Irvine and Baragar, 1971), yet are ultra-potassic (K20 abundance greater than the high-K field of Gill, 1981). Alkali-lime indexes (Peacock, 1939) typically fall in the alkali-calcic or alkalic field due to low CaO and high K20. Elevated Fe/(Fe '+ Mg) ratios and low Sr (usually <300 ppm; exceptions include the San Isabel and Oak Creek batholiths of Colorado and the Hualapai granite of Arizona which range to higher Sr) are
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typical. The rocks even have a common appearance; most are red, coarse grained, and seriate to porphyritic with large alkali feldspar phenocrysts. Several have pyterlitic to wiborgitic textures typical of rapakivi granites. Mafic minerals are usually interstitial, and not prismatic, due to their late crystallization. Accessory minerals, in addition to sphene and Fe-Ti oxides, include allanite, apatite, and fluorite. Principal differences among the granites are (1) the degree of alumina saturation and (2) the Fe-Ti oxide mineralogy (Thble 1). Most are metaluminous (Fig. 2C) and contain biotite + sphene f hornblende (only the more evolved, silicic members are peraluminous), and can be divided into either the ilmeniteor magnetite-series of Ishihara (1977). Yet, in major regions of the continent, the granites are uniformly peraluminous (also Fig. 2C) and contain biotite as the sole mafic phase coexisting with muscovite and other aluminous phases. Metalum inous, ilm enite-series granites The vast majority of anorogenic granites are metaluminous and many are ilmenite-series. The distribution of low-fo,, ilmenite-series granites occupy much of Laurentia (North America and Baltic shields combined) including most of the classic rapakivi massifs of Finland (Ramo and Haapala, 1990). Whether or not the rapakivi granites of Sweden and southern Greenland belong to this series is currently unknown, but the dominance of ilmenite as the main Fe-Ti oxide phase of anorogenic granites continues across into North America, including some of the large biotite-hornblende rapakivi granite intrusions in Labrador (Mistastin Lake and Snegamook granites of Emslie, 1980; R.E Emslie, personal communication, 1991) and in the Adirondack Mountains of New York (Tupper Lake and other 1.10 to 1.15 Ga granites of McLelland and Chiarenzelli, 1990; J.M. McLelland, personal communication, 1991). Continuing to the southwest, other occurrences in North America include the Wolf River batholith (Anderson, 1980) of Wisconsin, the Pikes Peak batholith (Barker et al., 1975) of Colorado, and the portions of the Sherman granite (Edwards, 1991) of Wyoming. Rapakivi granites of South Africa (Kerr, 1985; Kerr and Thomas, 1991) and Brazil (Dall’agnol et al., 1991) are examples of ilmenite-series anorogenic granite outside of Laurentia. Uniformly, the ilmenite-series, metaluminous granites are ultra-potassic (ultrahigh K group in Fig. 2A), alkaline for low silica members, and have the highest observed Fe/Mg (high Fe/Mg group, Fig. 2B) ratios as a result of lower Mg, irrespective of silica. Fluorite, indicative of high fluorine activity, is an ubiquitous accessory mineral. The Wolf River batholith is composed of older anorthosite and mangerite intruded by eight plutons of rapakivi syenogranite and porphyritic monzogranite (Anderson and Cullers, 1978). Within the 1.03 Ga Pikes Peak batholith (Barker et al., 1975) are at least four magmatic lineages including (1) gabbro, (2) monzonite, syenite, quartz syenite, and riebeckite granite, (3) fayalite granite, and (4) biotite granite. Though much different in age, the batholith shows compositional affinity to the 1.4-Ga Laramie anorthosite-syenite and Sherman granite of Wyoming (Eggler, 1968; Fountain et al., 1981; Edwards, 1991). Ratios
Role of anorogenic granites in tlze Proterozoic crustal development of NorthAmerica
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Y+Nb ppm
wt.Z SiO,
Fig. 2. Elemental composition of anorogenic plutons. A. Weight percent silica versus K2O (0 = high K plutons including those of St. Francois Mountains, New Mexico, central Arizona, Log Cabin and Sherman granites, and 1.4 Ga midcontinent drill hole samples, x = all others; fields for low, medium, high, and ultrahigh K2O from Gill, 1981). B. Weight percent silica versus FeO/(FeO + MgO) (0 = high Fe/Mg ilmenite-series granites, x = intermediate Fe/Mg metaluminous magnetite-series granites plus data from the Silver Plume and St. Vrain batholiths; 0 = low Fe/Mg peraluminous granites plus 1.3 Ga midcontinent drill hole samples; field boundary for calc-alkaline and tholeiitic from Miyashiro (1974); all Fe as FeO). C. Weight percent silica versus molecular proportions of Al203/(CaO + Na2O K2O) ( X = data for two mica granites, o = data for biotite hornblende granites; peraluminous and Y using classification of Pearce et al. metaluminous boundary from Shand, 1947). D. Rb versus Nb (1984) (x = data for high Nb peraluminous granites, o = data for high Nb metaluminous granites, and 0 = data for low Nb granites of the St. Francois Mountains, the Oak Creek and San Isabel batholiths, and midcontinent drill hole samples; field labels are SYNCOLG = syn-collisional granite, WPG = within plate granite, VAG = volcanic arc granite, and ORG = ocean ridge granite. Sources of data include Eggler (1968), Barker et al. (1975), Condie (1978), Bickford et al. (1981), Anderson (1983), Anderson and Thomas (1985), Anderson and Bender (1989), Cullers et al. (1992a, b), and Anderson (unpublished).
+
+
LL. Anderson and J. Morrison
270 TABLE 1 Age of mid-Proterozoic plutons - continental U.S. a Exposed intrusions peraluminous
metaluminous
(M
Wisconsin
1. Wolf River batholith
Missouri
2. Silvermine granite (1.48) 3. Butler Hill granite (1.46) 4. Breadtray granite 5. Hawk Park granodiorite (1.48) 6. Graniteville granite (1.27) 7. Munger granite porphyry (1.38)
(m
Oklahoma
8. Tishomingo granite (1.37) 9. Blue River gneiss (1.40) 10. Troy granite (1.40) 11. Spavinaw granite (1.37)
Texas
12. Lone Grove granite (1.06) 13. Town Mountain granite (1.05) 14. Van Horn rhyolite (1.49) 15. Red Bluff granite (1.03)
Wyoming
16. Laramie anorthosite ( 9 17. Sherman granite (1.38) 18. Log Cabin batholith (1.39) 19. St. Vrain batholith (1.42) 20. Silver Plume batholith 21. Oak Creek batholith (1.44) 22. St. Kevin granite (1.39) 23. Curecanti granite (1.39) 24. Trimble granite (1.39) 25. Granite of Unaweep Cyn. (1.44)
26. Vernal Mesa granite (1.451 27. Eolus granite (1.46) 28. San Isabel batholic., 1 29. Pikes Peak bathc1i:n (1.03) 30. Electra Lake gabbro (1.42)
New Mexico
31. Ladron granite (1.32) 32. Obscura granite (1.34) 33. Capitol Peak granite (1.35) 34. Tusas Mountain granite (1.47)
35. Rana granite (1.44) 36. Sandia granite (1.44) 37. Priest granite (1.55) 38. Pedernal granite (1.47) 39. Los Pinos granite (1.48) 40. Sepultura granite (1.35)
Arizona
41. Dos Cabezas granite (1.38) 42. Stockton Pass granite (1.41) 43. Tungsten King granite (1.42) 44. Oracle granite (1.44) 45. Ruin granite (1.44) 46. Dells granite (1.40) 47. Lawler Peak granite (1.41) 48. Sierra Estrella granite (1.38) 49. Ak-Chin granite
50. Continental granodiorite 51. Fort Huachuca granite 52. Holy Moses granite (1.34) 53. Hualapai granite (1.37)
Colorado
(m
Nevada
'a
(m (w
54. Gold Butte granite (1.43) 55. Beer Bottle Pass granite (1.43) 56. Newberry granite (1.43) 57. Davis Dam granite (1.43)
Role of anorogenic granites in the Proterozoic crustal development of North America
271
TABLE 1(continued) Exposed intrusions peraluminous California
metaluminous
58. Dead Mountains granite (1.43) 59. Homer granite (1.43) 60. Marble granite (1.43) 61. Parker Dam granite (1.40) 62. Bowmans Wash qtz. monz. (1.41)
in Ga by U-Pb (underlined) and by Rb-Sr (italics). Peraluminous based on presence of primary muscovite and biotite o r A/CNK >1.05; metaluminous based on presence of biotite, sphene, fhornblende or A/CNK <1.05.
a Age
References: Silver and McKinney (1962); Steiger and Wasserburg (1966); Muehlberger et al. (1966); Peterman et al. (1967,1968); Silver and Barker (1967); Hansen and Peterman (1968); Mose and Bickford (1969); Bickford et al. (1969); Bickford and Mose (1975); Bickford and Cudzilo (1975); Van Schmus et al. (1975, 1987); Mukhopadhyay et al. (1975); Subbarayudu et al. (1975); Barker et al. (1976); Swan (1976); Kessler (1976); Shake1 et al. (1977); Silver (1978); L.?:Silver, in Stewart and Carlson (1978); White (1978); Register and Brookins (1979); Condie and Budding (1979); Brookins et al. (1980); Keith et al. (1980); Erickson (1981); Silver et al. (1981); Hoppe et al. (1983); Thomas et al. (1984); Wobus (1984); Anderson and Bender (1989); J.E. Wright, personal communication; M.E. Bickford, unpublished.
of FeO/(FeO + MgO) are generally in excess of 0.85 and range up to 0.98. These rocks have the highest abundances of Nb and Ga, the latter ranging to values greater than 20 ppm upwards to 95 ppm. This and other compositional features (high Ti, K, Ba, REE, and within plate granite abundances (Fig. 2D) of Y , Nb, and Rb) make the Wolf River, Pikes Peak, and Sherman indistinguishable from the classic rapakivi massifs of Finland (Sahama, 1948; Savolahti, 1956; Haapala and Ramo, 1990).
Metaluminous, magnetite-series granites Metaluminous, high-fo,, magnetite-series granites occupy a sizeable portion of North America and include those of the central and southern midcontinent, several in southern Colorado and New Mexico, and all occurrences in the lower Colorado River region of southern Nevada, western Arizona, and southern California. Magnetite-series rapakivi granites have also been observed in China (Jianhua et al., 1991), Estonia (firs et al., 1991), and Siberia (Moralev and Glukhovsky, 1991). A few of the North American examples are ultra-high K, but others, including those of the St. Francois Mountains, the Log Cabin batholith, granites in New Mexico, and 1.4 Ga plutons sampled by drill hole in the midcontinent are less potassium enriched and are “high K” (Fig. 2A). Compared to the ilmenite-series granites, many of these plutons are less Fe rich. Ratios of FeO/(FeO + MgO) typically range between 0.80 and 0.88 (intermediate Fe/Mg group, Fig. 2B). They also are less enriched in Ba and Ga. The plutons are largely metaluminous and contain biotite f hornblende. The more evolved portions of
J.L. Anderson and J. Morrison
272
these granites trend into peraluminous compositions but the rocks never contain muscovite or other markedly aluminous phases. Compared to the ilmenite-series granites described above, REE abundances also tend to be lower, such as in the St. Francois complex, with LREE ranging up to less than 200 times chondrite (Cullers et al., 1981). Fluorite occurs as an accessory mineral in some of these granites but is absent in many others. The following three examples depart from some aspects of the above summary. The 1.36 Ga San Isabel batholith (Cullers et al., 1992a) is fundamentally metaluminous and is the only magmatic epidote-bearing, mid-Proterozoic pluton known in the continent. The batholith has ultra-high K, yet it has the lowest silica range (56-65 wt.%) and thus, has the highest concentration of compatible elements (Ca, Mg, Fe, and Sr (310-570 ppm)). Its alkali-lime index is, like most others, alkali calcic ( 4 4 ) . The subsurface 1.34-1.39 Ga rhyolites of the southern midcontinent are also distinctive. Although potassic and possessing many of the compositional features of &type granites, the rocks have conspicuously low ratios of FeO/(FeO MgO), mostly between 0.64 and 0.83 (low F e M g group, Fig 2B) and lower abundances of Nb and Y. Finally, the 1.46 Ga Eolus pluton (Barker, 1969; Collier, 1989) of the Needle Mountains of Colorado has the very wide range in composition from 59.9 to 74.4% Si02 and has one of the higher Fe and lower K trends. Collier (1989) characterises the batholith as calc-alkaline but his data indicate that the suite has an alkali-lime index of -53, or alkali-calcic.
+
Peraluminous granites
Fundamentally peraluminous granites form a distinct subprovince from central Colorado (Silver Plume-type after the Silver Plume granite) to New Mexico and central Arizona (Oracle-type after the Oracle granite). These granites contain biotite or two micas, often lack sphene and fluorite, and have monazite as an accessory phase. Magnetite is the principle Fe-Ti oxide. Ilmenite is less abundant than magnetite or is absent. Magmatic sillimanite occurs in the Silver Plume-type granites (Anderson and Thomas, 1985). The Silver Plume-type and Oracle-type granites, in addition to those of New Mexico, form a distinct peraluminous province (Fig. l ) , bordered on the north by the metaluminous Sherman granite of Wyoming and on the southwest by numerous metaluminous plutons of the lower Colorado River region of western Arizona, Nevada, and California (Anderson and Bender, 1989). The southwestern petrologic boundary has been further characterized by sharp changes in Nd and Pb isotopic ratios (Bennett and DePaolo, 1987; Wooden and Miller, 1990). In general, these granites tend to have lower K20 (Fig. 2A) and ratios of FeO/(FeO + MgO) less than 0.84 (low Fe/Mg group, Fig. 2B). The Oracle-type granites have peraluminous compositions as a consequence of higher N 2 0 3 and lower K 2 0 (high K versus ultra-high K) a t any level of Si02; the plutons also have relatively high Mg and lower abundances of Ga and La (Anderson and Bender, 1989). Limited data for peraluminous granites of New Mexico are similar. In contrast, the peraluminous Silver Plume-type granites
Role of anorogenic granites in the Proterozoic crusfaldevelopment of NorthAmerica
273
are more potassic, yet maintain slightly higher ratios of N C N K (1.03-1.13) due to marginally higher A 1 2 0 3 and lower CaO, they also have lower Mg and higher Ba, Rb, and R E E at any corresponding level of silica. An unusual peraluminous intrusion is the 1.44 Ga Oak Creek batholith of Colorado (Cullers et al., 1992b). SiOz ranges from 56 to 73 wt.% and the intrusion is metaluminous only in the most mafic portions, the remainder being peraluminous at SiO2 greater than 58 wt.%. This high degree of alumina saturation at low silica is not found in any other Proterozoic pluton of the continent. The batholith also has the highest observed A 1 2 0 3 , which ranges up to 17.9 wt.%. Typically ultra-high K with an alkalic alkali-lime index, trends for Ti, Fe, and Ca as a function of silica are conspicuously low. Primitive members have relatively high Ba and Sr (up to and in excess of 400 pprn). Abundances of
Th, ZGand Nb
In their characterization of A-type granites, Whalen et al. (1987) discussed the enrichment of Zr and Nb. Likewise, Wilson and Akerblom (1982) have noted the high U concentrations of Proterozoic granites relative to those of other ages in Sweden. These elements, along with Th, deserve special mention due to rather striking abundance variations. The data reported below come from Silver et al. (1981), Doe et al. (1983), Anderson and Bender (1989), M.E. Bickford (written commun., 1990), Cullers et al. (1992a, b), and this study. The average A-type granite concentrations of Zr and Nb are reportedly 528 and 37 ppm, respectively (Whalen et al., 1987). Concentrations of Z r in these Proterozoic granites show a strong negative correlation with silica and thus are lower (< 250 ppm) in rocks with greater than 70% Si02. Yet independent of the silica dependence, the lowest abundances (< 200 ppm) occur in the peraluminous Silver Plume- and Oracle-type granites. The highest Zr (up to 1400 ppm) occurs in the low silica portions of the Wolf River, Oak Creek, and San Isabel batholiths. Granites of the southern midcontinent (including the St. Francois complex) have the lowest abundances of Nb (< 26 ppm, shown as low Nb group in Fig. 2D). Most of the other granites have Nb ranging up to 45 ppm; greater than 50 ppm Nb exists for portions of the Wolf River batholith and the Marble granite (of California). The most enriched is the Dells granite, a highly evolved intrusion in western Arizona. Together with their high Y and Rb, these plutons characteristically have compositions plotting in the “within-plate” field as defined by Pearce et al. (1984) (Fig. 2D). Average U contents in granites are often considered to be in the range of 3 to 5 ppm whereas enriched granites have greater than 10 ppm (Wilson and Akerblom, 1982). Uranium exhibits some positive covariation with silica in the anorogenic granites but abundances are conspicuously low (1-5 ppm) in most of the granites of the lower Colorado River region and central Arizona. Notably high concentrations exist in the Silver Plume (to 8 pprn), Log Cabin (to 9 pprn), Dells (to 10 ppm), Graniteville (to 12 pprn), Marble (to 13 pprn), and Wolf River (to
JL.Anderson and . I Morrison
274 10 8 6 4
2 (0
$
30
0
0
5
10
15
20
5
10
15
20
c
1 0 1
0
;:
Th (PPm) 5 0 2 >67%
14 12 10
-
--
m, ,,., 0
50
100
,v,,m,, ,m,m,,m., , , 150
200
1
18
0
50
100
150
200
Fig. 3. Abundances of U and Th in Proterozoic anorogenic granites separated into rocks with less than and greater than 67 wt.% Si02. Sources of data include Anderson and Bender (1989), Cullers et al. (1992a, b) and J.L. Anderson (unpublished).
16 ppm) granites. Ram0 (1991) reports U concenfrations up to 19.6 ppm in the Wiborg rapakivi massif but abundances in most of the other anorogenic granites of Finland range from 1.3 to 7.4 ppm. The complete distribution of U is shown in Fig. 3, exclusive of some exceptionally uraniferous examples, the Lawler Peak and Dells granites of Arizona (19.8 and 40 ppm U, respectively) (Silver et al., 1981) and granite sampled by drillhole in Illinois (16 to 64 ppm U) (Doe et al., 1983). Thorium abundances in A-type granites average near 20 ppm and vary widely (Whalen et al., 1987). The range in many of the granites considered in this study
Role of anorogenic granites in the Proterozoic crustal development of North America
275
is from 10 to 40 ppm (Fig. 3) but higher concentrations were determined for the Silver Plume batholith (40 to 100 pprn), subsurface granite retrieved from northern Illinois (90-140 pprn), the Oak Creek batholith (90 to 180 pprn), and the Marble granite (140-180 pprn).
INITIAL Sr AND Nd ISOTOPES
An extensive data set exists for the initial 87Sr/s6Sr(see data and references in Tible 2) of the anorogenic granites. The mean initial s7Sr/s6Sr ratio is 0.7056 f 0.0031 for plutons of the western U.S., which is similar to the 0.7051 f 0.0025 average value for all plutons of this age in the continent. No systematic difference occurs between metaluminous and peraluminous granites nor is there any regional variation. The Sherman and Sandia granites and the Pikes Peak batholith exhibit a range of initial 87Sr/86Sr.The Pikes Peak batholith is well studied, and a positive correlation exists between silica and initial s7Sr/86Sr(Barker et al., 1976). Lower ratios exist for gabbro (0.7044) and higher ratios have been determined for syenite and quartz syenite (0.7052-7063) and granite (0.7067-0.7117). Overall, the data imply a source that is not very old and/or radiogenic, yet the above noted ranges indicate that some degree of crust interaction may have occurred during magmatic evolution. Nd isotopic data (Table 2) come from DePaolo and Wasserburg (1976), DePaolo (1981), Farmer and DePaolo (1984), Nelson and DePaolo (1985) and Bennett and DePaolo (1987). As expected for felsic igneous rocks, most of the initial 143Nd/ 144Ndratios are lower than chondritic values at 1.4 to 1.5 Ga leading to a predominance of negative 6Nd values between -0.2 to -5.3. A few have higher ENd values, ranging +0.8 to +4.8. In contrast, Nd model crust-formation ages are uniform with a restricted range (from depleted mantle; DePaolo, 1981) (TDM) for specific regions of the continent (Nelson and DePaolo, 1985; Bennett and DePaolo, 1987). Model ages for granites of the central midcontinent as far west as central Arizona are mostly in the range of 1.7 to 1.9 Ga. Higher TDMages (1.8 to 2.1 Ga) occur for 1.4 Ga plutons in the Lake Superior region, the Sherman granite of Wyoming, and numerous plutons in Colorado, western Arizona, Nevada, and southern California. Based primarily on TDMdata for Proterozoic to Tertiary plutons, Bennett and DePaolo (1987) delineated three provinces in the western U.S. having model crust-formation ages of 1.7-1.8 Ga (central Arizona and New Mexico), 1.8-2.0 Ga (northern Arizona to Colorado), and 2.0-2.3 Ga (western Arizona and California). The TDMages for the 1.4-1.5 Ga granites largely support this conclusion, except that the ages for those of western Arizona and California (1.83 to 2.08 Ga) are somewhat lower than the designated 2.0 to 2.3 Ga model age of that province. Collectively, the Nd and Sr data, depicted in Fig. 4, have a range not unlike that observed for coeval anorthosite. This is not taken to imply that the granites are derived from the same magma systems that form anorthosite. It
JL.Anderson and Morrison
276 TABLE 2 Sr and Nd isotopic composition of midProterozoic plutons Locality
"~r/%ri
Wisconsin
Wolf River
0.7045
Michigan Illinois
TDM -5.3
2.30
DH: 81-12
-2.8
2.27
DH: UPH-3
-1.8
1.94
Nebraska
DH: LEM-1
0.7022
+1.2
1.77
Kansas
DH: DG-3 DIH: MI-4
0.7017
-1.8 -1.6
1.78 1.90
Missouri
DH: L6-4 DH: L5-2 Butler Hill
0.7035 0.7029
f2.6 -0.9 +4.8
1.63 1.97 1.47
Oklahoma
DH: BR DH: TG
0.7061
+0.8 -1.2
1.81 1.98
Texas
Town Mountain
0.7060
f2.3
1.34
Wyoming
Sherman
0.7036-0.7065
-3.4
2.11
Colorado
Log Cabin St. Vrain Silver Plume Vernal Mesa Curecanti Pikes Peak Eolus San Isabel
0.7031 0.7025 0.7030 0.7038 0.700 0.7067-0.7117 0.7043 0.7030
-1.5 -2.3
1.94 1.81
-1.2
1.51
Dos Cabezas Rana Sandia Priest Pedernal Los Pinos Ladron Oscura
0.7051 0.7113 0.7027-0.7060 0.7029 0.7027 0.7078 0.7107 0.7060
+1.3
1.72
Stockton Pass Oracle Ruin Dells Lawler Peak Sierra Estrella Holy Moses Hualapai
0.7101 0.7065 0.7065
+1.0
1.80
New Mexico
Arizona
+2.8 -0.2 0.7050 0.7094 0.7032
-2.78
1.86
Nevada
Gold Butte
-1.13
1.83
California
Newberry Davis Dam Homer Marble Parker Dam
-2.23 -2.72 -3.97 -4.50 -2.12
2.08 1.99 1.89 2.01 1.97
0.7042
Role of anorogenic granites in the Proterozoic crustal development of North America
277
does require, however, that the granite source not be very radiogenic or old as implied by the range of TDMages. The most consistent interpretation involves a predominately Proterozoic crustal source, a conclusion also supported by Hf isotopic data (Patchett et al., 1981). Combined isotopic and elemental data for those granites with 1.7 to 1.9 Ga T D model ~ ages are best explained by derivation from Early Proterozoic crust produced during preceding orogenic growth of the craton. Bennett and DePaolo (1987) account for the model ages in excess of 2.0 Ga resulting from limited (10 to 30%) contribution of Archean crustal material. Two of the high TDMgranites (Wolf River batholith with TDMat 2.3 Ga and the Sherman granite at 2.1 Ga) partially intrude Archean crust. Ram0 (1991) reports similar high TDMages for two 1.54 Ga rapakivi intrusions (lowest open square data points in Fig. 4b), the Salmi (TDM= 2.46 to 2.82 Ga) and the Sotjarvi (2.32 Ga), both of which were emplaced marginally into Archean rocks. In contrast, the 1.54 to 1.65 Ga Finnish rapakivi granites yield distinctly lower TDMages, ranging 2.0 to 2.2 Ga (Ramo, 1991). Low Nd model ages ( 5 1.6 Ga) are limited to a four examples (Fig. 4b) but remain provocative. Included are data for a 1.47 Ga diorite (TDM = 1.6 Ga) and a 1.49 Ga granite (Butler Hill granite, TDM= 1.47 Ga) from Missouri. The remaining low model ages are from two 1.0- to 1.1-Ga plutons, the Pikes Peak batholith (TDM = 1.51 Ga) of Colorado and the Town Mountain granite of Texas (TDM = 1.37 Ga). Together the data may imply the existence of a previously unrecognised 1.4 to 1.6 Ga, mantle-derived crustal terrane that was imaged by these partial melting events (Nelson and DePaolo, 1985). However, this conclusion is flawed for two reasons. First, the crust formation ages assume ultimate derivation from a strongly depleted mantle. If the ultimate mantle source were less depleted or undepleted (chondritic), then the above model ages are erroneously old. Secondly, melting invariably fractionates Nd and Sm precluding crustal-derived magmas from giving accurate mantle-separation ages. Residual pyroxene, garnet, hornblende, and zircon retain Sm over Nd leading to apparent crust-formation ages that are erroneously young. In contrast, residual feldspar and
NOTES TO TABLE 2 4 143 144 Ndi-,,,,,,)/(143Nd/144NdT_CHUR)I -1, a numerical expression of the initial f N d ( T ) = 10 [( Nd/ Nd isotopic composition of the sample of age T relative to a chondritic reservoir. a
TDM = age of separation in Ga from depleted mantle (DePaolo, 1981) assuming single stage history. DH = drill hole sample followed by sample number.
References for Sr data include: Steiger and Wasserburg (1966); Muehlberger et al. (1966); Peterman et al. (1967, 1968); Hansen and Peterman (1968); Bickford et al. (1969); Mose and Bickford (1969); Bickford and Cudzilo (1975); Subbarayudu et al. (1975); Mukhopadhyay et al. (1975); Swan (1976); Kessler (1976); Barker et al. (1976); White (1978); Register and Brookins (1979); Brookins et al. (1980); Keith et al. (1980); Erickson (1981); Davis et al. (1982). References for Nd data include: DePaolo (1981); Farmer and DePaolo (1984); Nelson and DePaolo (1985); Bennett and DePaolo (1987).
. L. l Anderson and J. Morrison
278 6
A
2
+
A
v
T
Z
o
W -2
-4
-6
-8 0.700
0.702
0.706
0.704
0.710
0.708
0.712
0.714
87sr/'sSr Nd Evolution - Rapakivi Granites & Anorthosites Granite o North America Finland Mafic Rocks HarpLake A Mealy Mtn Harpdikes x Marcy rn Fmland Other 1.2
1.4
1.6
1.8
2.0
Age (Ga)
Fig. 4. (a) €Nd and s7Sr/86Srdata for Proterozoic anorogenic granites compared to fields for anorthosites (taken from Ashwall and Wooden, 1985). Symbols according to U-Pb age: squares = 1.4 to 1.5 Ga, circles = 1.3 to 1.4 Ga, and triangles = 1.0 to 1.1 Ga. Sources of granite data given in Table 2. (b) 6Nd versus age for Proterozoic rapakivi granites and temporally related mafic rocks including anorthosite, leuconorite, gabbro, and diabase. Sources for North American granite data given in Table 2. Data for Finland granite and anorthosite from Ramo (1991); Marcy, Mealy Mountain, Harp Lake, and Harp Dikes data from Ashwal and Wooden (1983) and Ashwal et al. (1986). Other mafic rock data includes norites from Greenland (Patchett and Bridgwater, 1984) and the Laramie anorthosite (Kolker et al., 1990).
biotite have the opposite effect. Nelson and DePaolo (1985) calculate that the offset due to fractionation should be less than 100 Ma but their analysis is not well constrained. If monazite or allanite were retained in the residue, then the fractionation of Sm to Nd would be even more profound. Alternative explanations for the low ages exist but are also non-unique. The Pikes Peak batholith is hosted in Early Proterozoic rocks and the Nd data are also
Role of anorogenic granites in the Proterozoic crustal development of North America
279
permissive with the view of Barker et al. (1975) that the Pikes Peak formed from mixing of cogenetic mantle-derived and crust-derived melts. The ?bwn Mountain granite intrudes Grenville-aged (1.1-1.2 Ga) crust and the Nd isotopic data could also be explained by partial melting of older crust mixed with material of the host terrane.
OXYGEN ISOTOPIC COMPOSITIONS
Oxygen isotope compositions may yield important information concerning both source characteristics and later contamination, particularly when interpreted in conjunction with radiogenic isotopic compositions. Interestingly, a review of the literature reveals a scarcity of oxygen isotope data from anorogenic granites. Although a few selected intrusions have been studied in detail (Wenner and Taylor, 1976; Shieh e t al., 1976; Heaman e t al., 1982; Shieh, 1983; Wu and Kerrich, 1986), no systematic oxygen isotopic survey of anorogenic granites has been conducted on a regional basis. We have recently begun such a survey and our initial results are reported in Table 3. The whole rock data are characterized by a large range in S ” 0 , from 6.2 to 11.5%0. However, because many anorogenic granites have epizonal features, it is crucial to assess whether post-crystallization hydrothermal alteration has altered magmatic 6l’O values. For example, Wenner and Taylor (1976) demonstrated that regionally extensive hydrothermal alteration in the St. Francois Mountains of Missouri has lead to elevated 6l’O values of the Bevos and Musco Group granites. Only a few samples retained apparently N 1 to 1.5 (Astz-fsp = “igneous” 6”O values as indicated by Aquartz--feldspar 6l8OqtZ- 6l8OfSp).In samples interpreted to have been hydrothermally altered, Astz-fsp values ranged from < +1 to -1.8. Thus, in the absence of S1’0 values for coexisting quartz and feldspar, caution must be exercised in interpreting whole rock S ” 0 values in Table 3. Petrographic examination of samples listed in Bible 3 has enabled identification of samples which contain mineralogic evidence of hydrothermal alteration. All whole rock 6lSO values are plotted versus wt.% Si02 in Fig. 5a. Three samples which have undergone hydrothermal alteration (indicated by near complete replacement of biotite by chlorite) are marked with a “c” (Fig. 5a) and are excluded from the compilation in Fig. 5b. There is additional evidence of hydrothermal alteration of samples plotted in Fig. 5a. Values for samples from individual plutons have been circled, and such “within-pluton” S*’O variations suggest that at least 3 plutons within this data set may have experienced a hydrothermal readjustment of their oxygen isotope ratios. For example, the hornblende-biotite Hualapai granite of Arizona has S ” 0 values of 9.2, 6.1, and 9.0 at Si02 contents of 66.7, 68.4, and 71.0 wt.%, respectively. This 3%0 range in 6”O is relatively large, and may not be a primary magmatic feature. Whole rock Sl’O values for anorogenic granites from localities in North America are plotted in Fig. 5 with the data reported in B b l e 3. Lithologies represented include granite, quartz monzonite, monzonite, monzosyenite and
J:L. Anderson and J. Mowison
280 TABLE 3
Oxygen isotopic composition of mid-Proterozoic plutons - western U.S. Locality 6180wr Si02 (wt.%) Colorado
St. Vrain
Silver Plume
Arizona
Oracle Ruin Ak-Chin Sierra Estrella Holy Moses Hualapai
Continental Ft. Huachuca Nevada
Gold Butte Beer Bottle Pass
California
Newberry Davis Dam Marble
Parker Dam
Bowmans Wash
11.6 10.8 10.6 6.8 * 10.4 10.9
72.18 71.24 69.49 72.49 70.55 67.47
10.5 11.0 10.3 10.3 11.2 10.2 7.6 8.4 6.1 * 9.1 9.1 7.7 7.9 11.1 12.0
67.30 67.89 71.27 70.27 73.43 70.15 67.05 68.28 68.44 66.67 71.00 64.43 64.24 67.55 63.52
9.1 9.2 8.1 9.6
66.06 66.76 65.99 67.77
10.8 8.9 * 10.1 11.4 10.4 8.8 9.7
65.20 64.03 70.36 66.49 67.84 68.35 69.50 66.33 67.42 69.41 69.88 61.32 60.85
10.0 8.4 9.1 8.9 6.5 * 8.2
* Chloritized sample.
trondhjemite. Also shown are ranges for (<6%0), “normal” (6 to 10%0) and “high” (>10%0) S1’0 granites as defined by Taylor (1978). These three ranges are arbitrary (Taylor, 1978), but they provide a working framework for interpreting S”O variations in granites.
Role of anorogenic granites in the Proterozoic crustal development of North America
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1.4 Ga Granites, Western US I
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Fig. 5. (a) Whole rock S ” 0 values for anorogenic granites of the western and southwestern U.S., including the St. Vrain and Silver Plume batholiths of Colorado; Ruin, Oracle, Sierra Estrella, Hualapai, and Holy Moses granites of Arizona; Beer Bottle Pass and Newberry granites of Nevada; and the Davis Dam, Homer, Parker Dam, Bowrnans Wash, and Marble granites of California. Where multiple analyses exist for a pluton, the data are circled to indicate the observed range in composition (see text). (b) Histogram of whole rock S1’0 values for anorogenic granites in N. America. Ranges for “low”, “normal” and “high” 6”O granites (Taylor, 1978) are also shown. Data from the following localities are plotted: (1) the Loon Lake monzonite and quartz-monzonite of Ontario (Shieh et al., 1976; Heaman et al., 1982), (2) the Algonquin granite and monzosyenite of Ontario (Wu and Kerrich, 1986), (3) the Illinois Deep Hole granite (Shieh, 1983), (4) the Pikes Peak batholith of Colorado (Barker et al., 1976), (5) the Wolf River batholith of northern Wisconsin (Kim, 1989), and (6) new data for plutons of the western and southwestern U.S. (given i n detail in (a)).
Oxygen isotopic compositions are most useful’ when S1$0 values fall outside the “normal” range. Values of S’$O < +6%0 are likely to result from either (1) interaction between a normal-180 magma and low 6 l 8 0 country rocks (e.g. previously hydrothermally altered by low-l80 meteoric fluids) or (2) melting of hydrothermally altered 10w-’~O rocks. Granites with S ” 0 values > +10 are probably derived from (1) melting of sedimentary rocks or (2) exchange between a normal-180 magma and sediments via either magmatic assimilation or exchange with a fluid derived from sediments.
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Examples of high 6 l 8 0 plutons include the two-mica Oracle-type and Silver Plume-type granites of the western U.S. Their average 6 l 8 0 (10.7 f 0.4) is only marginally higher than that observed for biotite granite (10.4 f 1.2), but significantly higher than that observed for biotite-hornblende granite (9.1 f.1.0). In keeping with their silica and peraluminous composition, the 6 l 8 0 is consistent with derivation from a source containing metasedimentary material. In contrast, the Loon Lake quartz monzonite of Ontario, the only “high” S l 8 0 granite, has S ” 0 values between 8.9 and 13.9 (mean = 11.3). Shieh et al. (1976) and Heaman et al. (1982) attribute the elevated and highly variable S l 8 0 values of the quartz monzonite to interaction with a mixed H20-C02 fluid that was derived from the country rocks. The Loon Lake monzonite forms the core of the complex and has S l 8 0 values that range from 8.8 to 9.7 (mean = 9.4). These values, in conjunction with a high WRb and an initial 87Sr/86Srvalues of 0.7036 f0.0006, are interpreted to indicate a lower crustal source for the Loon Lake complex. The three “low” S ” 0 values are from two different localities. One of the low values (4.3%0) is from the Algonquin granite, also from Ontario (Wu and Kerrich, 1986). The observed range in the oxygen isotopic composition of the potassic suite of the Algonquin granite (4.3 to 9.3) is interpreted to result from isotopic exchange of 10w-’~Owaters during emplacement (Wu and Kerrich, 1986), despite the preservation of magmatic Aqtz-fspvalues. The two other low values (5.7, 5.7) are from the Wolf River batholith (Kim, 1989). The oxygen and sulfur isotopic compositions indicate that the Wolf River granites were derived from partial melting of a 1ow-l80, lower crustal source region. The magmas probably underwent variable degrees of interactions with Penokean plutonic rocks. The two low values from the Wolf River batholith are likely the most pristine and thus rocks. lowest PO The majority of anorogenic granites (81%) fall within the “normal” range. In the absence of radiogenic isotope data, there is some ambiguity involved in interpreting these values. The ambiguity arises because multiple scenarios can account for the 6l80 values. For example, variable mixing between mantle-derived basalts and gabbros (6l80 M +6) and crustal rocks (e.g. 6lSO x 8-10) could explain any of the normal values. Alternatively, melting of meta-igneous crustal rocks which can range in 6l80 from 6 to 10%0 could also result in “normal” granites.
CONDITIONS OF CRYSTALLIZATION
Temperature Utilizing two-feldspar, garnet-biotite, and Fe-Ti oxide thermometry, Anderson (1980), Anderson and Thomas (1985), Anderson and Bender (1989), and Anderson (unpublished data) have determined crystallization temperatures of granitic plutons from Wisconsin to California. Results for individual plutons include: Wolf
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River (640-790"C), Silver Plume (740-760"C), San Isabel (745-79OoC), Oak Creek (63O-67O0C), Ruin (701-703"C), Sierra Estrella (622-647"C), Fort Huchuca (649725"C), Parker Dam (616-66loC), Marble (616-759"C), Hualapai (732-794"C), and Gold Butte (615632°C). Thermometry for pyroxene-bearing charnockitic rocks yield higher crystallization temperatures. For monzosyenites of the Laramie complex, Fuhrman et al. (1988) have determined a crystallization temperature of 950-1050°C. Likewise, Anderson (1980) calculated a 860-970°C range for mangerites of the Wolf River batholith.
Depth of emplacement Many of the plutons have epizonal features but may have not have intruded to near-surface levels that are typical of the midcontinental granites. Pressure estimates from country rocks within contact aureoles potentially provide quantitative constraints on emplacement depth. Unfortunately, the general case is that aureole development around these granite intrusions has either been found to be lacking or has not been recognised. Clearly, this enigma of "apparent" lack of contact metamorphism is a viable subject for more study. More conclusive information has been derived from crystallization barometry. This approach is appropriate for granites of peraluminous composition, containing garnet and two micas, and those of metaluminous composition, containing hornblende, biotite, sphene, and magnetite (reviewed by Anderson, 1987b, and Zen, 1989). With increasing pressure, magmatic muscovite becomes more siliceous, garnet more calcic (via exchange with plagioclase), and hornblende more aluminous. In addition, magmatic epidote becomes stable in metaluminous granites a t pressures > 4 to 6 kbar (Zen and Hammarstrom, 1984). Evolved portions of the Oak Creek batholith of Colorado contain garnet and two micas. Anderson (in Cullers et a]., 1992b) calculated the emplacement pressure to be 2.7 to 3.9 kbar based on several calibrations for the pressure-sensitive element exchange among garnet, biotite, plagioclase, and muscovite. Many of plutons contain near-solidus hornblende, allowing an estimation of crystallization pressure with the Al-hornblende barometer (calibrations of Hollister et al., 1987 and Johnson and Rutherford, 1989). Unfortunately, amphiboles in most of the granites are Fe-rich. The barometer was based on hornblende with an Fe/(Fe + Mg) ratio between 0.42 and 0.58; high-Fe amphiboles typically have high-alumina contents potentially leading to erroneously high pressure estimates (Anderson and Bender, 1989). Only hornblende of the Hualapai (western Arizona), San Isabel (southern Colorado), and Silvermine (Missouri) granites fall within this designated Fe/(Fe + Mg) range. For the Hualapai granite, the calculated pressure is 3.5 f 0.6 kbar, not unlike the 1.7 Ga metamorphic conditions for the lower Colorado River region (Thomas e t al., 1988; Young e t al., 1989). For the San Isabel granite, the calculated pressure is 6.2 f 0.8 kbar, consistent with the occurrence of magmatic epidote in this pluton. In contrast, hornblende from the Silvermine granite of the
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St. Francois Mountains (Anderson, unpu51ished) yields an anticipated low pressure of 1.1f 0.5 kbar. Plutons of the St. Francois complex intrude and are overlain by broadly coeval silicic' volcanic rocks. Based on the composition of iron-rich olivine and pyroxene (with quartz) in syenites of the Laramie complex, Kolker and Lindsley (1989) estimate emplacement of this composite anorthosite-charnockite intrusion to be 4.0 to 4.5 kbar. Using average crustal densities, these results show that for many 1.3 to 1.5 Ga plutons, the emplacement depth was on the order of 10 to 16 km. The San Isabel batholith is an exception. At a calculated depth of 23 f 3 km, this intrusion is the only mid-crustal pluton known for the continent-wide expression of this enormous magmatic event. This may relate to its unusual composition and the fact that it is the only reported batholith of its age to contain magmatic epidote.
Water and oxygen jiqpcity Water contents of anorogenic granitic magmas are generally viewed as low (Anderson, 1983) due to the paucity of pegmatites and/or hydrothermal activity and the late crystallization of the hydrous phases. Exceptions to this generalization include some of the two-mica granites of New Mexico (Wobus, 1984), Arizona, and Colorado. Other peraluminous granites are relatively dry, including the Silver Plume granite (Colorado) for which Anderson and Thomas (1985) calculate a fHzo of 433 to 495 bar ( P H of~ 487 ~ to 560 bar) relative to an uncertain total pressure of 3 to 4 kbar. In these rocks, muscovite and biotite crystallized late and occur interstitial to, or poikilitically enclosing, other mineral phases. The metaluminous granites, though having sufficient water and/or fluorine to stabilize biotite and hornblende (rather than pyroxene), can be consistently shown to be relatively dry intrusions. For example, the P H of~the~ Isabel batholith has been determined to be on the order of 0.5 to 2 kbar relative to a total pressure in excess of 6 kbar (Anderson, in Cullers et al., 1992a) The most widely varying intensive parameter is oxygen fugacity Cfoz),which ranges over four orders of magnitude for the granitic plutons. Effects of this variation occur in the Fe-Ti oxide mineralogy and in the Fe/(Fe Mg) ratio of the mafic silicates. Low fo, granites have ilmenite as the dominant or sole Fe-Ti oxide and iron-rich mafic silicate minerals. Examples are uncommon among the Middle Proterozoic complexes of the North American continent but include the Pikes Peak batholith (Colorado) and the Wolf River batholith (Wisconsin). Both contain ilmenite f magnetite and iron-rich biotite and amphibole with Fe/(Fe Mg) ratios greater than 0.80 (Barker et al., 1975; Anderson, 1980; Anderson, 1991). Calculatedfo, is on the order of bar a t a average temperature of 720°C. Figure 6 depicts the relative fo, of 1.4-Ga plutons of the western U.S., compared to that determined for the Wolf River batholith. Higher fo2 conditions result in the presence of magnetite as the dominant Fe-Ti oxide (Ishihara, 1977) and lower Fe/(Fe + Mg) ratios of the mafic silicates, a standard characterization that applies to all of these Middle Proterozoic plutons. At an average temperature
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Role of anorogenic granites in the Proterozoic crustal development of North America
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Fig. 6 . Cxystallization conditions in terms of temperature andfo2 for 1.4 Ga granites (Anderson, 1983; Anderson and Bender, 1989) compared to the low fez, ilmenite-series granites of the Wolf River batholith (Anderson, 1980).
of 720"C, the Silver Plume granite crystallized at an average fo, of bars, the metaluminous granites of California and Nevada at u - ' ' . ~ bars, the two-mica, Oracle-type granites of Arizona at bars, and the Hualapai granite of western Arizona at 10-13.2 bars (Anderson and Thomas, 198.5; Anderson and Bender, 1989). The Hualapai granite magma crystallized at the most oxidizing conditions calculated for any ca. 1.4 Ga granite in the continent and it has the most Mg-rich mafic silicates (in spite of the high Fe/Mg of the rock) including biotite with a Fe/(Fe Mg) as low as 0.27. The most reduced fo2 conditions occur in the Laramie anorthosite-syenite complex which Fuhrman et al. (1988) and Frost et al. (1988) place at 1.5 to 2.0 log units below the quartz-fayalite-magnetite buffer.
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Assessment of liquidus temperature based on zircon saturation Experiments by Watson and Harrison (1983) show that the solubility of Zr in a silicate melt decreases with (on an atomic basis) increasing Si/(K Na 2Ca) and Al/(K Na + 2Ca) and provide a basis for calculating the saturation temperature of zircon. Zircon is a near-liquidus phase in many anorogenic granites, thus estimation of its temperature of initial crystallization enables not only assessment of liquidus conditions but also minimum temperatures achieved during magma formation. Figure 7a shows calculated zircon saturation temperatures, plotted against ppm Zr, for several of the anorogenic suites including the Wolf River, Oak Creek, and San Isabel batholiths, the St. Francois complex, the peraluminous granites of central Arizona (labelled as Arizona), and the metaluminous granites of southern California, Nevada, and western Arizona (labelled as California). At a fixed Zr concentration, more felsic rocks could have higher temperatures of saturation, but as shown in Fig. 7a, the temperatures trend to lower values because
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286 Zircon Saturation 1 0 0 0 1 . .
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San Isabel Wolf River -WR-cumulate California St. Francois Arizona Oak Creek OC-cumulate Silver Plume
Fig. 7. Zircon saturation temperatures (Watson and Harrison, 1983) plotted against concentrations of (a) Zr and (b) Ba. Sources of data include Cullers et al. (1992a, b), M.E. Bickford (unpublished), Anderson and Cullers (1978), Anderson and Bender (1989), and Anderson (unpublished).
Z r abundances generally decrease with increasing Si02. The lower temperatures approach what has been determined independently for solidus conditions by the element-exchange thermometry described above. The focus here is on results for the more intermediate, higher Zr compositions in which zircon crystallization is early, and the calculated temperatures (>85OoC) more indicative of liquidus conditions. The assumption is that no restitic, inherited, or cumulate zircon occurs. U-Pb geochronology (references in B b l e 1) does not indicate the existence of restitic or inherited zircon, but several geochemical studies have shown that the more primitive portions of some intrusions formed by accumulation of solids, including zircon. For example, the Wolf River granite of the Wolf River batholith has an observed range of SiO2 from 61.3 to 75.2%, yet the lower silica portions (61.3 to 69.9% SiOz), have formed by up to 30% accumulation of two feldspars, biotite, hornblende, and zircon (Anderson and Cullers, 1978). This process has resulted in positive Eu anomalies and enrichment in Ba, Sr, and Zr. Similar
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findings have been made for the Oak Creek batholith (Cullers et al., 1992b). The general case for these anorogenic granites is that any rock with greater than 250 ppm Sr, 1400 pprn Ba, or 700 pprn Z r may represent partial accumulation and thus, the derived temperature of zircon saturation may be erroneously high. Figure 7b plots calculated temperature versus Ba. O n both Figs. 7a and 7b, cumulate portions of the Wolf River (labelled WR-cumulate) and Oak Creek (OC-cumulate) are designated. Omitting results reflective of accumulation, it is apparent that estimated liquidus temperatures are very high, typically in excess of 850°C and up to 910°C. The implication is that these granites represent very hot magmas and that their mode of formation was at temperatures far in excess of their solidus and greater than that achievable by wet melting of average crust.
ANOROGENIC OR EPIZONAL?
The term anorogenic implies the lack of any relation to orogeny and thus has all the failings of any genetic classification. It should be used with caution where knowledge of the complete deformation history of the crust is precluded by lack of deeper exposure depths. Even in orogenic settings, low-H20 magmas will have the capacity to rise far from zones of crustal disturbance and, depending on the nature of magma generation and source composition, may have the standard geochemical and physical attributes of an A-type magma series. The mid-Proterozoic igneous complexes of North America formed on an immense scale involving a transect across the entire continent with significant variations in exposed crustal depth. With exceptions of portions of the Grenville Province (Daly and McLelland, 1990), nowhere has Proterozoic orogenic deformation, metamorphism, or calc-alkaline magmatic activity been found to have an age between 1.3 and 1.5 Ga. These facts, plus the data offered above showing that these igneous suites in general have typical A-type elemental abundances, confirms the conclusion that they are anorogenic and unrelated to any plate margin activity. Alternatively, Nelson and DePaolo (1985) determined that the 1.1Ga rapakivi granites of the Llano uplift of central Texas have a Nd-depleted model age of 1.37 Ga and concluded that these and other granites formed as an “inland manifestation of subduction activity”. The evidence, based on Sm/Nd systematics, is indirect and somewhat unconstrained as explained earlier. The 1.37-Ga model age assumes derivation from a depleted mantle or from youthful crust derived from a depleted mantle where no Sm/Nd fractionation occurred during partial melting. If the ultimate mantle source was instead only mildly depleted (near the observed ENd of +2.3, .instead of +5.0), then the calculated crust formation age could be equivalent to the young (Grenvillian) crust hosting the Llano granites. Additionally, the age could be the result of a heterogeneous crustal source composed of both Grenvillian and some older component, one that was, for example, of Early Proterozoic age (Patchett and Ruiz, 1987).
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SOURCE OF GRANITIC MELTS
Twenty-seven determinations of ENd (see Table 1 and references therein) range from +4.8 to -5.3. Nelson and DePaolo (1985) concluded that while a majority of the batholiths investigated in their study were derived “almost purely from the crust”, those with higher ENd (e.g., Butler Hill and Town Mountain granites) are indicative of magma derived from the mantle. Nelson and DePaolo further argue that the range of 6Nd implies mixing of crustal and mantle components. Alternatively, we suggest that all of these granites are crustal derived and that the range of ENd results from large variations in Sm/Nd in the crustal source, variations in age of the crust at the time of melting, and, as previously described, incorporation of small components of much older (e.g., Archean) material. The experimental studies of Wyllie (1984, 1988) have demonstrated that large volumes of granite cannot be derived directly from the mantle by partial melting. The high 6Nd (>2.0) can be explained by melting youthful, mantle-derived crust with a high Sm/Nd ratio. Based on major and trace element compositions, Condie (1978, 1991), Anderson and Cullers (1978), and Cullers et al. (1981, 1991) have presented mathematical models showing that many of the elemental features of these granites are consistent with melting of the lower crust, one of granulitic metamorphic grade comprised of tonalitic to granodioritic composition. The inferred granulite grade would have been achieved at the time of melting, if not before. A critical issue, recently reviewed by Creaser et al. (1991), centers on the melting origin of A-type granites of which these Proterozoic granites are typical examples. Collins et al. (1982) argued that the low HzO, high F, and high HFSE abundances of A-type granites resulted from partial fusion of melt-depleted granulite. Both Anderson and Bender (1989) and Creaser et al. (1991) opposed that model for several reasons. A melt-depleted granulite would be depleted in H20 and, if residual amphibole and biotite occur, enriched in E Yet, melt extraction would also eliminate alkali feldspar and deplete or eliminate quartz and biotite, all of which are required source phases to account for the typical anatectic or minimum melt composition of many of the granites and serve, by calculation, to explain abundances of K, Rb, and Ba (Anderson, 1983). Moreover, melt depletion would decrease the Fe/Mg ratio whereas these granites are characterized by high to very high Fe/Mg. Similar arguments can be made on the basis of Ca (enriched by melt extraction, these granites are not calcic) and R E E (depleted, whereas the observed R E E are high). The nature of partial melting is perhaps the key factor separating the differing origins of I-type (Chappell and White, 1974) and A-type granites, since both can arise from melting meta-igneous crust. The broader availability of water during orogenesis clearly would promote higher degrees of partial melting which, under wetter conditions, would also include derivation from crustal sources more mafic than tonalite. The role of mantle-derived magmas amongst I-type suites and mixing of mantle- and crustal-derived magmas is often quite convincing. The
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MASH (melting-assimilation-storage-homogenization) of Hildreth and Moorbath (1988) may be very important in the formation of I-type magmas but appears to have no role in the origin of A-type magmas. Under anorogenic conditions, there is no obvious source of externally derived fluids and the amount of melting depends critically on the abundance and composition of hydrous phases, particularly biotite. A source most consistent with available data is meta-igneous tonalitic to granodioritic crust. The low water content of the granites is a direct result of limited water in the source (probably < l%),bound in biotite and hornblende until breakdown temperatures were exceeded and melting occurred. All of the melting is envisioned as vapor-absent and the percentage of melting controlled by the ability of the hydrous phases to participate in melt-generating reactions. Unlike melting under wet conditions, the A T between the solidus and liquidus of a tonalite is large, on the order of 250°C or more. Thus, the amount of melting would be limited. Based on reasonable amounts of hydrous phases in a tonalite or granodiorite, Creaser et al. (1991) calculated the amount of melting to be 15 to 45% with H 2 0 in the derived melt less than 2.6%. Similar amounts of melting have been experimentally determined by Wyllie (1988) and modelled, based on elemental abundances, by Anderson and Cullers (1978) and Cullers et al. (1981). The existing fo2 evidently also plays a major role with the low fez, ilmenite-series anorogenic granites being the most enriched in K, other LIL elements, and Fe/Mg. The forward progress of melt-forming, biotite consuming reactions would be limited by the ambient oxygen level in the source; under low fez, the biotite breakdown and the amount of melting would be more limited relative to that produced by the same source under higher fo2. In conclusion, we suggest that the high K, Rb, Ba, REE, HFSE, F, and Fe/Mg of unevolved members of most anorogenic granites is a direct result of the low degree of melting of intermediate meta-igneous crust produced during a prior Proterozoic orogeny. Less silicic intrusions, such as the San Isabel batholith of Colorado or the Bowmans Wash quartz monzodiorite of California requires a less felsic or dioritic source. Some metasedimentary component presumably occurred in the source of the peraluminous granites of Colorado and Arizona. These plutons are often less MgO) and higher potassic, have less Nb, Y , and E lower ratios of FeO/(FeO Sl80.Several appear to have crystallized under wetter conditions. Reasonably, the metasedimentary material would have had a greater amount of hydrous phases yielding a greater degree of partial melting allowing some dilution of elemental abundances that otherwise are typical of A-type granites. The Sr and Nd isotopic data do not differ for these plutons, which constrains the amount and radiogenic character of the inferred metasedimentary contribution.
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ORIGIN OF ANOROGENIC MAGMATISM: RIFTS, PLUMES, AND SUPERSWELLS
The origin of the anorogenic magmatism that swept across the continent a t 1.4 to 1.5 Ga followed by repeated events a t 1.3 to 1.4 and 1.0 to 1.2 Ga, remains one of the more enigmatic plutonic and volcanic episodes of the Proterozoic Laurentian craton. Similar Proterozoic igneous activity has been documented on nearly every other continent between periods of orogenic growth. The amount of crustal readjustment is profound, and in many areas 15 to 40% of exposed Proterozoic crust are composed of these intrusions or their volcanic equivalents. Much higher percentages exist for the central midcontinent based on drill hole data (Bickford et al., 1986). The fact that the Proterozoic granitic igneous activity never extends far into Archean portions of the continent must be of fundamental importance. An expression of the event in the Archean shield is recorded, but only by emplacement of mafic magmas, including the impressive 1.27-Ga Mackenzie and the 1.24-Ga Sudbury dike swarms (Fahrig, 1987). Large-scale production of granitic magma in this region is clearly missing. Outside the Archean craton, anorthosite massifs are voluminous in the Grenville Province and Laborador but elsewhere are less abundant. Diabase dike swarms are widespread, however, and notable occurrences include dikes in Missouri, the Colorado Front Range, and throughout much of the Southwest. The swarms tend to have a consistent northerly trend indicative of a mild, but pervasive state of regional extension. Numerous workers have called attention to the compositional affinity of the anorogenic igneous activity to that of younger suites associated with extension and rifting (Barker et al., 1975; Anderson and Cullers, 1978; Emslie, 1978 and 1991; Bickford et al., 1986). Yet, with exception of the Keweenawan mid-continent gravity high (Behrendt e t al., 1988), nowhere is there evidence for extension maturing into well documented rifts. The abundant production of crust-derived granitic magmas points to a central problem. In absence of subduction-related mechanisms, continental crust intrinsically lacks heat for in-situ melting. As argued by Morse (1991), the heat required to melt large volumes of crust must be imported. An obvious source is the mantle, specifically, the same melts that yielded anorthosite and related mafic rocks. Anderson (1987a), Anderson and Bender (1989) and Hoffman (1989), have called upon a model involving mantle diapirism. Anderson (1987a) and Anderson and Bender (1989) coined the phrase “mantle and crustal overturn” and subsequently, Hoffman (1989) described “mantle superswells”, to refer to a similar model. After a 600-Ma period of orogenic quiescence, many continents underwent rapid orogenic growth from 1.9 to 1.65 Ga before entering another period of relative stability (Condie, 1990). The new crust is envisioned as fertile, not having experienced significant prior melting events, and thus would contain a potential low-melting fraction susceptible to an influx of mantle-derived heat or magma. Likewise, the underlying mantle, now isolated from ocean-crust-formation processes, is also less depleted relative to suboceanic mantle. The evidence for the
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mantle being less depleted can be seen in Fig. 4. Remarkably, out of the more than 40 initial Nd isotopic determinations plotted for anorthosite, gabbro, and diabase taken from North America to the Baltic, not one falls on the depleted mantle curve of DePaolo (1981). This is in contrast for the many depleted-mantle compositions observed for mafic rocks (gabbro to basalt) emplaced in the older orogenic suites (Patchett and Arndt, 1986; Huhma, 1986; Chauvel et al., 1987). That some of the low 6Nd may be the result of crustal contamination is inescapable (Ashwal et al., 1986) but, given the large volumes of some of these mafic to anorthositic intrusions, it seems that the much this subcontinental mantle was not depleted (and later contaminated by crust assimilation) and, in part, may have been chondritic to enriched. Chondritic mantle would have higher incompatible element abundances, which could lead to radiogenic heating and diapiric rise. Ascent of a gravitationally destabilized mantle plume could lead, depending on the P-T path of decompression, to the formation of large gabbroic magma chambers which, in the absence of rifting, would underplate o r intrude into the lower crust. As originally envision by Barker et al. (1975), a succession of melting events would result beginning first in the mantle (gabbro to anorthosite) and pass upward into the lower crust causing a profound thermal disturbance and formation and separation of potassic granitic magma. The process would continue until both the mantle and the relatively young Proterozoic crust achieved a more stable gravitational and thermal configuration. The eventual result is that the mantle becomes depleted in a thermally unstable component and the crust becomes differentiated into a mafic lower crust of residual composition and an upper felsic crust derived from intrusion and extrusion of granitic magma.
ACKNOWLEDGEMENTS
The authors acknowledge unpublished data shared by Bob Cullers and Pat Bickford, and the support of NSF grants EAR 91-9105636 and EAR 89-04060. Critical reviews by Kent Condie, Bob Cullers, and Eric Christiansen were most helpful and are very appreciated.
REFERENCES Aberg, G., 1988. Middle Proterozoic anorogenic magmatism in Sweden and worldwide. Lithos, 21: 279289. Anderson, J.L., 1980. Mineral equilibria and crystallization conditions in the late Precambrian Wolf River rapakivi massif, Wisconsin. Am. J. Sci., 280 289-332. Anderson, J.L., 1983. Proterozoic anorogenic granite plutonism of North America. Geol. SOC.Am., Mem., 161: 133-152. Anderson, J.L., 1987a. The origin of A-type Proterozoic magmatism: A model of mantle and crustal overturn. Geol. SOC.Am., Abstr. Prog., 1 9 571.
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Shieh, Y.-N., 1983. Oxygen isotope study of Precambrian granites from the Illinois Deep Hole Project. J. Geophys. Res., 88: 7300-7304. Shieh, Y.-N., Schwarcz, H.P. and Shaw, D.M., 1976. An oxygen isotope study of the Loon Lake Pluton and the Apsley Gneiss, Ontario. Contrib. Mineral. Petrol., 54: 1-16. Silver, L.T, 1978. Precambrian formations and Precambrian history in Cochise County, southeastern Arizona. In: J.E Callendar, J.C. Wilt, J.C. and R.E. Clemons, Land of Cochise. N.M. Geol. SOC., 29th Field Conf., pp. 157-163. Silver, L.T and Barker, E, 1967. Geochronology of Precambrian rocks of the Needle Mountains, southwestern Colorado: Part 1, U-Pb zircon results. Geol. SOC.Am., Abstr. Progr., Spec. Pap., 115: 204205. Silver, L.T. and McKinney, C.R., 1962. U-Pb isotope age studies of a Precambrian granite, Marble Mountains, San Bernardino County, California. Geol. SOC.Am., Spec. Pap., 7 3 65. Silver, L.T, Williams, I.S. and Woodhead, J.A., 1981. Uranium in granites form the southwestern United States: Actinide parent-daughter systems, sites, and mobilization. U.S. Dept. Energy, Open File Rep., GJBX-45, 381 pp. Steiger, R.H. and Wasserburg, G.J., 1966. Systematics in the Pbzo8-Thuz, Pb210-U235,and Pbzo6-Uu8 systems. J. Geophys. Res., 71: 6065-6068. Stern, R.J. and Hedge, C.E., 1985. Geochronologic and isotopic constraints on late Precambrian crustal evolution in the eastern Desert of Egypt. Am. J. Sci., 285: 97-127. Stewart, J.H. and Carlson, J.E., 1978. Geologic Map of Nevada, scale: 1 : 500000. Subbarayudu, G.V., Hills, A.E and Zartman, R.E., 1975. Age and Sr isotopic evidence for the origin of the Laramie anorthosite and syenite complex, Laramie Range, Wyoming. Geol. SOC.Am., Abstr. Progr., 7 1287. Sutton, J. 1963. Long-term cycles in the evolution of continents. Nature, 198: 731-735. Swan, M.M., 1976. The Stockton Pass Fault, an Element of the Texas Lineament. M.S. Thesis, University of Arizona, Tucson, Ariz., 119 pp. Taylor, H.P., 1978. Oxygen and hydrogen isotope studies of plutonic granitic rocks. Earth Planet. Sci. Lett., 38: 177-210. Thomas, J.J., Schuster, R.D. and Bickford, M.E., 1984. A terrane of 1350-1400 m.y. old silicic and volcanic and plutonic rocks in the buried Proterozoic of the midcontinent and the Wet Mountains, Colorado. Geol. SOC.Am., Bull., 9 5 1150-1157. Thomas, W.M., Clarke, H.S., Young, E.D., Orrell, S.E. and Anderson, J.L., 1988. Precambrian granulite facies metamorphism in the Colorado River region, Nevada, Arizona, and California. In: W.G. Ernst (Editor), Metamorphism and Crustal Evolution of the Western United States, Rubey Volume VII. Prentice Hall, Englewood Cliffs, N.J., pp. 527-537. Van Schmus, W.R., Medaris, L.G. and Banks, P.O., 1975. Geology and age of the Wolf River batholith. Geol. SOC.Am., Bull., 86: 907-914. Van Schmus, W.R., Bickford, M.E. and Zietz, I., 1987. Early and Middle Proterozoic provinces in the central United States. In: A. Kroner (Editor), Proterozoic Lithospheric Evolution. Am. Geophys. Union, Geodyn. Ser., 1 7 43-68. Watson, E.B. and Harrison, TM., 1983. Zircon saturation revisited: temperature and compositional effects in a variety of crustal magma types. Earth Planet. Sci. Lett., 6 4 295-304. Wenner, D.B. and Taylor, H.P., 1976. Oxygen and hydrogen isotope studies of a Precambrian graniterhyolite terrane, St. Francois Mountains, southeastern Missouri. Geol. SOC.Am., Bull., 87: 15871598. Whalen, J.B., Currie, K.L. and Chappell, B.W., 1987. A-type granites: geochemical characteristics, discrimination, and petrogenesis. Contrib. Mineral. Petrol., 95: 407-419.
Role of anorogenic granites in the Proterozoic crustal development of North America
299
White, D.L., 1978. Rb-Sr isochron ages of some Precambrian plutons in south-central New Mexico. IsochronNest, 21: 8-14. Wilson, M.R. and Akerblom, G.V., 1982. Geological setting and geochemistry of uranium-rich granites in the Proterozoic of Sweden. Mineral. Mag., 46: 233-245. Windley, B.E, 1984. The Evolving Continents. Wiley, New York, N.Y., 2nd ed., 399 pp. Windley, B.E, 1989. Anorogenic magmatism and the Grenville orogeny. Can. J. Earth Sci., 26: 479-489. Wobus, R.A., 1984, An overview of the Precambrian geology of the Tusas Range, north-central New Mexico. N.M. Geol. Soc. Guidebook, 35th Field Conf., Rio Grande Rift, Northern New Mexico, pp. 193-198. Wobus, R.A. and Hedge, C.E., 1980. Rb-Sr isochron age of Precambrian plutons of the San Pedro Mountains, north-central New Mexico. Isochron/West, 27: 19-25. Wooden, J.L. and Miller, D.M., 1990. Chronologic and isotopic framework for Early Proterozoic crustal evolution in the eastern Mojave Desert Region, SE California. J. Geophys. Res., 95: 20,133-20,146. Wooden, J.L., Ashwal, L.D., Wiebe, R.A. and Emslie, R.F., 1987. Regional Pb isotopic systematics in Proterozoic intrusives, Nain province, Labrador. EOS, Trans. Am. Geophys. Union, 6 8 1519. Wu, T-W. and Kerrich, R., 1986. Combined oxygen isotope-compositional studies of some granitoids from the Grenville Province of Ontario, Canada: implications for source regions. Can. J. Earth Sci., 2 3 1412-1432. Wyllie, RJ., 1984. Constraints imposed by experimental petrology on possible and impossible magma sources and products. Philos. Trans. R. Soc. London, Ser. A, 310 439-456. Wyllie, P.J., 1988. Magma genesis, plate tectonics, and chemical differentiation of the Earth. Rev. Geophys., 26: 370-404. Young, E.D., Anderson, J.L., Clarke, H.S. and Thomas, W.M., 1989. Petrology of biotite-cordieritegarnet gneiss of the McCullough Range, Nevada, I. Evidence for Proterozoic low pressure fluidabsent granulite grade metamorphism in the southern Cordillera. J. Petrol., 30: 39-60. Zen, A-En, 1989. Plumbing the depths of batholiths. Am. J. Sci., 289: 1137-1157. Zen, A-En and Hammarstrom, J.M., 1984. Magmatic epidote and its petrologic significance. Geology, 12: 515-518.
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301
Chapter 8
PROTEROZOIC GRANULITE TERRANES S.L. HARLEY
INTRODUCTION
The ubiquitous presence of granulites in continental shield areas indicates that high-grade metamorphism has been a fundamental process in the development and stabilisation of Precambrian deep crust. Recent controversy concerning granulite metamorphism and lower crustal processes has focussed, firstly, upon the largescale tectonic implications of variations in granulite pressure (P)-temperature ( T ) regimes and pressure-temperature-time paths (P-T-t); and secondly, definition of the physical and chemical processes responsible for the low water activities generally characteristic of granulites. Whereas some studies (e.g. Bohlen, 1987) have interpreted a general uniformity of P - T conditions and paths for granulites, it has recently been emphasised that granulite terranes display diversity in these features. A spectrum of tectonic models or settings is required to explain, for example, contrasting near-isothermal decompression (ITD) and near-isobaric cooling (IBC) post-peak P - T histories (Ellis, 1987; Harley, 1989), particularly in the light of structural and geochronological data bearing on the relative timing and timescale of metamorphism. The fundamental question of whether low water activities in high-grade terranes have been produced through the passage of an externally derived C02-rich fluid (e.g. Newton et al., 1980; Touret, 1985, 1986), extraction of water-bearing partial melts, metamorphism of initially dry rocks, or combinations of all of these (Valley and O'Neil, 1984; Lamb and Valley, 1985; Valley et al., 1990) can also be usefully discussed on the basis of data obtained from Proterozoic granulite terranes. It is clear that further progress in understanding the geological processes related to granulite metamorphism requires an integration of P-T and P-T-t studies with investigations of geochemical (e.g. partial melting, fluid migration) and deformational processes in terranes which show varied P - T paths, deformational styles, and exhumation histories. These major issues are addressed here through a consideration of Proterozoic granulite terranes, which have proved critical in the development of current ideas on deep-crustal processes. Several of the best-described granulite terranes in terms of data on P - T conditions, fluid activities, lithological/structural chronologies and isotopic features are Proterozoic in age: for example, the Adiron'dacks, Furua, Finnish Lapland, SW Norway, Sri Lanka and Willyama complexes. In this chapter I present a compilation, description and assessment of the diverse range
302
S.L. Harley
of Proterozoic granulites with an emphasis on those features, parameters and observations of relevance to the general issues of granulite genesis and crustal development. It is not, however, the intention here to present a terrane by terrane account of granulites: such relevant information is to be found in significantly greater detail in the literature cited in Imble 1.
MODES OF OCCURRENCE AND TIME-SPACE DISTRIBUTION
Granulite facies terranes form major components of most continental Precambrian shield areas and are widely, though variably, distributed in space and time within the Proterozoic (Table 2). They are represented to varying extents in most of the major crustal formation or tectonic episodes on all continents. Several modes of granulite occurrence and varieties of terrane are recognised, principally on the basis of scale of exposure, uniformity of metamorphic grade, and nature of the boundaries with adjacent terranes. These include: (a) Discrete granulite blocks or belts (10-103 km2) separated from adjacent, generally lower-grade or different aged domains by younger normal faults or steep shear zones. These fault-bounded granulites could represent isolated uplifted blocks of deep-crustal material or dismembered pieces of formerly extensive terranes, as may be the case for the Eastern Granulites of Imnzania (Coolen, 1980; Key et al., 1989) and the varied granulite blocks of Central Finland (Hollta, 1988). (b) “Straight belts” or mobile zones mainly reworking older basement complexes within or between cratonic blocks. Examples include the Torngat Orogen, Labrador (Mengel and Rivers, 1991), the Laxfordian of NW Scotland (Park, 1981), and the Isortoq region of the Nagsuggtoqidian in Greenland (Glassley and Sorenson, 1980; Kalsbeek et al., 1984, 1987). (c) Thrust-bound granulite sheets, slabs, or imbricated slices, often up to several kilometres in thickness, emplaced onto or against either younger and lower grade nappes or older cratonic foreland areas. The extensive Grenville Front Tectonic Zone (GFTZ) and adjacent parautochthonous belt of eastern Canada (Davidson, 1985; Rivers et al., 1989; Figs. 1, 6) is one important region in which granulites have been thrust against and imbricated with a diverse suite of foreland lithotectonic units (e.g. Indares and Martignole, 1989), although in this case the granulites themselves are domains within and related to the much larger high-grade Grenville Province. Examples of terranes which are more strictly slabs partly bounded by major thrusts include the ’Lapland Granulite Complex or Inari Craton (Hormann et al., 1980; Barbey and Raith, 1990), the Magondi Belt and Masoso granulites of Zimbabwe (Peloar and Kramers, 1989; Treloar et al., 1990), the Albany Mobile Belt of Western Australia (Beeson et al., 1988), the Musgrave Complex (Moore and Goode, 1978), the Pan-African Yaounde Complex of Cameroon (Barbey et al., 1990), the Imataca granulites of Venezuela (Swapp and Onstott, 1989), and the Iforas Granulite Unit of Mali (Boullier and Barbey, 1988). Thrusting and emplacement of the granulite terranes onto other units has
Proterozoic granulite terranes
303
Fig. 1. (a) Tectonic domains and divisions in the Grenville Province of Canada and northeastern U.S.A., modified after Rivers et al. (1989) and Gower (1990). WLT is the Wilson Lake allochthonous terrane; inset A is the Adirondack region detailed in Fig. 2; inset B is the Ontario region detailed in (b). The domain classification is that used in the text. (b) Tectonic domains (outlined by dashed lines) and peak pressure estimates in kilobars (shown as ticked lines) in the SW Ontario region of the Grenville Province, based on Davidson (1985), Anovitz and Essene (1990) and Pattison (1991). Cover sequences distinguished by vertical lined ornament. Tectonic domains as follows: BR = Britt; BU = Bunvash; CMS = Central Metasedimentary Belt; FR = Frontenac Axis; GH = Go Home; KI = Kiosk; M R = Moon River; MU = Muskoka; PE = Pembroke; PS = Parry Sound; RO = Rosseau; SE = Seguin.
w
TABLE 1
B
Pressure-temperature and related data for Proterozoic granulites
P (kbar)
No. TerraneIArealLocality
2200-1900 Ma granulites l a Finnish Lapland, syn-D l b Finnish Lapland, post-D
8.0 7.2 6.2 5-7 8-8.5 5.6-6.4 4.5-5.4 133~2 8- 9 9-10
2 Jequie, Bahai, Brasil 3 Imataca, Venezuela 4a Devon Island, Arctic 4b Ellesmere Island, Arctic 5 South Harris, Scotland 6a Eastern Ghats, Anantagiri(?) 6b Eastern Ghats, Paderu(?) 1900-1600 Ma granulites 7a Arunta, Anmatjira Ra. 7b Arunta, Reynolds Ra. 7c Arunta, Strangways Ra. 7d 7e 8 9 10 11 12 13 14 15 16 17 18 19
Arunta, Strangways (early?) Arunta, Mt Hay Broken Hill, Willyama Mukalo, Gascoyne Province NW Gawler Craton, S. Australia Savo Belt, Finland Central Finland, early Pielavesi, Finland Thrku, Finland W. Uusimaa Complex, Finland Lofoten-Vesteralen, Noway Sth. Ketilidian Belt, Greenland Smithson Bjerge, Greenland Isortoq, Nagssugtoqidian Belt
2.5 4-5 ~
T ("C)
850 830 760 800- 850 750- 800 750- 800 750- 800 825 950 900-1000 >750 750
8.0
850- 920
5.3 7-8 5-6 9-10 10 5-6 8-7 5.5 5-6 3-5 9-12 3-4.5 7.5 7-8
750 800- 850 760- 800 730- 780 1000 730 750 800- 880 800 700- 825 850- 950 700- 800 700 750
dPldT (barPC)
10 9 20-25 18 30-40 30 15-20 3, 50a 30
-loa IBC~ IBC, <3a
dT, dTldt ("C, "CIMa)
300 240 100 200,lO 100 100 250, > 4 275 150
100 >50
aHzO
Features
P-T references
<0.2 0.3 0.3 0.3
T R, my1 Mig Mig C,R+N, Mig T, Mig C, R + N C, R+N, Mig C, R+N, Mig c , R, Spr C, R, Spr
Barbey and Raith (1990), Raith and Raase (1986), Hormann et. al. (1980) Barbosa (1990) Swapp and Onstott (1989) Frisch (1988) Frisch (1988) Wood (1977), Cliff et al. (1983) Sengupta et al. (1990) La1 et al. (1987)
C, N, Mig C, N, Mig
Vernon et al. (1990) Warren and Stewart (1988), Clarke and Powell (1991a) Warren (1983a, 1983b), Warren and Hensen (1989) Norman and Clarke (1990) Harley and Watt (unpublished) Phillips and Wall (1981) Muhling (1988, 1990) Oliver (1989) Korsmann et al. (1988) Hollta (1989) Hollta (1988,1989) Hollta (1986) Schreurs and Westra (1986) Griffin et al. (1978) Dempster et al. (1991) Garde et al. (1984) Glassley and Sorenson (1980), Moecher et al. (1988)
low
0.3-0.5 0.3
90,l 250
C, N, Mig
burial <3a 45 ITD 30 <4,8 12 0-3 0-3 ITD?
S , my1
200, 3 100
0.3-0.6
loo? 200,lO 400, 4 150
0.3 0.1-0.4 very low 0.4-0.6
S, myl, Mig C, Mig T, R, D, my1 R+N, Mig R+N, Mig R+N, Mig N, Mig C, N, T, Mig C, R+N, C, N, Mig R R, D, S, Mig
9
P %
2
9
20
21 22 23 24 25 26
Rinkian Belt, Baffin Is. Torngat Orogen, Labrador Kisseynew, Manitoba McCulloch Ra., western U.S.A. Questa Terrane, New Mexico Magondi Belt, Zimbabwe ArendaVBamble, Noway
3-4 9-10 5-6 4.5 11 6-7 7.3 f 0.5
700800750 725 800750740-
750 850
825 800 860
C , Mig
25 <5 22 25
200,3-5 300, 3 60 250
0.4
200
C , R, D, S C, N, T, Mig
0.03-0.2
<0.3
C, N, Mig C, Mig, Se C, R, T, Mig C, R+N, Mig
Henderson et al. (1989) Mengel and Rivers (1991) Gordon (1989) Young et al., 1989) Grambling et al. (1989) Treloar and Kramers (1989) Lamb et al. (1986), Visser and Senior (1990)
"n
aR
g 3 K' E
Pre-Grenville Polycyclic Terranes (Labradoran) 27 28
Grenville Front, N. Labrador Lake Melville Terrane, Labrador 29a Wilson Lake, Ontario 29b 30 R. Baskatong Terrane, Quebec 1250-950 Ma granulites 31 Namaqualand, Bushmanland 32 Mozambique 33 Madagascar, N belt(?) 34a Musgrave Ra., Australia 34b 34c 34d 35
Musgrave Ra., post-shear Musgrave Ra., West area Giles Complex, Musgraves Albany, W. Australia
36 37
Fraser Ra., W. Australia Eastern Ghats, late event
4-6 10-12
650- 750 700- 770
10-12
900- 980
10.3
790
5.0 12 6.0 10-12
~
850 850? 1000 >860
6.0 <8.0 6.0 5-6
670 750? 800 750- 800
5-6 7-8
800- 8.50 800- 850
7*1 5-7 8 8-10 8.5 8
730 f 50 800 750 700- 800 750 790
>loo
IBC IBC 7-10, 50 0-2 0-3
a
0-3 0-2
a
1OO?
130 100,2-3 80-100 350.2
0.2-0.4
later D, R, T later D, R, T
Owen et al. (1988) Gower and Erdmer (1988)
R+N, Spr N, T ap, R, T
Currie and Gittins (1988), Arima et al. (1986) Indares and Martignole (1990a)
C, R, T, Mig R, N, T
Waters (1988,1989, 1990) Sacchi et al. (1984) Nicollet (1990) Ellis and Maboko (1992) Moore and Goode (1978) Maboko et al. (1989) Clarke and Powell (1991b) Ballhaus and Berry (1991) Stephenson (1984), Beeson et al. (1988) Harley (unpublished) Grew and Manton (1986)
T, Ec N
200
60, 3
>30
<100
2
300
16
100
10-22 20 8,u)
70-150 100 150-60
low
N R R, T D N
Grenville Domains 38 39 40 41a 41b 42
Oaxaca, Mexico Long Range, Newfoundland Gagnone Terrane, Labrador Val-d'Or, Quebec (GFTZ) R. Dozois Terrane, W. Quebec R. Baskatong Terrane, Quebec
am, C, N, Mig pa, Mig, T pa, R, T pa, R, T, C ap, R+N, T
Mora and Valley (1985) Owen and Erdmer (1989) Rivers (1983) Indares and Martignole (1989) Indares and Martignole (1990a) Indares and Martignole (1990a)
a
TABLE 1 (continued)
W 0 o\
No. Terrane/Area/Locality
P (kbar)
T (“C)
dPldT (barPC)
au20 dT,dTldt (“C, “C/Ma)
Features
P-T references
725- 800 625- 725 650- 775 >900?
13 6-12 6 ITD 6
150 50-150 180
am, C, T, Mig am, C, T, Mig am, C, Mig am, C, Spr am, C, Mig pa, T pa, T R, Ec pa?, S, Ec pa, R,D pa,R,D,T
0.1 -0.3 0.1 -0.5 0.1 -0.2 0.07-0.18
ap, c , R, T ap, R, T am, C, T, Se am, C, T, S C C, N C, N, R, Mig C, N, Mig
Indares and Martignole (1990b) Indares and Martignole (1990b) Indares and Martignole (1989) Herd et al. (1986) Perreault and Martignole (1988) Anovitz and Essene (1990) Davidson (1990) Grant (1989) Anovitz and Essene (1990) Anovitz and Essene (1990) Anovitz and Essene (1990) Anovitz and Essene (1990) Anovitz and Essene (1990) Anovitz and Essene (1990) Anovitz and Essene (1990) Edwards and Essene (1988) Edwards and Essene (1988) Bohlen et al. (1985) Lamb and Valley (1988)
0.2
C, N, T R+N c,R, D R, D, T R+N, D, Mig R+N C, Mig C+R, S, Mig C, Mig N, S, my1
Groenewald and Hunter (1991) Grew (1978, 1981) Black et al. (1987) Ellis (1983) Clarke et al. (1989) Fitzsimons and Thost (1992) Stiiwe and Powell (1989a) Fitzsimons and Harley (1991) Thost et al. (1991) Nichols and Beny (1991)
Grenville Domains (continued) 43a 43b 43c 43d 44 45 46 47a 47b 47c 48 49 50 51 52 53a 53b 54a 54b
R. Cabonga Terrane, Quebec Mont-Laurier Terrane, Quebec Morin Terrane, Quebec St. Maurice, Quebec NEQuebec Burwash Domain, Ontario Rosseau subdomain, Ontario Britt subdomain, Ontario Britt subdomain, Ontario Kiosk subdomain Pembroke subdomain Parry Sound Domain Muskoka-Seguin subdomains Central Metasedimentary Belt Frontenac Axis, Ontario Adirondack Lowlands, SW .Adirondack Lowlands, NE Adirondack Highlands Oregon Dome area
7.2-8.8 4.2-6.0 6-8 9-10 4.5-6.5 8.5-9.5 13-15 14-16 10-11 10 8.5-9.5 10-12 10-12 6-8 5.5-6.5 5.4 6-8 7-8.5 7.5
>800
700650715700800 700750700700700 680760 725725-
750 850 875 770 780 800 800 800
200
40-50?
130
10
200,2-4
14
200,2-4
720 820 750
2-5 2-5
a
a
200 200.1-2
East Antarctic
I I@-950 Ma granulites 55 Sverdrupfjella, E. Antarctica 56 Molodezhnaya, E. Antarctica 57a Rayner Complex, Enderby Land 57b Rayner Complex, Kemp Land 58 Mawson Coast, MacRobertson Land 59 Northern Prince Charles Mts. 60a Pxydz Bay, Larsemann Hills 60b Ptydz Bay, Brattstrand Bluffs 60c Pxydz Bay, Bolingen Is. 60d Reinbolt Hills
8
5-6 6-8 8-10 4 6-7 4.5 6 5-7 7
ITD
800
>700 750700750800750850 725800
800 800 800 830 800
825
20-25 30 IBC IBC 20 17 20-30 20
125 150 >lo0 75 50 100
9
P
3
x 9
61 6% 62b 63
Rauer Group, E. Prydz Bay Bunger Hills, Antarctica Bunger Hills, late coronas Windmill Islands, Antarctica
7-9 4 6-7.5 5.5
900-700 Ma granulites 64a Masoso, Zimbabwe, syn-D 64b Masoso, Zimbabwe, post-D 65 Yamato Mts., Antarctica
700-500 Ma granulites 66 Lutzow-Holm Complex 67 68 69a 69b 69c 70 71a 71b 71c 72a 72b 72s 72d 72e 73 74 a
Furua, Tanzania Wami, Tanzania Madagascar, Vohibory Madagascar, Ampanihy Madagascar, Androyan Yaounde, Cameroon Kerala Khondalites(?) Kerala Charnockite event Sri Lanka, Highland Complex Sri Lanka, Highland Complex Sri Lanka, Highland Complex Sri Lanka, Highland Complex Sri Lanka, Kurunegala Leeuwin Block, W. Australia Ox Mts., Ireland
(SE) (SW) (E) (W)
850 800 750 780
30 -40 IBC
12% 2 7 f 2 4.5
725- 800 625- 700 750
50
6-8
750- 810
27-40 <8
10-11 8-9 9-11.5 7-9 4-6 10-12 5-6 5-6 5.5 f 1.9 7-9 5-7 7-9 5-6 6.9 f 1.2 5 10-12
800- 825 700- 820 750- 800 750- 850 850 800 770- 840 775 700- 750 800- 850 800- 850 760- 840 700- 800 738 f 60 >690 800- 900
100 50
0.3-0.4
R+N, D, Mig R+N,D N, D C, N
R, D, T, Mig
It-eloar et al. (1990)
N
Shiraishi et al. (1991)
200, l?
C, N, Mig
275,3-4
C, R, Ec
Motoyoshi et al. (1989) Hiroi et al. (1991) Coolen (1980), Andriessen et al. (1985) Nicollet (1990) Nicollet (1990) Nicollet (1990) Barbey et al. (1990) Santosh (1986,1987) Chacko et al. (1987) Santosh et al. (1990) Faulhaber and Raith (1991) Faulhaber and Raith (1991) Schumacher et al. (1990) Schumacher et al. (1990) Burton and O'Nions (1990) Wilde and Murphy (1990) Sanders et al. (1987)
100
20
ITD
R R, N, T
IBC+ITD
ITD? 14-17 80-100
200 50
10 10 20-30 15-20
200 200 100 100 100,6
0-3
Harley (1988), Stuwe and Powell (1989b) Stiiwe and Powell (1989b) Blight and Oliver (1977)
200
very low low 0.3-0.4
C, S, Mig C, R, Mig C, Mig R, T, S
0.18f 0.16
R R, S, mvl, Ec
3R
gff
'
2 5
E
g R
$2
. ,
Counter-clockwise P-T path inferred by source literature Late P-T estimates from shears or cooled intrusives
Guide to features: am =allochthonous monocyclic domain; ap = allochthonous polycyclic domain; C = cover sequences including metapelites common; D = deformed dykes present; Ec = eclogitic relics present; Mig = migmatitic features common in metapelites; my1 = high-grade syn- to post-peak mylonites; N = new synmetamorphic intrusives abundant; pa = parautotochthonous domain; R = reworking of older crust significant; S = strong shear fabrics, high grade shearing; Se = extensional shear zones/fabrics reported; Spr = sapphirine quartz association present; T = syn- to post-metamorphic thrusts present.
+
w
3
W 0
TABLE 2
03
Time-event chart for Proterozoic granulite formation U.S.A./Canada Scandinavia/U.S.S.R. Timespan
Africa
India/Sri Lanka
Australia
Eastern Ghats?
Gascoyne?
Antarctica
(Ma) 2100-1900
Ellesmere
Inari, Lapland Laxfordian
1900-1600
Nagssugtoqidian Ketilidian Rinkian, west U.S.A. Labradorian Torngat Orogen
Laxfordian Svecofennian (Savo, Pielavesi, Uusimaa, Turku)
1600-1300
New Mexico?
Kongsbergian
1250- 950
G renvil lia n (Quebec, Ontario, Adirondacks) Ottawan
Grenvillian (south Norway)
Namaqua/Kibaran (Namaqualand, Mozambique, Madagascar)
Ox Mts.
Pan-African: Mozambiquian south Madagascar FuruaNami
900- 550
Arunta Willyama Gascoyne
Adelia Land?
Eastern Ghats Sri Lanka?
Albany-Fraser Musgrave Complex
Rayner Complex Prydz-Prince Charles Mts. Bunger Sverdrupfjella
Pan-African: Sri Lanka Highlands Kerala?
Leeuwin
Yamato Mts. Lutzow-Holm
Eburnian (Zimbabwe)
Major events in South America include 2200-1900 Ma in east Brazil and Venezuela; Grenvillian analogues; and Brasiliano (Pan-African).
Proterozoic granulite terranes
309
been shown in some cases to have occurred as a late phase of the tectonism responsible for the high-grade metamorphism itself (e.g. Barbey and Raith, 1990; Treloar et al., 1990; Ellis and Maboko, 1992), whereas in other cases it is a consequence of rejuvenation through tectonic events unrelated to the granuliteforming episode (e.g. Iforas: Boullier and Barbey, 1988; Wilson Lake: Currie and Gittins, 1988). (d) Large-scale (lo4 to lo6 km2) terranes typified by the preservation of relatively uniform granulite assemblages, implying similar physical conditions over very large regions, and a lack of gradual transitions to lower-grade areas. Such terranes are often dissected by later retrograde shear zones, mylonites and faults, but there are generally no substantial differences in metamorphic grade between the blocks on either side of these transgressive zones. Some important examples are the Central Province of the Arunta Complex (Warren, 1983a; Warren and Hensen, 1989; Fig. 3), the Proterozoic of East Antarctica (Harley and Hensen, 1990), the Kerala Khondalite terrane (Chacko et al., 1987) and large tracts of the Grenville Province in SW Ontario and the Adirondack Highlands (Bohlen et al., 1985; Anovitz and Essene, 1990, Figs. l b and 2). (e) Moderate- to large-scale terranes (lo2 to lo6 km2) similar to type (c) above but preserving apparently undisturbed or little-disrupted gradations or transitions from lower to higher grade metamorphic conditions i.e. the granulite terrane forms part of a larger region of varied metamorphic grade in which regional P-T variations are mappable using mineral isograds or assemblage zones. Well-preserved prograde greenschist-amphibolite-granulite sequences and amphibolite-granulite transitions are reported from several Proterozoic terranes including, for example, the classic Willyama Complex (Binns, 1964; Phillips and Wall, 1981), the Adirondack Lowlands (Buddington, 1963; Bohlen e t al., 1985; Edwards and Essene, 1988; Fig. 2), the Central Metasedimentary Belt, Grenville Province (Anovitz and Essene, 1990; Fig. lb), the Anmatjira and Reynolds Ranges, Arunta Complex (Warren and Stewart, 1988; Clarke et al., 1990; Vernon et al., 1990; Fig. 3), the Sri Lanka Highlands (Cooray, 1962; Schumacher et al., 1990; Fig. 4), the Namaqua Province, Southern Africa (Waters, 1986, 1989), and the Lutzow-Holm Complex, East Antarctica (Hiroi et al., 1983; Motoyoshi e t al., 1989). Granulite metamorphism is also intimately associated with mid-Proterozoic anorthosite complexes such as the Nain Complex, Labrador (Berg, 1977) and the Rogaland Complex, Norway (Tbbi et al., 1985). However, as these granulites are low-pressure types of essentially contact-metamorphic origin, exhibiting zonal distributions of assemblages and temperature conditions concentric with the anorthosites, they are not considered further in this analysis. B b l e 1 is a summary of the main metamorphic and geological features of a large number of representative Proterozoic granulite occurrences, grouped in broad time divisions, and forms the basis for most of the discussion in this chapter. In the light of available geochronology, the Proterozoic can be divided into a number of time intervals in which granulite metamorphism has been prominent.
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Fig. 2. Metamorphic conditions in the Adirondack region, Grenville Province, U.S.A. (Bohlen et al., 1985; Edwards and Essene, 1988), an example of a terrane displaying variations in peak temperatures and pressures. Pressures are denoted by circles, and peak temperatures defined using isotherms (solid lines). Dashed lines A to G represent various isograds (Buddington, 1963; De Waard, 1965; Bohlen et al., 1985) as follows: A = garnet in semipelites; B = muscovite out; C = orthopyroxene in amphibolite; D = orthopyroxene in amphibolitic layers within metapelites; E = orthopyroxene in amphibolitic orthogneisses; F = garnet in quartz-deficient metagabbros; G = garnet in quartz-orthopyroxene rocks.
These time intervals are defined and the major “events” or crustal cycles involving granulites within each continent loosely correlated in Bble 2. This representation of the time/space distribution is necessarily somewhat biased and partly flawed, firstly because it is dependent upon geochronology in polymetamorphic terranes in which the relative extents and importance of earIier and later granulite metamorphic events are not yet fully established (e.g. Eastern Ghats: Aftalion et al., 1988; Grew and Manton, 1986; Mozambique Belt: Coolen, 1980; Sacchi et al., 1984; Nicollet, 1990; NE Grenville Province: Rivers et al., 1989; Gower, 1990; East Antarctica: Harley and Hensen, 1990), and secondly because peak metamorphic ages are not well-defined in many granulite belts. The Highland Series or Complex of Sri Lanka (Fig. 4) serves to indicate the necessity of adopting a detailed and multi-technique isotopic approach to define the age of metamorphism. Whereas early Rb-Sr and recent SHRIMP zircon U-Pb data have been interpreted in terms of a 1100-1000 Ma age for the main granulite event
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Proterozoic granulite terranes I
I
I
135’E 1
Amadeus Basin
100 krn 1320E
a
mainly granitoids amphibolite-greenschist f a c i e s
135’E
0
amphibolite f a c i e s granulite f a c i e s
Fig. 3. Metamorphic geology of the eastern Arunta Complex, Australia, modified after Shaw et al. (1984) and Warren and Hensen (1989) and featuring the distinctive tectonic zones noted in the text. Inset map of Australia shows location of the main map (lined area) and of the Willyama Complex ( W ) . Main map shows pressures in kilobars as ornamented circles. NZ = Northern Tectonic Zone; CZ = Central Tectonic Zone; SZ = Southern Tectonic Zone. Heavy solid lines are tectonic zone boundaries. The Anmatjira-Reynolds Ranges region lies to the north and east of the Ngalia Basin, in the N Z ; the Strangways Ranges are north of Alice Springs in the CZ.
(Kroner et al., 1987), further integrated studies have instead pointed to a 607-560 Ma, Pan-African, age for the main metamorphism (Holzl et al., 1988) and an even younger age for charnockite formation near Kurunegala (535 f5 Ma, Burton and O’Nions, 1990). Similar observations apply to the Kerala Khondalite terrane of southern India, where previous Late Achaean to Early Proterozoic ages have been rendered doubtful by new data indicating a much younger metamorphic age, and to several granulite blocks within the Mozambique Belt of eastern Africa and Madagascar (e.g. Andriessen et al., 1985; Chacko et al., 1987; Nicollet, 1990). Clearly, reported metamorphic ages of even well-studied granulite terranes must be regarded with some caution. For this reason the compilation given here (Bbles 1, 2) excludes many probable or potential Proterozoic granulite belts occurring in the former Soviet Union, China, South America and elsewhere. The time intervals used to outline gross correlations are deliberately long (100300 Ma, Table 2), and earth- or continent-wide coeval granulite “superevents”
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Fig. 4. Map of Sri Lanka showing the position of the Highland Series or Complex, cover (horizontal lined areas), and East and West Vijayan Complexes (stippled). Metamorphic pressures in the Highland series are indicated by the ornamented circles (Sandiford et al., 1988; Schumacher et al., 1990; Faulhaber and Raith, 1991), and define a marked E-W gradient with higher2 granulites exposed in the east (dashed isobars). Probable thrust contacts are shown with a toothed ornament.
of any particular type or origin are not implied during any of these intervals. Diachroneity in metamorphism and variations in styles of deformation occur between terranes grouped for convenience within the same time/event domain (e.g. Lapland vs. Laxfordian; Arunta vs. Gascoyne; Furua vs. Kerala). Nevertheless, and despite the caveats noted above, some generalisations can be made. A major period of crust formation, reworking, and granulite metamorphism occurred in the North Atlantic Craton, taken here to3nclude the Baltic, Greenland and Northern American shields, at 2100-1900 Ma (Barbey and Raith, 1990); relict or partially reworked granulites of similar age may occur in the Eastern Ghats (Aftalion et al., 1988) and granulite grade reworking affected older, mainly Archaean, crust in the Amazonian Craton of South America (Barbosa, 1990). Granulite terranes metamorphosed at 1850-1600 Ma are widespread in the North Atlantic Craton, Australia, and probably Africa (e.g. Treloar and Kramers, 1989), but have not been reliably reported from India or Antarctica. The Kongs-
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bergian granulite regions of SW Norway (e.g. Bamble/Arendal: Lamb et al., 1986; Visser and Senior, 1990) and terranes in western U.S.A. (Grambling et al., 1989) may be somewhat younger than 1700 Ma, but are certainly distinct in age from the important 1250-950 Ma group. Reworking of Archaean crust is significant in some of the 1850-1600 Ma granulite terranes (e.g. Nagsuggtoqidian Belt, Greenland - Kalsbeek et al., 1987; Torngat Orogen,.Labrador - Mengel and Rivers, 1991; Lofoten - Griffin et al., 1978; Gascoyne Province, Western Australia - Muhling, 1988, 1990), but accretion of juvenile crust is also evident (e.g. Ketilidian - Kalsbeek and Tmylor, 1985) and high-grade metamorphism of Proterozoic cover sequences common (Schreurs and Westra, 1986; Kalsbeek et al., 1987; Dempster et al., 1991). Proterozoic supracrustal sequences comprise a major component of the granulites in the Arunta and Willyama complexes of Australia (Stewart et al., 1984; Stevens et al., 1988; Fig. 3). The 1250-950 Ma time interval is notable for the vast granulite terranes of types (d) and (e) above, including much of the Grenville Province (Mora and Valley, 1985; McLelland et al., 1989; Rivers et al., 1989; Daly and McLelland, 1991), the East Antarctic Rayner Complex and correlatives (Grew, 1978, 1981; Black et al., 1987; Harley and Hensen, 1990), the Namaqua Province (Waters, 1986) and Kibaran granulites of Africa, and the Albany-Fraser and Musgrave complexes of Australia (Maboko et al., 1989). Widespread granulite overprinting and reworking probably also occurred in the Eastern Ghats at ca. 1000 Ma (Grew and Manton, 1986; Aftalion et al., 1988). These younger Proterozoic granulite terranes do not inherently consist of reworked older crustal material; new crustal accretion can be identified in most Proterozoic granulite terranes, including the Grenville Province (Gower, 1990; Daly and McLelland, 1991). Late Proterozoic granulites (<900 Ma) are also common, but the present data indicate that they are mostly restricted to the Gondwana fragments (Africa, South America, Antarctica, Sri Lanka). The most important group includes those formed in the 700-570 Ma interval, corresponding with the later phase of post-rift PanAfrican compression, magmatism and tectonism in Gondwana (Porada, 1989). This group includes the Highland Series of Sri Lanka (Holzl et al., 1988), Furua and associated Eastern granulites of Thnzania (Andriessen et al., 1985), and perhaps the Kerala Khondalites (Chacko et al., 1987): quite a varied suite in terms of lithologies, metamorphic P-T conditions, and regional settings (Tmble 1).
STRUCTURAL STYLES AND SEQUENCES
Given the often complex interplay of deformation, magmatism and melting in granulite metamorphism (e.g. Hopgood, 1984), and the local or regional significance of reworking of older crustal blocks in many Proterozoic granulite terranes, it is important to distinguish those structural features and sequences associated with any particular high-grade metamorphism. This can be achieved using either cross-cutting field relations or by study of rock sequences with little structural
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Fig. 5. Some geological features of a Proterozoic granulite terrane: the Rauer Group, East Antarctica. a. An example of high-grade reworking. Early flat-lying fabrics defined by mineral orientations and layering in a composite felsic-mafic orthogneiss are cut by upright, granulite grade, cleavages and shear zones in which the mesoscopic layering is severely attenuated. Mineral elongation lineations are nearvertical in the latter zones. Note early discordant leucosomes in the flat zones, and later leucosomes subparallel to the upright axial surfaces. Rauer Group, Antarctica. Lens cap is 50 mm across. b. High-grade reworking of flat-lying to reclined structures by upright folds and ductile shear zones, Rauer Group, East Antarctica. The folded mafic layers include isoclinally folded metamorphosed dykes. In the shear zone on the right side of the photograph the mafic layers are attenuated into lenses and sheaths with long axes defining a near-vertical stretching direction.
Proterozoic granulite terranes
3 15
inheritance. Clearly, the structural elements related to a particular granulite event are best defined in those belts dominated by supracrustal sequences which have only had a sedimentary record prior to metamorphism. These are typified by the monocyclic belts in the eastern parts of the Grenville Province, including the Central Metasedimentary Belt of Ontario and the Adirondacks (Rivers et al., 1989; Daly and McLelland, 1991), where mid-Proterozoic sedimentary sequences such as the Grenville Supergroup have been subjected to granulite metamorphism and deformation at 1100-1050 Ma. Other essentially “monocyclic” granulite belts or domains include the Willyama Complex (Stevens et al., 1988), parts of the Arunta Complex (Shaw et al., 1984), the Bushmanland Province of Namaqualand (Waters, 1989), the Prydz Bay paragneisses of East Antarctica (Sheraton et al., 1984; Fitzsimons and Harley, 1991) and Lutzow-Holm Bay (Motoyoshi et a]., 1989), to name but a few in which cover sequences form an important part of the metamorphic pile (Xible 1). A common structural evolution in these terranes involves the formation of isoclinal folds, often flat-lying but in some cases initially upright, which are then progressively refolded about recumbent isoclinal or asymmetric close folds and slide zones (e.g. Marjoribanks et al., 1980; Hopgood, 1984; McLelland and Isachsen, 1985; Schreurs and Westra, 1986; Clarke, 1988) to produce a recumbent gneiss complex. Subsequent large- and small-scale refolding about steep or upright open to isoclinal folds of several generations and orientations (e.g. Hopgood, 1984; McLelland and Isachsen, 1985), or reorientation and attenuation in upright highgrade ductile shear zones (Harley, 1987), are typical features of syn- to post-peak high grade structural histories but are highly variable in intensity on outcrop scales (Figs. 5a, 5b). Later deformation is often restricted to discrete mylonite zones, pseudotachylite arrays, and linear amphibolite to greenschist facies retrograde schist zones (RSZs) which divide granulite terranes into blocks and across which
c. An example of rhythmic and graded igneous layering preserved in a rock of dioritic composition, metamorphosed under granulite conditions. This ca. 1030 Ma orthogneiss pre-dates the late deformation defined by upright structures in (b) above and contains a foliation, yet preserves fine-scale igneous layering features. Rauer Group, East Antarctica. d. Graded igneous layering beautifully preserved in a gabbroic orthogneiss, Rauer Group, East Antarctica. This unit also preserves pre-deformational brittle faults (above the 20 cm long pencil) and locally rhythmic layering. Way-up is to the top right of the photograph. Such layering is preserved despite a profound deformational history: the whole mafic unit is isoclinally folded on a scale of 50-100 m and has subsequently been boudinaged on a similar scale and then refolded about open upright flexures correlated with the late structures in (b). This rock unit is older than that in (c). e. Deformed granulite-facies mafic dyke which locally cuts earlier layering and high-grade mineral fabrics in felsic-mafic composite gneisses, Rauer Group, East Antarctica. The later deformation affecting the dyke is near-coaxial with the pre-dyke one. A weak foliation occurs in the dyke parallel to the 16 cm long pen. Note amphibole-rich margins on the dyke. f. At least two, and probably three, different orientations of former mafic dykes dissect felsic orthogneisses at Lunar Island, Rauer Group (Harley and Fitzsimons, 1991). In this view, these dykes are deformed but are still discordant to the gneisses and each other. An upright zone of high shear strain, similar to (b), occurs on the left of the photograph. Length of outcrop approximately 60 m.
3 16 NW
S.L. Harley SE
Fig. 6. Schematic cross-section (length 300 km), here through the Grenville Province in Ontario (Rivers et a]., 1989), across a Proterozoic granulite terrane dissected by internal and bounding thrusts. MBBZ = monocyclic belt boundaly zone, which carries allochthonous monocyclic tectonic domains; ABT = allochthon boundaly thrust; GFTZ = Grenville Front Tectonic Zone. The foreland in this case is the Superior Province. Note that the MBBZ is interpreted to cut the ABT in this case (Rivers et al., 1989), and hence is out-of-sequence if NW-directed thrusting is inferred.
significant vertical displacements may occur (Willyama Complex: Marjoribanks et al., 1980; Phillips and Wall, 1981; Arunta Complex: Shaw et al., 1984; Fig. 3). Proterozoic granulite terranes in which early, pre- to peak-metamorphic flatlying fabrics have been identified, and in which recumbent large-scale structures (e.g. nappes) have been inferred include, among others, the Lapland granulites, the Willyama Complex, the Adirondacks, the Strangways Ranges in the Arunta Complex, and the Rayner Complex in East Antarctica (Hormann et al., 1980; Sheraton et al., 1980; Marjoribanks et al., 1980; Wiener et al, 1984; Shaw et al., 1984; Clarke, 1988). Small-scale structures and fabrics associated with the early, high-grade, deformation events may include layer-parallel foliations, stretching and mineral rodding lineations which may be colinear with mesoscopic fold axes, curved fold axes and sheath folds with axes in the extension direction defined by the stretching lineations, fold limb attenuation, and boudinage of competent layers (e.g. Park, 1981; Hopgood, 1984; Clarke, 1988). In some instances where sub-horizontal deformation has outlasted peak metamorphic grain growth, for example parts of the Arunta Complex (Norman and Clarke, 1990), protomylonitic and flaser fabrics overprint and rework the previous tectonite fabrics which suffer grainsize reduction even at granulite grades. Such textural features are indistinguishable from those related to reworking of older high-grade gneisses in other polycyclic terranes (e.g. Gascoyne Province: Muhling, 1988, 1990). The tectonic significance of horizontal structures in these granulite terranes is problematic, as they could reflect either compressional or extensional regimes (Park, 1981; Sandiford, 1989). Regionally consistent foldhhear fabric vergence relations, and sense of shear indicators concordant with movement directions inferred from stretching lineations, have been described from several terranes (Clarke, 1988; Barbey and Raith, 1990). These features, which are consistent with non-coaxial deformation such as progressive simple shear, are clearly related to syn-metamorphic thrusting in some terranes (e.g. Beloar, 1988; Barbey and Raith, 1990), and peak metamorphism is considered to be synchronous with compression in many low-pressure granulites (Loosveld and Etheridge, 1990). Progressive shear has also been inferred from the development of mineral elongation and stretching
Proterozoic granulite terranes
3 17
lineation fabrics in and boudinage of successive leucosomes formed in axial planes of early fold generations (Hopgood, 1984). However, consistent structural asymmetry is not apparent in many terranes, and the structures may be more consistent with coaxial flattening. Internal thrusting is now recognised as an important component of the highgrade structural history of many Proterozoic granulite terranes, whether of reworked type or not (Treloar, 1988; Indares and Martignole, 1989; Rivers et al., 1989; Waters, 1990). Within the western Grenville Province and in the many cases of thrust-bound or imbricated granulite slabs noted above (Figs. 1 and 6), mylonitic tectonite fabrics developed at granulite, or even eclogite grade are directly correlated with thrust surfaces (i.e. thick-skinned thrusting) and reverse ductile shears. In some examples these features are also synchronous with the emplacement of hot granulites onto footwall rocks which then undergo heating (e.g. Lapland: Barbey and Raith, 1990; Imataca: Swapp and Onstott, 1989; Magondi: Treloar, 1988). However, in other cases (e.g. Iforas: Boullier and Barbey, 1988; Zambesi Belt: Treloar et al., 1990) the fabrics, and even thrust geometries, are correlated in time with decompressional P-T paths. In the Grenville Province, for example, cooling and some decompression were occurring in the westerly granulite domains at 1150 Ma, while west-directed thrusting was occurring further east. This paradox can be resolved if the granulite units were unroofed either by backthrusting or extension and lateral spreading at higher crustal levels, synchronous with the observed deeper-level thrusts (Rivers et al., 1989). Evidence consistent with an extensional origin for horizontal structures is seen in some Proterozoic granulite terranes. Relatively flat-lying high-grade extensional shear zones, and extensional fabrics superficially similar to S-C fabrics (BCrthe and Brun, 1980), have been recognised from the low? granulites of the Ketilidian Belt, Greenland (Dempster et al., 1991), the Lapland granulites (Barbey and Raith, 1990) and in the Prydz Bay paragneisses (Fitzsimons and Harley, 1991). Barbey and Raith (1990) define an early set of west-facing isoclines and reverse ductile shear zones partially overprinted by leucosomes in migmatites from the Lapland Complex. These structures are in turn overprinted by east-facing folds and extensional shears bearing rodded leucosomes and high-grade mineral assemblages, interpreted to result from syn-metamorphic EW extension. In the Prydz Bay paragneisses, early fabrics and leucosomes are reoriented into and attenuated within subhorizontal granulite-grade shears, which show a regionally consistent shear sense (Fitzsimons and Harley, 1991). As a decompressional P-T history is deduced from the assemblage evolution associated with this deformation, these zones are interpreted as extensional structures. Successful interpretation of the tectonic settings of Proterozoic granulites requires such integration of structural chronologies or kinematic indicators and metamorphic P-T histories. Despite the often profound effects of the deformation events, there are some Proterozoic granulite terranes in which records of earlier sedimentary and igneous structures are preserved in low-strain domains or across zones transitional to lower grades. In the Willyama Complex, the early large-scale recumbent deformation
3 18
S.L. Harley
phase is defined by regional-scale structures, inferred to be downward-facing from inverted modified sedimentary structures including graded beds, cross bedding and syn-sedimentary slumps (Marjoribanks et al., 1980). Recognition of this, coupled with the availability of extensive drill core data, has led to the development of terrane-wide lithostratigraphic correlations in this terrane. Although there is disagreement on the details of what the precursor sediments or volcanogenics were, there is broad consensus on the validity of the lithostratigraphic approach (Stevens et al., 1988). A similar approach has been successfully applied to the paragneiss sequences of the Adirondack Lowlands and southern Highlands, where lateral sedimentary facies variations, pinch-outs, and unconformities have been deduced (Wiener et al., 1984; McLelland and Isachsen, 1985). Granulites of the Arunta Complex have been divided into three lithostratigraphic groups separated by unconformities, which are disrupted or transposed but can nevertheless be recognised as such in places (Stewart et al., 1984). In the Reynolds Range area for example (Fig. 3), a distinctive stratigraphic succession, the Reynolds Range Group, is laterally continuous from low to high metamorphic grades and preserves relict sedimentary structures (graded layering, carbonate debris flows). A folded and partly transposed unconformity between this cover and an underlying granulite basement metamorphosed in an earlier event is recognised even where the second metamorphism is at high grade (Shaw et al., 1984; Warren and Stewart, 1988). Examples of relict igneous structures,including delicate graded and rhythmic layering, preserved within othenvise strongly deformed orthogneisses are also reported from granulite terranes (Harley, 1987; Figs. 5c, 5d) and, like the sedimentary features, provide valuable facing criteria.
COMPLEXITY IN STRUCTURAL AND METAMORPHIC EVOLUTION THE ROLE OF REWORKING
Reworking of older crustal basement is an important feature of many Proterozoic granulite terranes (see lhble l), and structural sequences such as those described above may be superimposed on already complex and non-planar rock structures and relations. In the absence of cross-cutting relationships between younger intrusives and older basement gneisses, the extent of reworking may be impossible to evaluate without detailed isotopic data. For example, in the 1850 Ma Nagsuggtoqidian Belt, basement orthogneisses of Archaean age are indistinguishable in the field from isotopically juvenile Early Proterozoic orthogneisses (Kalsbeek et al., 1987). The distinction of lithotectonic domains or provinces, an approach developed extensively by workers in the Canadian Grenville Province (Davidson, 1985; Rivers et al., 1989), provides a useful framework for description and characterisation of many granulite terranes. The Grenville Province is composed of numerous shear/thrust bounded tectonic domains on the medium to large scale (lo's to 100's km; Fig. 1).Adjoining domains may be distinct in terms of metamorphic P-T
Proterozoic granulite terranes
3 19
conditions, structural histories and lithological constitution, although in some cases only minor differences are apparent (Culshaw et al., 1988; Indares and Martignole, 1990b; Anovitz and Essene, 1990). Within this framework, three domain/terrane types are recognised: (a) Parautochthonous terranes (pa, Bble 1) consisting mainly of reworked and partially overprinted gneisses derived from the adjacent foreland, as indicated by features such as progressively deformed dyke swarms which are found in pristine condition in the adjacent craton (Bethune, 1989). Internal thrusts, imbrication and variable metamorphic grades are features of some of these terranes (Rivers, 1983; Indares and Martignole, 1989). (b) Allochthonous polycyclic terranes (ap, Bble 1) which record high-grade metamorphic and deformational events pre-dating the Grenvillian ones, but which do not correlate with events in the foreland. The intensity of later overprinting is highly variable, and may be insignificant in some polycyclic terranes, suggesting that these regions had escaped significant burial in the later tectonism (e.g. Wilson Lake terrane, Labrador: Currie and Gittins, 1988). (c) Allochthonous monocyclic terranes, as described above, which have not experienced a major pre-Grenvillian tectonic event and include little reworked material. These terrane types are broadly grouped into larger-scale belts essentially aligned parallel to the trend of the Grenville Front (Fig. 1) and separated by major diachronous thrusts, or even folded extensional shear zones, along which large horizontal displacements (30-100 km) are inferred between 1150 and 1050 Ma (Rivers et al., 1989; Fig. 6 ) . The geometry and kinematics of the younger high-grade deformation can be well constrained within polycyclic and parautochthonous terranes if the overprinting is heterogeneous and non-pervasive in character and planar markers such as dykes are present in various original orientations. This approach has been extensively used to document the extent of reworking (e.g. Greenland: Escher and Watterson, 1974; Grenville Front Tectonic Zone: Davidson, 1985), and can define parautochthonous domains if desired. The presence of allochthonous polycyclic domains, on the other hand, are more difficult to demonstrate since their definition requires extensive isotopic data. 1250-1100 Ma mafic dykes suites in East Antarctica, which dissect Archaean cratons, are complexly deformed in reworked zones within the Rayner Complex and Rauer Group (Sheraton et al., 1980, 1987; Sheraton and Collerson, 1983; Harley, 1987). In the latter area, the deformation patterns of intersecting dyke generations (Figs. 5e, 5f) constrain the overall upright geometry of the main post-dyke high-grade deformation. This distinctive high-grade structural style involving dominantly upright folding, ductile shearing, and steep down-dip stretching lineations has also been recognised in the linear granulite belts or “mobile zones” which dissect earlier high-grade basement terranes (Glassley and Sorenson, 1980; Kalsbeek et al., 1984, 1987) or rework adjoining cratonic blocks (Mengel and Rivers, 1989). This structural style, which involves both shortening and simple
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shear, has in general been attributed to transpression because of asymmetry in deformed dyke geometries, and may in some cases post-date compressional folding and thrusting a t lower grades, as appears to be the case in the Torngat Orogen (Van Kranendonk, 1990). Similar early upright folding without large-scale horizontal displacements has also been recognised in some monocyclic terranes in which lithostratigraphic continuity from low to high grades can be demonstrated (Marjoribanks et al., 1980), although these structures may be overprinted by flat-lying ones. PRESSURE-TEMPERATURE CONDITIONS OF PROTEROZOIC GRANULITE TERRANES
P-T estimation: geothermobarometric approaches and uncertainties Two complementary but inherently different approaches, which yield concordant results in many high-grade terranes (Bohlen, 1987) but may be discrepant in others, are used to evaluate the P-T conditions and histories of granulites: quantitative geothermbarometry (GTB) on the one hand and detailed mineral assemblage/petrogenetic grid constraints on the other. With the development and application of internally consistent thermodynamic datasets for end-member minerals and exchange vectors (e.g. Holland and Powell, 1990) it is now possible to unify the two approaches for specific chemical/rock systems and obtain calculated P-T data, for example “average-pressure” estimates (Powell and Holland, 1988), defined by all the equilibria which limit the natural assemblage. It must be recognised, however, that uncertainties in mineral activity-composition relations (a-X) still limit the accuracy of these and conventional thermobarometric P-T calculations. Figure 7 compares alternative thermodynamic calibrations of the popular GAES barometer (Essene, 1989), which is based on the Mg-end member equilibrium relation between the phases garnet, orthopyroxene, plagioclase and quartz, a common assemblage in granulite facies felsic to mafic orthogneisses. The calculated end-member reaction (logK = 0) and lines of equal log(equi1ibrium constant) (log K ) are well-positioned and there is excellent agreement between independent calculations. However, different garnet activity models profoundly effect the calculated logK values and hence pressure estimates (Harley, 1989; Perkins, 1990), particularly where the garnets are Fe-rich. LogK values of 2.5 to 3.8 are relevant to this equilibrium for garnet-orthopyroxene-plagioclase-quartz assemblages in many terranes, including the medium to high-pressure granulite domains of SW Ontario and the Adirondack Highlands in the Grenville Province (shaded box, Fig. 7). Although this good calibrational agreement exists for a number of geobarometers (Essene, 1989), significant calibrational discrepancies are apparent in others. For example, logK lines calculated by different workers for the widely used Fe end-member equivalent of the GAES reaction, G U S (garnet-anorthite-ferrosilitesilica), differ by up to 1.5 kbar (equivalent to ca. 0.5logK units). Clearly,
Proterozoic granulite terrmes
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20
15
10
(kbar) 5
log K=-5- - - - - - - - - - - - - __
k3 600
700
800
900
T ("(2) Fig. 7. Positions of isopleths of fixed loglo K displaced from the magnesium end-member garnetorthopyroxene-plagioclase-quartzequilibrium, GAES (loglo K = 0). K expression is in terms of activities of components in the participating phases: a h , activity of anorthite in plagioclase; U E " , activity of enstatite in orthopyroxene; acrsand uprp, activities of grossular and pyrope components, respectively, in garnet. Solid and dashed lines are the isopleths of Essene (1989) and Newton and Perkins (1982), respectively. The dotted isopleths have been calculated using the dataset of Holland and Powell (1990). Note the good agreement between the calibrations. Shaded box defines the P-T conditions for many granulites, including those from the Grenville Province (Table 1).
alternative calibrations and mineral a-X models may be important sources of spurious variations in calculated P-T conditions, and must be accounted for when P-T comparisons are made between granulites. Despite these reservations, the geobarometry of granulite terranes is now reasonably precise. A number of reliable methods are available for assemblages in metamorphosed felsic, mafic and pelitic rock-types, and these are often in good agreement with dataset-based calculations. As a consequence, we can be reasonably confident that the large pressure range calculated for the spectrum of Proterozoic granulites (Bble 1; Figs. 8, 9) is real and that significant pressure differences occur even within terranes. Temperature estimates for the granulites listed in a b l e 1 have mostly been based on exchange thermometry involving the FeMg-l, NaK-l and other vectors in coexisting mineral pairs. Notwithstanding the discrepancies apparent between alternative thermometer calibrations (Bohlen, 1987; Harley, 1989; Perkins, 1990), it is well-documented that cation exchange often occurs down to closure temperatures significantly less than the peak conditions attained in the higher-T granulites (Ellis, 1980; Harley, 1985; Bohlen, 1987). As a consequence, because most geobarometers are also somewhat T-dependent, erroneous "peak" P-T estimates may result. Furthermore, spurious P-T paths might be generated when mismatch
S.L. Harley
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occurs between the closure temperature for exchange and that for growth or reaction of a mineral involved in a barometric net-transfer equilibrium (Harley, 1985, 1989; Perkins, 1990); the so-called “granulite uncertainty principle” (Frost and Chacko, 1989). Harley (1989) also describes a more subtle effect of mismatched cation exchange on geobarometry, that of feedback, wherein the barometer has significant compositional dependencies (e.g. dP/dXMg) for the exchanging phases. In this situation, continued XMgchange subsequent to barometer closure leads to a spurious, often low, pressure estimate. Geothermometry remains the major problem in definition of the physical conditions of granulite terranes. Uncertainties in temperature have important consequences for calculations of fluid activities and compositions, estimation of potential extents of melting, and definition of temperature-time histories and cooling rates. Only in cases where discrepancies between geothermometric estimates and temperatures implied by well-constrained mineral assemblages are obvious is the full extent of the problem realised (see below). The P-T spectrum from GTB
P-T data for Proterozoic terranes, based mainly on geothermobarometry of felsic to mafic orthogneisses and metapelites are listed in n b l e 1, summarised in the histograms of Fig. 8, and presented in P - T diagrams along with relevant P-T path information in Figs. 9 and 10. The granulites span a large recorded pressure range. Lower-P (3-6 kbar) shallow crustal types, implying extreme transient thermal gradients over relatively small (e.g. Uusimaa: Schreurs and Westra, 1986; Tbrku: Hollta, 1986; Anmatjira Ranges, Arunta Complex: Vernon et al., 1990) to very large (e.g. Namaqualand: Waters, 1989, 1990; Ellesmere Island: Frisch, 1988, Kerala: Santosh, 1987) regions comprise an important group which can be contrasted with markedly higher? terranes (8-12 kbar) which in some areas include eclogite facies metabasites (Grant, 1989; Davidson, 1990; Ellis and Maboko, 1992).
kbar)
650
700 7 5 0
800
850
900 950
T f‘C)
Fig. 8. Histograms of pressure and temperature estimates for peak metamorphism in Proterozoic granulite terranes and tectonic domains. Database and sources are given in Table 1.
Proterozoic granulite terranes
323
Significant P variations are recorded within several terranes. A baric gradient from 5-6 kbar in the Central Metasedimentary Belt to 8-11 and even 12-15 kbar in parautochthonous and polycyclic domains occurring further west has been documented in the Grenville Province in Ontario (Grant, 1989; Anovitz and Essene, 1990; Fig. lb). Although similar pressures are calculated for some adjacent tectonic domains, it is apparent that P-discontinuities also occur, particularly in and across sheared boundaries (e.g. Parry Sound domain: Davidson, 1990). Domain-related P-discontinuities are also recognised in the Grenvillian granulites in Quebec (Indares and Martignole, 1990a, b). Baric gradients over somewhat smaller P intervals have been documented from Finnish Lapland (Raith and Raase, 1986), the Adirondack Highlands (Bohlen et al., 1985) and the Rauer Group (Harley, 1988). In the Highland Series or Complex of Sri Lanka (Fig. 4), pressures vary from >9 kbar in the SE to 6 kbar in the W and NW (Schumacher et al., 1990; Faulhaber and Raith, 1991). Significant regional variations and probable discontinuities in peak pressures are also recorded in the Arunta Complex (Fig. 3), which includes a 3-5 kbar low? granulite belt in its Northern Rctonic Belt (Warren and Stewart, 1988; Clarke et al., 1990; Clarke and Powell, 1991a) and a contrasting higher2 Central zone (Warren, 1983b; Warren and Hensen, 1989). In many of the cases noted above, regional peak temperature variations (9 kbar) granulites also occupy a restricted temperature field (750-850°C), but the bulk of terranes which occur in the intermediate2 range (6-9 kbar) potentially show a wider temperature spectrum.
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Some assemblage constraints on Proterozoic granulites The geothermobarometric results can be compared with constraints imposed by diagnostic mineral assemblages in specific rock-types. Refinement of the threefold division of the granulite P - T regime based upon garnet-forming reactions in silica-saturated and -undersaturated metabasites (e.g. Green and Ringwood, 1967; Fig. 9) is now possible using petrogenetic grids delineating the stability fields of critical mineral assemblages in metapelites (Hensen and Green, 1973; Ellis et al., 1980; Grant, 1985; Hensen, 1986; Clarke et al., 1989), calc-silicates (Ellis, 1978; Warren et al., 1987; Harley and Buick, 1992) and silica-undersaturated aluminous gneisses (Waters, 1986; Hensen, 1987; La1 e t al., 1987). In addition, reaction textures in rocktypes for which reliable .geothermobarometry is lacking can often be interpreted within the framework of such petrogenetic grids and integrated with mineral zoning information to define dPldT vectors and hence build up P - T paths. 14
,
0 " " "
Fig. 9. Pressure-temperature estimates for Proterozoic granulites, grouped according to age divisions and regions. Stippled box: typical P-T uncertainty. Filled circles: 2200-1900 Ma and various granulites; half-filled circles: left half filled, 1900-1600 Ma granulites; right half filled, 1100-900 Ma Antarctic granulites; upper half filled, Grenville Province granulites; lower half filled, 900 Ma and younger granulites. Reactions A and B: incoming of garnet in quartz tholeiite and alkali basalt compositions respectively (Green and Ringwood, 1967). Pig20: approximate lower-T stability of pigeonite with X M = ~ 0.20 (Harley, 1989).An + Cc = Mei: lower-T stability of meionitic scapolite (Mei) with respect to anorthite (An) and calcite (Cc) in calc-silicate granulites (Moecher and Essene, 1990; Harley and Buick, 1992). Grs + Qtz = Wo An: calculated using Holland and Powell (1990), where Grs is grossular, Qtz is quartz. Grr + Sil = Spl + Crd + Qtz: stability of Fe-Mg spinel (Spl) with quartz, for a case where the spinel has 10% non-aluminous components (Waters, 1991).
+
Proterozoic granulite terranes
325
Low pressure (low P I T ) Proterozoic granulite terranes are characterised by distinctive mineral assemblages. The stability of hercynitic spinel with quartz (& garnet, cordierite, sillimanite, biotite) in Fe-rich metapelites is diagnostic of low-pressure granulite metamorphism, provided that the spinel is an Fe-Mg type with only minor Zn and magnetite component (Bohlen et al., 1986; Clarke et al., 1989). A concise review of the occurrence and implications of hercynite-quartz assemblages is provided by Waters (1991). Similarly diagnostic spinel-olivineplagioclase, fayalite-quartz, and wollastonite-anorthite assemblages may occur in mafic, Fe-rich felsic, and calc-silicate rock-types respectively. Early Proterozoic (2000-1700 Ma) examples of spinel-quartz and cordierite-rich granulites include Ellesmere Island and the Ketilidian of Greenland (Frisch, 1988; Dempster et al., 1991), the Svecokarelides of Finland (Schreurs and Westra, 1986; Hollta, 1986, 1989), the Reynolds Range region of the Arunta Complex (Warren and Stewart, 1988; Clarke and Powell, 1991a; Fig. 3) and the McCulloch Range, U.S.A. (Young, 1989). Younger examples include the 1200-1000 Ma Namaqua Province and East Antarctic granulites (Waters, 1986; Clarke et al., 1989; Stuwe and Powell, 1989a; Humphreys and van Bever Donker, 1990) and the potentially Pan-African Kerala Khondalite Belt (Santosh, 1987; Chacko et al., 1987) and Androyan group of South Madagascar (Nicollet, 1990). These low-P terranes are often typified by abundant migmatised metasupracrustal sequences and granitic gneisses. In most cases they are monocyclic belts in the broadest structural or lithotectonic sense, but internal complications such as the presence of polymetamorphosed former basement or successive supracrustal sequences with basementkover relationships are also evident (e.g. Arunta Complex, Reynolds Range region: Shaw et al., 1984; Warren and Stewart, 1988). The low? granulites imply extreme syn-metamorphic thermal regimes and gradients (up to 10O0C/km), are often typified by nearisobaric cooling retrograde P-T paths, and hence are of singular importance as explanations for their origin must account for the production of anomalously hot crust on a regional scale at shallow or mid-crustal levels. High-pressure (>9 kbar) granulite metamorphism is recognised by the diagnostic association garnet-clinopyroxene-quartzin metabasic rocks of intermediate Fe/Mg composition (Green and Ringwood, 1967). This association is recorded in several Middle to Late Proterozoic terranes including the Furua Complex, SW Grenville Province, Zambesi Belt, Musgrave Ranges and SE Sri Lanka Highland Complex (Coolen, 1980; Maboko et al., 1989; Anovitz and Essene, 1990; ?eloar et al., 1990; Schumacher et al., 1990), and in older (2000-1700 Ma) granulites of LofotenVesteraalen (Griffin et al., 1978), the Torngat Orogen (Mengel and Rivers, 1991), and Harris (Wood, 1975, 1977; Cliff et al., 1983). Compatible high2 metapelitic assemblages including kyanite have been recognised from many of these (e.g. Harris, SW Grenville) and other terranes (Yaounde: Barbey et al., 1990), and in a few instances the high2 assemblage orthopyroxene-sillimanite is preserved in local Mg-rich aluminous gneisses (e.g. Anovitz and Essene, 1990). 1200-1000 Ma eclogites and related garnet-clinopyroxene-plagioclase rocks are now recognised in the SW Grenville Province (Grant, 1989; Davidson, 1990) and occur in shear
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zones in the Musgrave Complex (Ellis and Maboko, 1992), corroborating and extending the range of high pressures calculated for these terranes. It is clear that high-P assemblages, previously thought to be rare in granulites, are present in many Proterozoic terranes. The P-T realm defined by the terrane-based data now strongly overlaps that developed from crustal xenolith studies (cf. Griffin and O'Reilly, 1987). Sapphirine-bearing metamorphic parageneses are important indicators of P - T histories and persuasive evidence for higher-T granulite metamorphism in a number of Proterozoic terranes. The rare association sapphirine-quartz (Hensen and Green, 1973; Ellis et al., 1980), and related orthopyroxene-sillimanite and magnesian spinel-quartz assemblages (as distinct from hercynite-quartz), have been described in magnesian paragneisses from the allochthonous polycyclic Wilson Lake terrane, Labrador (Arima et al., 1986; Currie and Gittins, 1988). Peak P-T conditions of 10-12 kbar and >9OO"C, followed initially by near-isobaric cooling to 800°C and then high-T decompression to only 4 kbar, have been deduced for this 1700 Ma terrane based on reaction textures in the sapphirinequartz rocks and presence of younger silica-undersaturated assemblages involving sapphirine, cordierite and spinel. In the Eastern Ghats terrane high-T sapphirine-spinel-quartz assemblages, orthopyroxene-sillimanite, and probable relicts after osumilite, are preserved locally, implying extreme peak granulite temperatures (>950°C) at 8-10 kbar (La1 et al., 1987; Sengupta et al., 1990). Reaction textures involving formation of orthopyroxene sillimanite or garnet between earlier sapphirine, spinel and quartz indicate that the initial post-peak P-T path in this case also involved near-isobaric cooling (Sengupta et al., 1990), succeeded by decompression at a later stage. However, the significance of this extreme metamorphism, and the regional context of the P-T path, is not yet clear as the ages of the assemblages in this terrane are poorly constrained. Although Grew and Manton (1986) deduce a ca. 1000 Ma age for sapphirine granulite from one locality, other workers have ascribed Early Proterozoic or even Archaean ages to the sapphirine-forming event, and the age given in Tmble 1 should be treated with caution. Poor age control is also a problem for interpretation of the high-T orthopyroxene-sillimanite granulites reported from North Madagascar by Nicollet (1990). Sapphirine occurs in assemblages involving orthopyroxene, spinel, cordierite, corundum and gedrite in silica-undersaturated aluminous gneisses in several Proterozoic granulite terranes. This lower- to medium-P association (i.e. <8-9 kbar), generally indicative of temperatures greatet than 8OO"C, is described in detail from the Arunta Complex (Warren, 1983b), Namaqualand (Waters, 1986) and Madagascar (Nicollet, 1990). Sapphirine coexisting with magnesian garnet and orthopyroxene sillimanite in quartz-deficient assemblages is restricted to higher P-T conditions (Hensen, 1987; Harley and Hensen, 1990), and is correspondingly less common in Proterozoic granulites, being reported only in very localised magnesian aluminous gneisses from the Grenville Province (Herd et al., 1986) and Rauer Group, Antarctica (Harley and Fitzsimons, 1991).
+
+
Proterozoic granulite terranes
327
Independent assemblage evidence on the peak temperatures attending metamorphism of medium-P Proterozoic granulites is now available from calc-silicates. The wollastonite-scapolite-quartz association implies temperatures in excess of 800°C provided that the scapolite is high in meionite content, as is the case in the terranes considered by Harley and Buick (1992) (e.g. Rauer Group, Antarctica). Temperature conditions implied by these and other T-sensitive assemblages including many of the sapphirine parageneses are often higher than those calculated from conventional exchange thermometers but are more reliable because they are less influenced by resetting and closure effects. Peak temperatures based on such assemblages are given in Xible 1 for those terranes or localities where the assemblage evidence is clear.
PRESSURE-TEMPERATUREPATHS Methodology and uncertainties in P-T path determination
The determination of P - T paths is a fundamental prerequisite to interpretations of tectonic settings of granulite metamorphism (Ellis, 1980, 1987; Harley, 1985, 1989; Bohlen, 1987, 1991; Sandiford, 1989). The best documented and most widely accepted P-T paths have been constructed using a range of rock-types and based upon GTB calculations for initial (“peak”) and secondary, generally retrograde, assemblages combined with interpretations of textures and mineral zoning relations in terms of dPldT vectors. Examples of this integrated approach applied to Proterozoic granulite terranes include the Adirondacks (Bohlen et al., 1985), the SW Grenville Province (Anovitz and Essene, 1990) and Namaqualand (Waters, 1986, 1989), and on a smaller scale the Rauer Group, Antarctica (Harley, 1988; Harley and Fitzsimons, 1991) and Uusimaa Complex, Finland (Schreurs and Westra, 1986). As advocated in previous work (Harley, 1989), P-T paths should be deduced using all possible constraints. Although having the apparent virtue of yielding quotable numbers GTB methods are, as noted in the previous section, affected by mismatch and the less obvious problem of feedback. Both of these effects can lead to spurious P - T points, and hence yield point-to-point P-T “paths” which are artifacts of the technique. This problem, discussed at some length in the literature (Harley, 1989; Frost and Chacko, 1989; Perkins, 1990), is avoidable in terranes where a variety of lithologies and assemblages is present. Specific examples of textural interpretation as applied to Proterozoic granulites can be found in the references in Xible 1. The textures most commonly used to define retrograde P-T paths are those in which one phase is pseudomorphed by finer-scale symplectites or where local coronas are developed between earlier coarse phases. Classic examples include garnet coronas between pyroxenes and plagioclase in metabasites and intermediate gneisses, produced by reactions such as:
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+ Plag = Grt + Qtz, and Opx + Plag = Grt + Cpx + Qtz
Opx
These have been used to deduce and quantify P - T paths involving near-isobaric cooling in several Proterozoic higher-P terranes (e.g. Griffin and Heier, 1969; McLelland and Whitney, 1977, 1980; Griffin et al., 1978; Bohlen et al., 1985) and successfully modelled using diffusion kinetics (e.g. Johnson and Carlson, 1990). Pyroxene-plagioclase symplectites on resorbed garnet, usually ascribed to the reactions noted above proceeding in an opposite sense, have been instrumental in defining near-isothermal decompressional P - T histories for several Proterozoic terranes (Ellis, 1983; Harley, 1988; Schumacher et al., 1990; Peloar et al., 1990; Mengel and Rivers, 1991). A few Proterozoic terranes preserve textural evidence for both garnet growth and breakdown as a consequence of more complex P-T histories (Schumacher et al., 1990). Within metapelitic rock types the scope for texturally constrained P-T vectors is greater as the numbers of phases and potential reactions are large. Again, garnet coronas and symplectites after garnet are the most commonly described textures, used in combination with Fe-Mg zoning patterns in reactant phases to deduce the sense of P-T change. For example, orthopyroxene + cordierite symplectites replacing garnet, and cordierite + spinel on garnet previously coexisting with sillimanite, have been used to define decompressional P-T paths for many Proterozoic terranes (Ellis, 1983; Schreurs and Westra, 1986; Santosh, 1987), via reactions such as: Grt
+ Qtz = Opx + Crd,
Grt
+ Sil = Crd + Spl
and
Deductions from such textures are now supplemented by inferences from complex textures involving spinel, cordierite, biotite, garnet, sillimanite and other phases in Fe-rich metapelites (e.g. Vernon et al., 1990) or sapphirine and associated phases in Mg-A1 rich gneisses (Hensen, 1986, 1987). Proterozoic granulites with isobaric cooling histories documented from textures in metapelites and Mg-A1 gneisses include the Arunta Complex (Warren, 1983b), Namaqualand (Waters, 1986), parts of East Antarctica (Clarke et al., 1989) and the Wilson Lake terrane (Currie and Gittins, 1988). Decompressional and more complex P-T histories have been interpreted using contrasting textures developed in similar rock-types in many other Proterozoic terranes (e.g. La1 et al., 1987; Frisch, 1988; Stiiwe and Powell, 1989a; Harley and Fitzsimons, 1991). In many cases the consistency of these texturally determined P-T paths can be tested through consideration of other rock-types in which definitive textures also develop. For example, granulite-facies calc-silicates provide complementary constraints on P-T paths through textures between grossular garnet and the generally higher-T association wollastonite-scapolite (Harley and Buick, 1992).
Proterozoic granulite terranes
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Transitional granulite terranes and prograde P- T trajectories The approaches to P - T path determination described above are in general of limited use in defining prograde P - T histories. However, the dPldT sense of the prograde path is an important constraint on tectonic models as it controls whether granulites are characterised by “clockwise” (England and Richardson, 1977) or “counter-clockwise” (Bohlen, 1987) P - T histories. The former involves first burial and then heating: peak temperature conditions (T,,,) are attained at pressures usually somewhat less than the peak pressures (P,,,) seen on the prograde path. In the latter case heating to peak conditions is synchronous with or more important than burial, implying that heat is supplied rapidly to the site of metamorphism (Loosveld and Etheridge, 1990). This generally results in T,, attained at P,,,,,, and Tmaxapproached from the lower-P direction with a positive prograde dP/dT path. Several Proterozoic granulite terranes show regional scale transitions from lower-grade zones defined principally through mineral isograd mapping. These terranes are significant because lower-grade mineral parageneses may be overprinted by granulite assemblages within the transitional zones. The sense of prograde dPldT can be deduced using these relicts, and inclusions in high-T porphyroblasts. Examples include the Broken Hill region (Binns, 1964; Phillips and Wall, 1981), the NW Adirondacks (Buddington, 1963; D e Waard, 1965; Bohlen et al., 1985; Edwards and Essene, 1988; Fig. 2), Bamble, Norway (Lamb et al., 1986) and Namaqualand (Waters, 19S6). Areally restricted and fine-scale variations characterise the “thermal domes” of Finland (Schreurs and Westra, 1986; Hollta, 1988) and the low-P Anmatjira/Reynolds Range regions of the Arunta Complex (Warren and Stewart, 1988; Vernon et al., 1990), where grades vary from greenschist to granulite over distances of less than 20 km. In the Lutzow-Holm Complex, Antarctica, a continuous transition from amphibolite through upper-amphibolite and into granulite facies is defined by assemblages and isograds in pelitic, calcareous, mafic and ultramafic rock-types (Motoyoshi et al., 1989 and references therein; Hiroi e t al., 1991). In many cases (e.g. Broken Hill, Adirondacks, Reynolds Ranges, Namaqualand, Finland) lower-grade zones have been metamorphosed at similar pressures to the granulites and it is apparent that the prograde history has involved either isobaric heating or heating with only a small increase in pressure (Phillips and Wall, 1981; Waters, 1990). Features considered diagnostic of such prograde histories, and counter-clockwise (CCW) P - T paths, include sillimanite replacing andalusite and cordierite pseudomorphed by spinel-quartz (Waters, 1990). In the Adirondack Highlands further evidence for a pre-granulite low? history is provided by the formation of wollastonite skarns on anorthosite contacts (Valley, 1985; Valley et al., 1990), although this low-P contact metasomatism probably pre-dates the main granulite event by 80 Ma and may therefore be unrelated to it. Not all transitional terranes conform to this counter-clockwise (CCW) model, however. High-P relics and inclusions (e.g. kyanite) implying considerable crustal
S.L. Harley
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t 2'
Low-Pressure Granulites
' 7 d 0 '
'
'
800 '
'
T ("C)
'
'
'
900
'
1
'
3 700
800
900
T ("C)
Fig. 10. Pressure-temperature paths deduced for Proterozoic granulites. Open circles are peak P-T estimates, numbered according to the list in Table 1. Stippled triangular areas bounded by arrowed lines in (a) and (b) indicate typical uncertainties on dP/dT slopes. a. Paths dominated by near-isobaric cooling (IBC). Dashed arrowed lines indicate prograde P-T paths in cases where "counter-clockwise" (CCW) histories have been deduced.
Proterozoic granulite terranes
331
thickening preceding the thermal peak, supporting clockwise P-T paths, are reported from the Lutzow-Holm Complex (Motoyoshi et al., 1989) and Sri Lanka (Schenk, pers. comm., 1991). Eclogitic relics within the SW Grenville Province also testify to significant burial without concomitant heating in that complex granulite region (Davidson, 1990). Those low-P monocyclic granulite terranes from which most of the CCW P-T trajectories have been determined probably represent one special style of granulite development (see below).
Types of retrograde P-T paths Post-peak P - T paths for Proterozoic granulites are depicted in Fig. 10. The paths are characterised in Table 1 using a slope parameter (dPldT) calculated over a given cooling interval (AT). Uncertainties in the cooling interval associated with mineral assemblages of different generations leads to dPldT uncertainties of up to *lo-15 barPC, although these are somewhat less for those paths which are constrained through reaction grids or dataset approaches as well as GTB. Postpeak P-T paths have also been defined based on late fluid inclusions (Santosh, 1987; ?buret and Hartel, 1990; Lamb et al., 1991), a potentially important technique which complements petrologic data. Rather complex P-T histories are described at conditions near the thermal peaks in some terranes (e.g. Hollta, 1988; Stuwe and Powell, 1989b). In general, however, simpler P-T paths characterised by near-isobaric cooling (IBC, low dP/dT) or near-isothermal and adiabatic decompression (ITD, high dP/dT) are prominent, as recognised in previous work (Harley, 1989). There is no correlation between age and P-T path type. IBC paths, here distinguished as those with average dP/dT gradients of less than 15 barPC (Figs. 10a and lla), are particularly common in, but not restricted to, post-1900 Ma granulites. IBC paths are not confined to any particular P or T domain (Fig. 10a). IBC is also documented as a first phase of the retrograde P-T histories of some high-T (>85OoC) terranes such as the polycyclic Wilson Lake terrane (Currie and Gittins, 1988). In three cases (Figs. 10a, 1Oc) some compression has been correlated with cooling, but the amount of such compression is not resolvable given the dPldT uncertainties noted above. ITD retrograde paths vary from essentially adiabatic (>40 barPC) to moderate dPldT, and characterise both low- and high-P terranes (Fig. lob). The extents of high-T decompression are highly variable, with deeper and hotter ITD terranes b. Paths dominated by initial near-isothermal decompression (ITD). Dashed path for 66 relates to a suggested “clochwise” prograde history. c. Composite P-T paths, most of which probably reflect superposed metamorphisms or reworking events. d. P-T paths in low-pressure granulites. The stability of spinel-quartz as given in Fig. 11 after Waters (1991) is shown for reference. Numbers in circles and paths with open arrows: terranes or areas for which IBC and IBC CCW paths have been described. Numbers in hexagons and paths with filled arrows: granulites in which decompression has been important.
+
S.L. Harley
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probably showing greater decompression (AP). A plot of dPldT against cooling interval, T, for all granulites featured in the P - T diagrams (Fig. l l b ) shows that most of the population (80%) define total decompressions of less than 3 kbar, equivalent to 10 km unroofing. There is no clear relationship between peak metamorphic pressure and dPldT (Fig. lla). In particular, significant high-T decompression (i.e. >1.5 kbar) has been documented for several l o w 2 granulite belts including Prydz Bay (Harley and Hensen, 1990), Ellesmere Island (Frisch, 1988) and Kerala (Santosh, 1987). Some low- and medium-P granulites (e.g Larsemann Hills: Stuwe and Powell, 1989a) preserve composite P-T paths with an initial ITD stage, documented by garnet breakdown, followed by IBC and new garnet formation. The preservation of the ITD and composite paths, shown in Fig. 12d along with the IBC paths considered by many workers to be typical of low-P granulite metamorphism,
c5
-
10-
a Y
8-
n
64.
PP
6 lb
I
1
io
3b
4b
50
20
30
40
50
'
!::I 2o01
'I
! 150 i
50
I
I
0
10
dP/dT (bar/'C) Fig. 11.Analysis of theP-T path data in terms of dP/dT slopes (in barPC). a. Demonstration that t h e P T path slope, dP/dT, is independent of pressure and age or regional divisions of Proterozoicgranulites. Symbols as in Fig. 9. IBC: true isobaric cooling; AD: adiabatic decompression. Qpical uncertainties on P and dP/dT indicated by the cross. b. Plot of the cooling intervals over which defined dP/dT slopes have been determined, against dP/dT. Contours defining the extents of unroofing or decompression (in kilobars) show that most of the granulites record decompressions of less than 3 kbar (10 km).
Proterozoic granulite terranes
333
illustrates that low-pressure granulite metamorphism is not always characterised by IBC paths and counter-clockwiseP-T histories. Composite retrograde P-T paths are recognised from several other areas (Fig. lOc). Deep-level IBC from high temperatures to 800-700°C followed by ITD through 2-5 kbar has been described from Wilson Lake (Currie and Gittins, 1988) and the Musgrave Complex (Maboko et al., 1989), and has been suggested for parts of the Highland Complex, Sri Lanka (Schumacher et al., 1990). In the first two examples the terranes are either bound or dissected by thrusts and mylonite zones, and it is feasible that the complex P - T histories have resulted from reactivation of deep-seated crustal segments along these zones. Very distinctive retrograde P-T paths involving significant compression (>2 kbar) without much cooling, and hence a large negative dP/dT gradient, have been proposed for the Bunger Hills, Antarctica, and Strangways Range, Arunta Complex (Stuwe and Powell, 1989b; Norman and Clarke, 1990). Such paths may result from reworking, rather than recording the effects of one metamorphic event. This situation is well-documented in the Adirondack Highlands, where an early low-P contact metamorphic event which produced wollastonite skarns at anorthosite-marble and syenite-marble contacts is distinguished on geochronological (McLelland and Chiarenzelli, 1990) and 0-isotopic (Valley and O'Neil, 1982; Valley, 1985; Cartwright and Valley, 1990) grounds from the later, unrelated, medium-P regional granulite metamorphism. Rates of metamorphism and cooling: P-T-time constraints The geometry of a retrograde P-T path is only a partial characterisation of the post-peak history. Determination of detailed tectonic models for granulites with specific P-T paths inevitably requires complementary age data defining the cooling rate, dTldt (Anovitz and Chase, 1990; Oxburgh, 1990), as this parameter is strongly dependent upon the effective conductive lengthscale of the relaxing thermal system (i.e. lithosphere, Oxburgh, 1990). Such dTldt information is generally developed from data on the apparent ages of minerals, interpreted in terms of different closure temperatures for isotopic exchange in the minerals and systems concerned (e.g. Mezger, 1990). This approach is subject to uncertainty arising from the secondary dependence of closure temperatures on the cooling rate itself, and from the definition of diffusion distances and effective grainsize, but well-constrained cooling curves have now been obtained for a select few granulite terranes using a variety of isotopic methods (e.g. Burton and O'Nions, 1990; Mezger, 1990). Cooling rate data (dT/dt in "C/Ma) for the first 100-300°C of cooling in selected Proterozoic granulites are plotted against the P-T path gradient, dPldT, in Fig. 12. The best constrained are mostly slowly cooled terranes with dT/dt in the range 1-5"ClMa irrespective of the exact shape or slope of the P-T path. The more rapid initial cooling of the Imataca Complex, Venezuela (Swapp and Onstott, 1989) may reflect its uplift and thrust emplacement onto adjacent lower-
S.L. Harley
334
dT/dt (OC/Ma) Fig. 12. P-T path slopes plotted against cooling rates, dT/dt, for those Proterozoic granulites (numbered as in Table 1) where cooling rates are constrained by isotopic data. This diagram is contoured for apparent exhumation rates (mmbear, numbered culves).
grade rocks, while that of the Pielavesi block, Finland, may be related to local magmatic heat sources rather than the regional setting (Hollta, 1988). Slow cooling of the majority of these granulites requires that the sources of the extreme thermal anomalies responsible for the metamorphism are external to the observed terranes, and probably to the crust itself (Oxburgh, 1990; Sandiford and Powell, 1991). As argued by Oxburgh (1990), if only intracrustal melting is considered to contribute to the thermal budget then the cooling rates will be rapid and events short-lived, particularly in the case of low-P granulites. The thermal domes of Finland (Schreurs and Westra, 1986; Hollta, 1988) and Anmatjira Ranges of Australia (Vernon et al., 1990) may be examples of this behavior. Slower cooling over tens of millions of years (Fig. 12) implies thermal relaxation with a thickness lengthscale greater than the crustal or metamorphic thickness. Oxburgh (1990) has modelled this situation in terms of subcrustal lithospheric thinning, which is broadly equivalent to the “asthenospheric perturbation” model put forward by other workers (Stuwe and Powell, 1989a; Clarke et al., 1990), and has demonstrated that the dT/dt data can indeed be successfully explained through thermal relaxation of the whole lithosphere back to its pre-attenuation thickness. In contrast, Anovitz and Chase (1990) have successfully applied a crustal thickening-thinning model (e.g. Harley, 1989; Sandiford, 19S9) to the dP/dT and dTldt data for thrust- and shear-bound tectonic domains within the SW Grenville Province (Fig. lb), without recourse to sub-crustal thinning. Contours of equal exposure or exhumation rate are also plotted on Fig. 12. An average exposure rate of 0.25 mm/yr, considered a maximum for erosion acting alone (England and Thompson, 1984), is exceeded only by the Imataca Complex and perhaps the Musgrave Complex and Tbrngat Orogen. It should be noted, however, that faster exposure rates are possible in many of the ITD granulite terranes if the mineral cooling-age data post-dates the decompression. The initial
Proterozoic granulite terranes
335
post-peak cooling rate is in general not well-constrained by isotopic data, so that most of the exhumation could have occurred early on in the relatively short cooling history associated with ITD. As a consequence, the data presented in Fig. 12 may be biased to slower cooling and exhumation rates simply because of the resolution of the available methods.
ROLES AND SIGNIFICANCE OF FLUIDS AND MELTS
A key factor in the formation and stabilisation of granulite facies high-grade terranes is the desiccation of the lower crustal environment to produce low U H ~ O conditions appropriate to relatively anhydrous assemblages. The major question in this respect is whether the low water activity conditions have been inherited from volatile-absent precursors, produced as a consequence of the extraction of water-bearing partial melts, or imposed through the presence of a water-poor fluid diluted by other volatile components. Evidence from Proterozoic granulites, which have an important bearing on this issue, will be considered in the following section.
Fluids in granulite metamorphism: the case from Proterozoic granulites Low U H ~ O metamorphic conditions (aHzo = 0.03-0.5) have been calculated for many Proterozoic granulites using mineral assemblages in felsic (e.g. biotite-orthopyroxene-kfeldspar-quartz)and pelitic (e.g. biotite-sillimanite-garnetkfeldspar-quartz) rock-types (Tmble 1). However, further evidence bearing on the nature and composition of any fluid phase, if present, has only been obtained in a limited number of terranes. These terranes and their specific features relevant to fluid processes in the granulite facies are given in Tmble 3. A popular school of thought holds that regional granulite metamorphism is often accomplished through the pervasive introduction of COz fluids into the lower crust. Such “carbonic metamorphism” (Newton et al., 1980) has been promoted as an explanation for the in-situ patchy replacement of felsic migmatites by charnockite, or “incipient charnockitisation” in the Achaean and Proterozoic amphibolite-granulite transition zones of southern India (Janardhan et al., 1982; Hansen et al., 1987; Santosh et al., 1990). A key piece of evidence used in support of this fluid-present model is the common observation of C02-rich fluid inclusion populations of appropriately high density in granulite minerals, both on a terrane-wide scale (?buret, 1971, 1985, 1986; Coolen, 1980; Santosh, 1986) and in the incipient charnockites (Hansen et al., 1984; Santosh et al., 1990). For example, Touret (1971, 1985) documented variations in fluid inclusion compositions with grade and a dominance of high-density C02-rich fluid inclusions in the granulite facies zone in the Bamble area, SW Norway. The CO2 fluid was in this example suggested to be derived from the passage of and deepseated crystallisation of mafic magmas, a model elaborated on by Frost and
w w
TABLE 3
Q\
Constraints on fluid processes in Proterozoic granulites Terrane/Area/Locality and Reference Adirondack Lowlands (Powers and Bohlen, 1985; Edwards and Essene, 1988; Valley et al., 1990)
U H ~ O
acoa
Features/Methods
Original interpretation
E l , E2 and fo2controls
fluid absence; low U H ~ Ovia melting; infiltration to give contact skarns
<0.2
El, 12, 14; wollastonite skarns
early contact skams; fluid absent
0.08-0.5 0.01-0.17
Adirondacks: Willsboro (Valley and O’Neill, 1982, 1984; Valley, 1985) Adirondacks Highlands (Lamb and Valley, 1984, 1985)
(0.2
<0.2
E l , E2
fluid absence
Adirondacks Highlands: various (Lamb et al., 1987, 1991) (Valley et al., 1990)
(0.2
<0.2
El, E2 and fo2,Fl,F2,I4 A l , El , E2, E3, F1, I2,14, I5
fluid absence or channelisation as lowfHao andfo, fluid absence or channelisation
<0.3
E l , E2
layer by layer buffering
>0.95 0.75-0.95
c1 C1, C = -7%0
metamorphism in the presence of a COz-rich fluid
C1, F1, F2,€3,F4 wollastonite present
COz at peak, from deep magmas; later resetting of fluid inclusions on retrograde P-T path
Arendal/Ba mbl e (Touret, 1985)
F1, F4
C02 from melts derived from a mantle source
Lofoten, Norway (Baker and Fallick, 1988)
11, large C-0 isotopic shifts
C02 from melts derived from a mantle source
Al, A2, E l , F1, I3
localised infiltration of some COz to give isotopic front; near-isochemical charnockitisation
Adirondacks: Oregon Dome (Lamb and Valley, 1988) Finnish Lapland (Armbruster et al., 1982; Raith and Raase, 1986; Hormann et al., 1990)
%
Uusimaa, Finland (Schreurs and Westra, 1986; Touret and Hartel, 1990)
Kerala: Ponmudi (Hansen et al., 1987; Harris and Bickle, 1989)
0.15 If: 0.14
0.2-0.4
(0.4 0.2-0.35
*!E?2,
9
Kerala general (Jiang et al., 1988) Kerala general (Santosh, 1986) Kerala: Mannantala (Santosh et al., 1990) S r i Lanka: Kurunegala
A l , I6
fluidhock ratio not constrained by
F1, F2, F3
isotopic ratios inclusions trapped along ITD path but
A l , El, F1, F2, F3, 13, I7
0.2-0.3 0.18 f 0.16
(Fiorentini et al., 1990) Rauer Group and Prydz Bay (Harley and Buick, 1992) Furua, Tanzania (Coolen, 1980; Moecher and Essene, 1990)
aB 8
8
3
early COz, so fluid-present? C02 advection from charnockite, infiltration via C isotope shift
2 5
* 2. E
Al, A2, A3, F1
charnockite formed through melting, but some melt lost; no fluid flushing
El, 12, 14, I5
no fluid advection on the large scale
(Burton and O'Nions, 1990) S r i Lanka general
'a
2
a
5
2
0.2-0.4
<0.45
C2, El, E3
fluid absence at peak wollastonite scapoli te
0.2
<0.5
El, E2, E3, F1, F2
no fluid advection
+
Key ro features: A1 = charnockite patches observed; A2 = bulk analysis of charnockite-host pairs; A3 = trace element, REE and radiogenic isotopes compared; C1 = volatile-saturated and CO2-rich cordierite; C2 = volatile-poor cordierite with low XCO,; El = mineral equilibria calculations for U H ~ O or aco,; E2 = C-0-H fluid speciation calculated from simultaneous mineral equilibria; E3 = locally buffering mineral equilibria or unusual implied local fluid X ; F1 = CO2-rich fluid inclusion populations; F2 = heterogeneity in H20-CO2 fluid inclusions; F3 = common late, low density fluid inclusions; F4 = brine inclusions in metapelites; I1 = homogeneous C and 0 isotope compositions in and across layers; I2 = sharp isotopic gradients and discontinuities across layers, heterogeneity in G O ; I3 = smooth asymmetric isotopic gradients or fronts; I4 = retention of pre-metamorphic 0-isotope values (>24%) in calcites; I5 = isotopically light C in graphite (i.e. organic) or cordierite; I6 = 0-isotopes show lithology-related differences; I7 = C-isotopes measured in fluid inclusions.
W W 4
338
S.L. Harley
Frost (1987). The interpretation of fluid inclusion data is, however, far from straightforward. Identification of syn-metamorphic fluid inclusions which have not suffered later modification is difficult and criteria for recognising “primary” inclusions are not always clear (Touret and Hartel, 1990). In addition, in most of the Proterozoic granulites examined a wide range of C02-inclusion densities are reported (Tmble 3), so that inclusion barometry gives highly scattered results when compared to mineral P-T estimates (Lamb, 1990). It is therefore not reasonable that all the C 0 2 inclusions are faithful samples of a peak metamorphic fluid in these cases; C02-rich fluid inclusion populations may be a necessary, but certainly not sufficient, condition for fluid infiltration in granulite formation. The general interpretation that high-density fluid inclusions are representative samples of fluids attending granulite metamorphism has also been questioned because of contradictions between the fluid inclusion chemistry and independent mineral assemblage and fo, constraints bearing on fluid compositions in the Adirondacks. Lamb et al. (1987) documented apparently primary, high-density, C02-rich fluid inclusions in granulite samples which in some cases also contained mineral assemblages restricting both fo, and ace, to low values. In these cases calculatedfco, is too small for a C02-rich fluid to be present. This contradiction implies that the high-density inclusions were, at least in the Adirondack case, entrapped after the peak of metamorphism. This is a feasible process given the isobaric cooling P-T path determined for this terrane (Harley, 1989; Lamb et al., 1991). Lower-density post-peak C02 inclusions would be expected in the many Proterozoic terranes where the retrograde P-T path involved near-isothermal decompression (Uusimaa: Schreurs, 1985; Schreurs and Westra, 1986; Kerala: Santosh, 1986). C02-rich fluid inclusions in samples from incipient charnockite patches, lenses and stringers in the Kerala Khondalites show shifts of up to +3%0 in carbon isotopic compositions compared with those in adjacent host gneisses. Sample traverses indicating that this isotopic shift extends beyond the charnockite/host contact have been modelled using advective front theory to infer that some CO2, derived from an external source, has infiltrated from the charnockite into the host (Harris and Bickle, 1989; Santosh et al., 1990). This isotopic evidence for fluid infiltration does not constrain the amount of fluid flow along the charnockite lens itself but does support the concept of fluid-present granulite formation on a local scale. C - 0 isotopic data suggesting more widespread C 0 2 fluid infiltration under granulite conditions has been documented for the Lofoten-Vesteraalen area by Baker and Fallick (1988), where extensive isotopic shifts to lighter C and 0 (+lo to -7 in 613C and +26 to +10 in @*O) are seen in granulite marbles when compared with amphibolite correlatives. Such extreme coupled C-0 depletions can only be explained by exchange with a large volume of externally derived CO2. The association of isotopically shifted marbles, charnockite/mangerite intrusives and geochemically depleted granulites in the Lofoten region is attributed to a deeper mantle perturbation and its related magmatic activity (Baker and Fallick, 1988).
Proterozoic granulite terranes
339
The volatile composition and content of cordierite has also been used as an indicator of fluid regimes in some Proterozoic granulites. C02-rich (up to 2.3 wt.%) cordierites in the Lapland granulites (Armbruster e t al., 1982) imply equilibrium with near-pure C 0 2 fluid, and their 6I3C value (-7%0) suggests that this fluid was ultimately mantle-derived. On the other hand, Vry et al. (1988) report low C02-H20 contents and isotopically light C (-26 to -15%0) in cordierites from some Proterozoic terranes (e.g. Kerala, Grenville), which does not support this as a general model and indeed restricts the amount of any infiltrating C 0 2 to very low fluid-rock weight ratios (<0.01; Vry et al., 1988, 1990). Uniformly low total volatile contents and light S13C values (-15 to -10%0) are also recorded for cordierites in migmatitic metapelites from the Prydz Bay region of East Antarctica, indicating a negligible influx of mantle-derived C 0 2 during an inferred partial-melting event producing cordierite (Harley et al., in Prep. ). Evidence such as the presence of distinctive low-variance, fluid-buffering mineral assemblages, and heterogeneity in the stable isotope composition of minerals in adjacent lithologies, supports the view (e.g. Valley et al., 1983, 1990; Waters and Whales, 1984; Valley and O’Neil, 1984) that granulite facies metamorphism has progressed under vapour-absent conditions in many Proterozoic terranes. Local-scale variations in U H ~ Owhich do not relate to variations in grade have been calculated using metapelite assemblages in the Adirondacks (Edwards and Essene, 1988; Valley et al., 1990) and Namaqua Province (Waters and Whales, 1984), among others. Assemblages which constrain both aHZO and acoZ,or finely interlayered assemblages restricting either of these variables, have been shown to either imply local buffering or general fluid absence (Lamb and Valley, 1985; Valley et al., 1990; Moecher and Essene, 1990). Mixed volatile equilibria in Adirondack marbles, for example, imply sharp gradients in the compositions of any fluids, and the close spatial association of wollastonite-calcite-quartz assemblages buffering acoZ to low values and biotite-quartz-Kfeldspar-orthopyroxene ones buffering U H ~ Oto <0.1 suggest fluid absence. General fluid absence is inferred for granulite metamorphism in the Rauer Group, Antarctica, where wollastonite-scapolite assemblages implying acOz < 0.4 occur within 15 mm of orthopyroxene-Kfeldspar-quartzlayers restricting aHZO< 0.3 (Harley and Buick, 1992). Wollastonite may be viewed as an indicator of the absence of a free fluid phase in granulite metamorphism in many Proterozoic terranes (e.g. Valley et al., 1990), although it is recognised that channelised and transitory fluid access may occur and is not precluded by heterogeneous aHZ0,acOz and fo, data. Although fluid advection has been invoked in some local-scale isotopic studies of Proterozoic granulites (Santosh et al., 1990), detailed traverses across lithological contacts in the Adirondacks (e.g. Valley and O’Neil, 1984; Cartwright and Valley, 1990) and Sri Lanka (Fiorentini et al., 1990) have shown that pervasive fluid fluxing is not consistent with measured oxygen (and sometimes carbon) isotopic heterogeneities between interlayered granulites. Pre-metamorphic stable isotope signatures and contrasts may be very well preserved in some cases. For
340
S.L. Harley
example, sharp discontinuities in O-isotope composition across wollastonite skarnmarble contacts near anorthosites and syenites in the Adirondacks are interpreted to represent contact metasomatic fronts little altered or readjusted in the later, fluid-absent, regional granulite facies metamorphism (Valley, 1985; Cartwright and Valley, 1990). Furthermore, in both the Adirondack and Sri Lanka case studies calcites with sedimentary O-isotope values (> +25%0) are preserved in granulite facies marbles and graphites with very light C-isotope values are also recorded, arguing against the pervasive infiltration of large quantities of externally derived fluid. The weight of combined phase equilibrium, isotopic and fluid compositional data from the most comprehensively studied Proterozoic granulite terrane, the Adirondacks, point to fluid absence under peak metamorphic conditions. This conclusion is consistent with the evidence from other Proterozoic granulites from Antarctica and Sri Lanka, but it is clear that more terranes must be studied using a multi-method approach before the role of C02 infiltration as a general cause of regional granulite metamorphism can be clarified or discounted. Migmatites and the role of partial melting in Proterozoic granulites
Melting in the granulite facies has long been promoted as a viable mechanism for the depletion of high-grade terranes in incompatible elements and simultaneous reduction of water activity consequent upon melt extraction prior to cooling (e.g. Fyfe, 1973; Thompson, 1982). Indeed, melting is probable in metapelitic rocks even under fluid-absent conditions, as recent experiments (Vielzeuf and Holloway, 1988; LeBreton and Thompson, 1988) have shown that dehydration melting of natural biotites (Bt) of intermediate Mg/(Mg + Fe) (X,,) begins at 700-800°C and may produce large volumes of melt at 8504375°C. In contrast to pervasive fluid “flushing” models for granulite metamorphism, partial melting and melt extraction can be envisaged as processes operating on a variety of scales and independently in different lithologies (e.g. Grant, 1985; Vielzeuf and Holloway, 1988; Waters, 1988), inducing local water activity gradients and variations in mineral assemblages (Waters and Whales, 1984). Migmatite structures have been extensively described in granulite metapelites and intermediate to felsic gneisses in many Proterozoic granulite terranes (e.g. Halden et al., 1982; Hopgood, 1984; Park and Dash, 1984) and are clearly important field features in many more. Detailed structural, textural and petrological studies of metapelitic granulites from several Proterozoic terranes including the Namaqua Province (Waters, 1988), Willyama Complex (Powell and Downes, 1990), Arunta Complex (Vernon and Collins, 19%; Vernon et al., 1990), Yaounde (Barbey e t al., 1990) and east Antarctica (Stuwe and Powell, 1989c) testify to the importance of local dehydration (vapour-absent)-melting reactions in the prograde evolution of the metapelitic assemblages. Some examples of structural and textural features in Proterozoic granulites considered to reflect the formation, passage, and crystallisation of partial melts are given in Fig. 13.
Proterozoic granulite terranes
341
Several approaches have been used to define the melting reactions and processes in garnet-bearing metapelites: (a) geochemical analysis and bulk comparisons of adjacent leucosomes, mesosomes and potential restites (Barbey et al., 1990); (b) mapping of discordant vein relations; (c) definition of interstitial-xenomorphic and phenocrystic textural relationships between high-T phases in leucosomes and adjacent hosts (e.g. Vernon and Collins, 1988; Vernon et al., 1990); (d) mass or phase balance calculations wherein the modal proportions of solid products of proposed peritectic melting reactions are compared with calculated abundances (Waters, 1988). In most cases, the migmatites have been interpreted using semi-quantitative or qualitative petrogenetic grids (Grant, 1985) developed around the reaction: Bt
+ Sil + Qtz = Grt + Crd + Kfs + Liquid,
or in terms of multivariant melting equilibria:
+ Sil + Qtz = Grt + Kfs + Liquid, Bt + Sil + Qtz = Crd + Kfs + Liquid
Bt
or
For example, Waters and Whales (1984) interpreted interlayered cordieritegarnet-Kfeldspar-quartz and biotite-sillimanite-quartz assemblages in terms of divariant dehydration melting, with the biotite-bearing assemblage stabilised a t a higher U H ~ Othan the garnet-cordierite one. Powell and Downes (1990) have considered garnet-Kfeldspar-quartz segregations in the Willyama Complex and interpreted their mantled morphology in terms of peritectic melting via the garnet-forming reactions above, followed by melt crystallisation over a dispersed area outside of the segregation. This has a bearing on the major problem for dehydration-melting as a mechanism for stabilisation of granulite assemblages, namely the preservation, without retrogression, of anhydrous assemblages formed in the melting interval. In an idealised closed system these should react with the hydrous melt to re-form hydrated phases (Waters, 1988). Indeed, in the case of the Arunta Complex in the Strangways Range region, Warren and Hensen (1989) have argued that a regional retrograde biotite-forming event (regional hydration) resulted from the back-reaction of partial melts with their host granulites subsequent to limited segregation. The general observation that segregations and leucosomes have compositions which do not correspond to those of minimum melts (e.g. Waters, 1988; Powell and Downes, 1990) is consistent with the loss of at least some final melt fraction, and vapour, from the site of earlier partial melting, which would prevent back-reaction of the prograde assemblages and result in assemblage-related aI120 gradients. A n important aspect of petrogenetic studies of melting in Proterozoic granulites up to the present is that most of the case studies are drawn from (k spinel, biotite) the low-pressure terranes: cordierite-garnet-sillimanite-quartz
342
S.L. Harley
Proterozoic granulite teiranes
343
migmatites appear to be typical rock-types in most of these. However, melting in the granulite facies is not restricted to these low-P terranes. Textural and geochemical features attributed to partial melting have been described from the higher-pressure kyanite-bearing migmatitic metapelites of Yaounde (Barbey et al., 1990), and orthopyroxene-feldspar segregations in magnesian metapelites from the Rauer Group are considered to reflect melting at 10 kbar by Harley and Fitzsimons (1991). Furthermore, while most studies have concentrated on migmatisation in metapelites, it should be noted that net-vein and agmatite structures, orthopyroxene-phyric leucosomes and segregations, and in a few cases detailed geochemistry and modal melt modelling, indicates that partial melting is also important in the evolution of Proterozoic medium- and high-pressure mafic granulites ( B i t and Harley, 1988; Pattison, 1991). Lastly, a dehydration-melting model has been proposed for the genesis of patch charnockites from Kurunegala, Sri Lanka (Thble 3 ) . Burton and O’Nions (1990) interpret the charnockite as a product of biotite-amphibole melting in the surrounding gneisses, rather than a result dehydration due to C 0 2 streaming. The charnockite is not considered to represent the actual melt composition, but is inferred from its systematic relative depletion in incompatible elements and Sm, Rb and Pb to be partly residual. The incompatible element-enriched melt fraction absent from the charnockite patch has instead been distributed and reacted in the selvedge or transitional domain between the charnockite and host. This type of charnockite development, however, is not comparable with the higher-temperature formation of relatively dry charnockitic magmas, either related to anorthosite emplacement (e.g. McLelland, 1989; Emslie and Hunt, 1989) or emplaced in differentiated intrusive suites in the mid- and lower-crust (Young and Ellis, 1991).
Fig. 13. Some migmatitic features in Proterozoic granulites. a. Patch leucosome segregation in migmatitic mafic granulite, Rauer Group, East Antarctica. The leucosome contains coarse euhedra of orthopyroxene (dark phase) and rare garnet, set in plagioclase and minor quartz; the mesosome contains amphibole, clinopyroxene, orthopyroxene and plagioclase. The coarseness of the phases, lack of continuity between patches on a larger scale, local loss of hornblende, and restriction of such patches to mafic units a r e consistent with local partial melting. The segregation is considered to include most of the solid and liquid products of the melting reaction, modified by the loss of some melt and by later subsolidus annealing. b. Layered low-P granulite-facies paragneisses, Reynolds Range, Arunta Complex, Australia. The coarser, migmatitic, layers a r e pelitic in composition and contain cordierite, sillimanite, quartz, K-feldspar and minor biotite; the liner layers a r e psammitic to seniipelitic and contain cordierite and biotite in addition to quartz. Coarse layers, although usually concordant with the psammitic layers, locally grade into discordant veins and patches and are considered to have undergone pervasive melting whereas the psammitic layers did not. Biotite selvedges occur on the edges of the psammitic layers. This type of migmatite is similar t o that described from Namaqualand by Waters and Whales (1984). c, d. Examples of discordant, cordierite-bearing leucosome veins cutting cordieritesillimanite-biotite metapelites, Reynolds Range, Arunta Complex. T h e high concentration of euhedral to subhedral cordierite in the veins cannot be produced by crystallisation of a reasonable melt composition (e.g. Vielzeuf and Holloway, 1988). These cordierites, which have low total volatile contents of less than 0.9 wt.% and X H ~ in O the range 0.64-0.7 as measured by SIMS (Harley, unpubl. results) a r e instead considered to mainly be the solid products of dehydration-melting reactions.
344
S.L.Harley
SOME REMARKS ON TECTONIC MODELS FOR PROTEROZOIC GRANULITES
The preceding sections document and summarise a remarkable diversity and range in structural characteristics, conditions of metamorphism, pressuretemperature paths and cooling histories, and fluid/melt relationships in Proterozoic granulites. This diversity in the major attributes of granulites cautions against applying a single simple model for their formation, as discussed in detail by Harley (1989). In this concluding section I review the recent developments and trends in tectonic modelling as applied to the interpretation of Proterozoic granulites. Structural constraints described above indicate that many large-scale Proterozoic granulite terranes have evolved within a broadly collisional context. For example, Rivers et al. (1989) suggest a diachronous collisional event, or oblique collision, involving microcontinental slabs and late closure between larger crustal units to explain the structural features and chronology of events in the Grenville Province of Labrador and Ontario. However, it is now clear that tectonic models involving collision with homogeneous thickening of the lithosphere are not capable of producing the high temperatures typical of most Proterozoic granulites for reasonable values of critical model parameters including thermal conductivity, crustal heat productivity, and basal heat flux (England and Thompson, 1984, 1986; Bohlen, 1987, 1991; Harley, 1989). This result is particularly striking in the case of the low-pressure/high-temperature terranes, where the extreme syn-metamorphic thermal gradients require a large-scale heat source or “thermal perturbation”. Conductive relaxation models also, in general, fail to produce dispersed and steep P-T paths such as the common ITD types described herein (Harley, 1989; Anovitz and Chase, 1990), and they cannot account for the counter-clockwise (CCW) P-T paths considered to be typical of many low-pressure and some medium-pressure Proterozoic granulites (Bohlen, 1987, 1991). A further weakness in the simple collision model is that the thermal peak and main granulite metamorphism postdates the major compressive deformation by a large time interval (e.g. 60-80 Ma), whereas granulite metamorphism is constrained to be syn- to post-deformational in many examples, particularly the low-pressure belts (Loosveld and Etheridge, 1990). Higher-pressure granulites such as the Zambesi Belt (Treloar e t al., 1990) and the Grenville Province in SE Canada (Rivers et al., 1989; Anovitz and Chase, 1990) show considerable overlap in time between high grade metamorphism, magmatism, and compressional or extensional deformations. It is, therefore, now generally accepted that regional granulite metamorphism requires an input of extra heat over and above that available from a thickened crust-lithosphere system, and that in many cases the thermal input has been coeval with major compressive, or in some cases extensional (e.g. Harley, 1989; Sandiford, 1989; Barbey and Raith, 1990) deformation. Bohlen (1987, 1991) has argued for magmatic accretion as the essential cause of granulite metamorphism in those terranes characterised by counter-clockwise, isobaric cooling (CCW-IBC) P-T paths. In this model, the high heat input is contributed by underplated basic magmas, while the compression results from loading by extrusives and granites a t
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shallower crustal levels. This process could be effective in magmatic arcs or in collapsed continental rifts where thinned crust is thickened by thrusting combined with magmatism (Waters, 1990). However, magmatic accretion will only effect large-scale granulite metamorphism if the volume of accreted magma is similar to or greater than the volume of previous crust, i.e. a major period of crustal growth, as perhaps seen in Lofoten (Griffin e t al., 1978) is required (Wells, 1980; Oxburgh, 1990). Clearly, magmatic accretion is not a viable mechanism for granulite metamorphism in terranes dominated by the reworking of older crust (Harley, 1987, 1989; Mengel and Rivers, 1991). A further implication of the magmatic accretion model is that the heating/ cooling cycle, and hence timescale of metamorphism, is relatively short, 10 to 30 Ma, compared with cooling rates deduced for many granulite areas (Loosveld and Etheridge, 1990; Oxburgh, 1990). Slow cooling requires that the heat source undoubtedly needed to produce most granulites must lie external to the crust undergoing metamorphism. In a few cases where rapid grade variations are observed and relatively rapid cooling rates calculated, as in the “thermal dome” situation (Schreurs and Westra, 1986; Hollta, 1988; Vernon et al., 1990), the case for magmatic accretion as the principal cause of granulite metamorphism may be plausible. In most regional granulites, however, the high-grade metamorphism and magmatism are probably both consequences of lithospheric-scale thermal processes. Similar cooling-rate arguments have been used to discount simple crustal extensional models for low-pressure granulites (Loosveld and Etheridge, 1990). In the case of homogeneous extension, the extension factors required to produce even upper amphibolite facies rocks in a typical crust are very high and as a consequence the thermal pulse is short-lived and decays much more rapidly than seen in real cases. In addition, the high-temperature metamorphism will be synchronous with extensional deformation. There has recently been a convergence of opinion and modelling on the formation of granulites, and in particular the low-pressure terranes characterised by CCW-IBC paths. The earlier approaches, which assumed a homogeneous distribution of strain throughout a deforming lithosphere have been superseded by models in which decoupling occurs between the crust and lithospheric mantle undergoing a thickening deformation. It is now considered that lithospheric thickening need not accompany crustal thickening, and that lithospheric thinning might even be possible during crustal compression. With this extra degree of freedom it is possible to duplicate most of the observed P-T paths and obtain appropriately slow cooling rates. The principal controls on thermal histories are the relative extents of crustal and lithospheric thickening and the timing of any lithospheric thinning relative to the onset of crustal thickening. The range of possibilities and thermal/mechanical consequences of decoupling the responses of crust and mantle to compression are elegantly portrayed by Sandiford and Powell (1990, 1991). The first case of importance for granulite formation in the Proterozoic is that in which both the crust and lithosphere thicken initially, but then the sub-crustal
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lithosphere thins while the crust is still thickening. This “type 2” deformation path of Sandiford and Powell (1990) develops from and generalises previous tectonic models relating granulite metamorphism to the extensional collapse of thickened crust (Ellis, 1987; Harley, 1988, 1989; Sandiford, 1989; Anovitz and Chase, 1990), triggered by detachment of the lithospheric thermal boundary layer (Houseman et al., 1981; Sonder, 1987). In this “Tibetian” tectonic style the compressional event leading to the development of an overthick crust is overprinted by a thermal event caused by the effective thinning of the lithosphere from below. Granulite metamorphism will therefore post-date compressional deformation but may be pre- to syn-extensional and correlate with collapse of the orogen (Harley, 1989; Sandiford, 1989). Crustal thinning and high-temperature decompression will in this case lead to syn- or post-extensional ITD paths, post-extensional IBC paths and potentially to P-T paths which are initially dominated by decompression and later by cooling (Ellis, 1987; Harley, 1989). This type of orogenic collapse model for granulite development has been applied in some detail to the SW Grenville Province (Anovitz and Chase, 1990), where P-T paths are characterised by combined decompression and cooling, slow cooling rates prevail (1-2”C/Ma) and late extensional structures are thought to occur. The model may also be applicable to other terranes showing high-temperature ITD paths (e.g. Harley, 19S8), and to those preserving IBC histories, if a “clockwise” P-T evolution can be demonstrated and an extensional origin for the high-grade deformation structures established. Crust-mantle behavior involving essentially total decoupling provides the basis for recent models developed to explain the low-pressure granulites and their broadly syn-deformational thermal peaks. In these models the lithosphere only thickens marginally, if at all, during crustal thickening (Loosveld and Etheridge, 1990; Sandiford and Powell, 1991), or the lithospheric mantle thins relative to the crust (Oxburgh, 1990). In all cases the absolute amount of crustal thickening will be limited by the reduced strength of the crust as it is rapidly heated by the syn-thickening mantle thermal perturbation (Sandiford and Powell, 1991), and crustal doubling will not occur. The metamorphic consequences of coeval crustal thickening and advective thinning of the subjacent mantle include low-pressure granulite metamorphism, counter-clockwise (CCW) P-T paths resulting from the early delivery of a significant thermal pulse, and slow (1-3”C/Ma) post-peak cooling as this is controlled by the decay of the whole lithosphere back to a steady-state thermal structure (Oxburgh, 1990; Loosveld and Etheridge, 1990). The recent models briefly described above, and those noted by Harley (1989), are appealing for the origin of many granulite terranes and provide a framework for further studies. The versatility of the models is a strength but it is also a weakness in that we do not have sufficient critical data such as cooling rates, sense of shear indicators and timing relationships in most terranes to evaluate the details of the thermal model. The geological circumstances in which large-scale thickening followed by collapse occurs and those associated with limited thickening and coeval lithospheric thinning are not resolved.
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In conclusion, Proterozoic granulite terranes are many and varied. No one model can explain all the features and variations, although it is apparent that compression, and probably collision, has often played an important role in their development. The most recent qualitative and quantitative models recognise the need for a major input of heat external to the crust, and as a consequence allow the sub-crustal lithosphere to thin, either in absolute terms or relative to the crust, at some stage in the orogenic evolution. We do not yet understand the reasons for these proposed variations in the behavior of the crust-lithosphere system.
ACKNOWLEDGEMENTS
I thank Kent Condie for his invitation to write this chapter, and for his subsequent patience and forebearance with a late and verbose contributor. Discussions with Ian Fitzsimons, Christian NicolIet, Volker Schenk, M. Santosh and John Valley are greatly appreciated. This is also a contribution to IGCP project 304 “Lower Crustal Processes”.
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Valley, J.W. and O’Neil, J.R., 1982. Oxygen isotope evidence for shallow emplacement of Adirondack anorthosite. Nature, 300: 497-500. Valley, J.W. and O’Neil, J.R., 1984. Fluid heterogeneity during granulite facies metamorphism in the Adirondacks: stable isotope evidence. Contrib. Mineral. Petrol., 85: 158-173. Valley, J.W., McLelland, J.M., Essene, E.J. and Lamb, W.M., 1983. Metamorphic fluids in the deep crust: evidence from the Adirondacks. Nature, 301: 226-228. Valley, J.W., Bohlen, S.R., Essene, E.J. and Lamb, W.J., 1990. Metamorphism in the Adirondacks, 11. The role of fluids. J. Petrol., 31: 555-596. Van Kranendonk, M.J., 1990. Structural history and geotectonic evolution of the eastern Torngat orogen in the North River map area, Labrador. Current Research, Part C. Geol. Sum. Can., Pap., 90-1C 81-96. Vernon, R.H. and Collins, W.J., 1988. Igneous microstructures in migmatites. Geology, 16: 1126-1129. Vernon, R.H., Clarke, G.L. and Collins, W.J., 1990. Local, mid-crustal granulite facies metamorphism and melting: an example in the Mount Stafford area, central Australia. In: J.R. Ashworth and M. Brown (Editors), High-temperature Metamorphism and Crustal Anatexis. Unwin Hyman, London, pp. 272-319. Vielzeuf, D. and Holloway, J.R., 1988. Experimental determination of the fluid-absent melting relations in the pelitic system. Contrib. Mineral. Petrol., 98: 257-276. Visser, D. and Senior, A., 1990. Aluminous reaction textures in orthoamphibole-bearing rocks: the pressure-temperature evolution of the high-grade Proterozoic of the Bamble sector, south Norway. J. Metamorph. Geol., 8 231-246. Vry, J.K., Brown, P.E., Valley, J.W. and Morrison, J., 1988. Constraints on granulite genesis from carbon isotope compositions of cordierite and graphite. Nature, 332 66-68. Vry, J.K., Brown, P.E. and Valley, J.W., 1990. Cordierite volatile content and the role of COz in highgrade metamorphism. Am. Mineral., 75: 71-88. Warren, R.G., 1983a. Metamorphic and tectonic evolution of granulites, Arunta Block, central Australia. Nature, 305: 300-303. Warren, R.G., 1983b. Prograde and retrograde sapphirine i n metamorphic rocks of the central Arunta Block, central Australia. BMR J. Aust. Geol. Geophys., 8: 139-145. Warren, R.G. and Hensen, B.J., 1989. The P-T evolution of the Proterozoic Arunta Block, central Australia, and implications for tectonic evolution. In: J.S. Daly, R.A. Cliff and B.W.D. Yardley (Editors), Evolution of Metamorphic Belts. Geol. Soc. London, Spec. Publ., 43: 349-355. Warren, R.G. and Stewart, A.J., 1988. Isobaric cooling of Proterozoic high-temperature metamorphites in the northern Arunta Block, central Australia: implications for tectonic evolution. Precambrian Res., 40/41: 175-198. Warren, R.G., Hensen, B.J. and Ryburn, R.J., 1987. Wollastonite and scapolite in Precambrian calcsilicate granulites from Australia and Antarctica. J. Metamorph. Geol., 5: 213-223. Waters, D.J., 1986. Metamorphic history of sapphirine-bearing and related magnesian gneisses from Namaqualand, South Africa. J. Petrol., 27: 541-565. Waters, D.J., 1988. Partial melting and the formation of granulite facies assemblages in Namaqualand, South Africa. J. Metamorph. Geol., 6 387-404. Waters, D.J., 1989. Metamorphic evidence for the heating and cooling path of Namaqualand granulites. In: J.S. Daly, R.A. Cliff and B.W.D. Yardley (Editors), Evolution of Metamorphic Belts. Geol. Soc. London, Spec. Publ., 43: 357-363. Waters, D.J., 1990. Thermal history and tectonic setting of the Namaqualand granulites, Southern Africa: clues to Proterozoic crustal development. In: D. Vielzeuf and Ph. Vidal (Editors), Granulites and Crustal Evolution. Kluwer, Dordrecht, pp. 243-256.
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Waters, D.J., 1991. Hercynite-quartz granulites: phase relations, and implications for crustal processes. Eur. J. Min., 3 367-386. Waters, D.J. and Whales, C.L., 1984. Dehydration melting and the granulite transition in metapelites from southern Namaqualand, S. Africa. Contrib. Mineral. Petrol., 88 269-275. Wells, P.R.A., 1980. Thermal models for the magmatic accretion and subsequent metamorphism of continental crust. Earth Planet. Sci. Lett., 46: 253-265. Wiener, R.W., McLelland, J.M., Isachsen, Y.W. and Hall, L.M., 1984. Stratigraphy and structural geology, of the Adirondack mountains, New York review and synthesis. In: M.J. Bartholomew (Editor), The Grenville Event in the Appalachians and Related Topics. Geol. SOC.Am., Spec. Pap., 194: 1-55. Wilde, S.A. and Murphy, D.M.K., 1990. The nature and origin of late Proterozoic high-grade gneisses of the Leeuwin Block, Western Australia. Precambrian Res., 47: 251-270. Wood, B.J., 1975. The influence of pressure, temperature and bulk composition on the appearance of garnet in orthogneisses - an example from South Harris, Scotland. Earth Planet. Sci. Lett., 26: 299-311. Wood, B.J., 1977. The activities of components in clinopyroxene and garnet solid solutions and their application to rocks. Phil. Trans. R. SOC.London, 286A: 331-342. Young, D.N. and Ellis, D.J., 1991. The intrusive Mawson Charnockites: evidence for a compressional plate margin setting of the Proterozoic mobile belt of East Antarctica. In: M.R.A. Thomson, J.A. Crame and J.W. Thomson (Editors), Geological Evolution of Antarctica. Cambridge University Press, pp. 25-31. Young, E.D., 1989. Petrology of biotite-cordierite-garnet gneiss of the McCullough Range, Nevada, 11. P - T - u H , path ~ and growth of cordierite during late stages of low-P granulite-grade metamorphism. J. Petrol., 3 0 61-78. Young, E.D., Anderson, J.L., Clarke, H.S. and Thomas, W.M., 1989. Petrology of biotite-cordieritegarnet gneiss of the McCullough Range, Nevada, I. Evidence for Proterozoic low-pressure fluidabsent granulite-grade metamorphism in the southern Cordillera. J. Petrol., 3 0 39-60.
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Chapter 9
XENOLITHS IN PROTEROZOIC CRUST: EVIDENCE FOR REWORKING OF THE LOWER CRUST J. RUIZ
INTRODUCTION
Geochemical studies of Precambrian crustal rocks indicate that there was a significant increase in the growth rate of the continental crust during the Proterozoic (e.g. Nelson and DePaolo, 1985; Patchett and Amdt, 1986), and that rocks of Proterozoic age form the major portion of the present-day continental crust. Although much of the Proterozoic crust is buried by Phanerozoic rock sequences, there are large rock exposures of Proterozoic crust in shield areas. The deeper portions of Proterozoic crustal blocks can be investigated by studying crustal xenoliths entrained in mantle-derived rocks. These xenoliths are important because they may yield information about the tectonic and magmatic history of the deeper portions of Proterozoic crustal blocks. Granulite-facies xenoliths that equilibrated at pressures and temperatures consistent with those of the lower crust are found in alkalic basalts throughout the world (Fig. 1). These granulite xenoliths are accidental fragments of the crust entrained by mafic eruptions from the mantle. The xenolith suites typically contain a variety of crustal and mantle lithologies. Although the mantle xenoliths may not be representative of the mantle they sample because the specific processes that caused the xenolith-bearing magmatic event is unusual, the crustal xenoliths are probably representative of the types of rocks that are found below xenolith-bearing localities because they are accidental fragments plucked from the lower crust by the quickly ascending mantle-derived magmas. Many xenoliths are thought to represent samples of present-day lowermost continental crust (e.g. Griffin and O’Reilley, 1987a, b; Roberts and Ruiz, 1989; Rudnick, 1992), but low grade and some granulite-facies xenoliths probably represent middle parts or possibly upper parts of the continental crust sampled by the erupting magmas (e.g. Griffin and O’Reilly, 1986, 1987b). This paper will concentrate on xenoliths that because of geobarometric studies (e.g. Bohlen and Metzger, 1989), mineralogical and textural characteristics, are thought to be from the deepest parts of the continental crust. Figure 1 shows the location of xenolith districts that have been reported. Of these localities, 34% are on Phanerozoic crust, 52% erupted through Proterozoic crust and 14% through Archean crust. An excellent review of the geochemical and
362
J. Ruiz
Fig. 1. Proterozoic and Phanerozoic xenolith-bearing localities worldwide. A summary of the types of xenoliths and work done in each locality is given in Table 1. Abbreviations: A = Anakies, Victoria, BC = Boomi Creek, N.S.W., CH = Chudliegh Province, Queensland, CL = Chinese Peaks, California, Camp Creek, Arizona and Geronimo Volcanic Field, Arizona, CP = Colorado Plateau, CT = Calcutteroo, South Australia, D = Delegate, E = Eiffel, Germany, EA = El Alamein, South Australia, H = Hoggar, Algeria, I = Ichinomegata, Japan, KH = Kilbourne Hole, New Mexico, LA = Lashine, Tanzania, L E = Lesotho, South Africa, LK = Lace Kimberlite, South Africa, L O = La Olivina, Mexico, W = La Joya Honda and La Joyuela, Mexico, M = Mount Carmel, Israel, MB = McBride Province, Queensland, MC = Massif Central, France, ME = Mt. Erebus, Antarctica, MS = Man Shiel, Liberia, MV = Midland Valley, Scotland, N = Nunivak Island, Alaska, PA = Pali Aike, Chile, PH = Popes Harbour, Nova Scotia, SR = Snake River Plains, Idaho, U = Udachnaya, Siberia.
some physical characteristics of crustal xenoliths has recently been published by Rudnick (1992), and a very complete comparison of the geochemistry of exposed granulite-facies terranes and granulite xenoliths has been made by Rudnick and Presper (1990). A comparison of the pressures of equilibration of granulite xenoliths and exposed granulite-facies terranes has been made by Bohlen and Metzger (1989). Readers are referred to these publications for descriptions of xenoliths of Archean or Phanerozoic age and for comparisons in the geochemistry of granulite-facies xenoliths and exposed granulite terranes. The purpose of this chapter is to review some of the most important geochemical and lithological aspects of xenoliths erupted through Proterozoic crust.
LOWER CRUSTAL XENOLITH DESCRIPTIONS AND LOCALITIES
Granulite-facies xenoliths are generally found in mildly alkalic to alkalic basaltic tuffs and lavas in regions that are typically under extensional tectonic stresses. Specific tectonic settings of the basaltic magmatism have been summarized by Nixon (1987). The extensional regime may be very broad, such as the Basin and Range (e.g. Camp Creek, Arizona), or continental rifts, such as the Rio Grande Rift (e.g. Kilbourne Hole, New Mexico). It is significant that volcanism in most xenolith-bearing localities is extensive and that the crust that is being
Xenoliths in Proterozoic crust: evidence for reworking of the lower crust
363
sampled by the eruptives may have been severely changed by the volcanism prior to the xenolith-bearing eruptions. This possibility has been proposed by Johnson et al. (1990) based on the geoche'mistry of volcanic rocks erupted from caldera complexes. Lower crustal xenoliths younger than the upper crustal rocks have been documented by ion microprobe dating of zircons in some localities (Rudnick and Williams, 1989; Rudnick and Cameron, 1991). Shorter-lived magmatic events that are less likely to have affected the lower crust and that also collect xenoliths are kimberlites, lamprophyres and other highly alkalic magmas mostly found in cratons. One of the most serious problems with xenolith studies is making sure that the xenoliths are from the lower crust. Obviously an important feature xenoliths must have is equilibration pressures and temperatures that are consistent with those of the lower crust. In most cases these pressures and temperatures will produce granulite-facies rocks but if the crust is thin enough, the xenoliths may be amphibolite grade. This metamorphic grade criterium, however, can only be used in conjunction with the existence of mantle-derived rocks in the same suite of xenoliths. If no mantle rocks exist in the suite of xenoliths, the crustal samples could have been picked up from granulite- facies rocks tectonically emplaced a t higher levels in the crust (e.g. Leeman et al., 1985). Decompression features, such as preferential melting of mafic phases (Padovani and Carter, 1977b) and kelyphite formation on rims and in cracks in garnets (Garvey and Robinson, 1984) are interpreted to have resulted from rapid ascent of the xenoliths from deep crustal levels. Mineral ages of the xenoliths may also be a useful test of the source of the xenoliths. If the samples resided in the lower crust at the time of sampling by the basalt host, then xenolith mineral isochrons could give the age of xenolith emplacement because the xenoliths would have been at P-T's capable of resetting the isotope systems (e.g. Van Breemen and Hawkesworth, 1980; Padovani and Hart, 1981; Cliff, 1985). The isotopic systems of Rb/Sr and Sm/Nd may be useful (Reid et al., 1989), although the data are generally difficult to interpret. Because zircons can survive lower crustal conditions without being completely reset, the U-Pb system may be the best way of dating the lower crust (Rudnick and Williams, 1987; Rudnick and Cameron, 1991). In this regard it is interesting to note that although xenolith localities are in crustal blocks of Proterozoic age ( a b l e 1; Fig. l ) , recent studies (e.g. Rudnick et al., 1986; Kempton et al., 1990; Downes et al., 1991; Ruiz et al., 1991) argue that the lower crust is younger. The mineralogy of the lower crustal xenoliths is a function of bulk composition, Auid activity and pressure and temperature of equilibration. Rudnick (1992) has compiled a very extensive list of mineralogical characteristics of xenoliths with worldwide coverage. Table 1 shows the mineralogy of xenoliths erupted through Proterozoic crust. The most common xenoliths are mafic, and their mineralogies are dominated by plagioclase, pyroxene and garnet. Olivine, spinel, amphibole, biotite/phlogopite, quartz, and Fe-Ti oxides are also commonly reported. Scapolite and sapphirine are also reported from some localities (Meyer and Brookins, 1976;
w
a\
P
TABLE 1 Proterozoic lower crustal xenolith localities Locality
Host type
Xenolith type
Age hasalt
Reference xenolith
North America Navajo Volcanic Field, Colorado Plateau Camp Creek, Arizona Chino Valley, Arizona
K
L L
MGG, FG, Ec MGG, FG, Ec MGG, Am, Ec
25 Ma 25 Ma 25 Ma
1.2-1.9 Ga
MG, Int G. MG, An, Pg, FG
<3 Ma < 1 Ma
1.1-1.4 Ga 1.5Ga
Geronimo Volcanic Field, New Mexico Kilbourne Hole, New Mexico
AB
Engle Basin, New Mexico La Olivina. Mexico
AB
AB
MG MG, PG
<1 Ma
1.1 Ga-30 Ma
La Joya Honda, El Toro, La Joyuela, Santo Domingo, Mexico
AB
MG, PG
<1 Ma
1.1 Ga-30 Ma
Riley County, Kansas
K
MGG, Ec
AB
MG, An, PG, MMG
AB
<3 Ma
<230 Ma
Ehrenberg and Griffin, 1979 Esperanca et al., 1988 Arculus and Smith, 1979 Schulze and Helmstadt, 1979 Arculus et al., 1989 Kempton et al., 1990 Padovani and Carter, 1977a, b Davis and Grew, 1977 Padovani and Hart, 1981 James et al., 1980 Reid et al., 1989 Warren et al., 1979 Nimz et al., 1986 Ruiz et al., 1988a, b Roberts and Ruiz., 1989 Ruiz et al., 1988a, b Roberts and Ruiz, 1989 Selverstone and Stem, 1983
Europe Scotland
350Ma
360Ma
Van Breemen and Hawkesworth, 1980 Halliday et al., 1984, 1985 Hunter et al., 1984 Upton et al., 1983
.4
P13'
Africa
Hoggar, Algeria Lashaine, Tanzania
'4B '4B
MG, An, PG MGG, Ec, An
Lesotho, South Africa
K
MGG, MG, FG, E c
Leyleroup et al., 1982 Jones et al., 1983 Cohen et al., 1984 Davis, 1977; Rogers, 1977 Griffin et al., 1979 Harte et al., 1981 Rogers and Hawkesworth, 1982 Van Calstern et al., 1986
s
M
S'
2D 3
8
6.
Australia McBride Province, north Queensland,
0 2.
2
AB
MGG, MG, PG, FG
El Alamein, South Australia
K
MGG, Ec
Calcutteroo, South Australia
K
MGG, FG, Ec
< 3 Ma
300 Ma
Kay and Kay, 1981,1983 Rudnick and Taylor, 1987 Rudnick and Williams, 1987 Stolz and Davies, 1989 Stolz, 1989; Rudnick, 1990 Edwards et al., 1979 Arculus et al., 1989 McCulloch et al., 1982 Arculus et al., 1989
Abbreviations. Host types: AB = alkali basalt, K = kimberlite, L = latite; xenolith types: An = anorthosite, Ec = eclogite, FG = felsic granulite, MG = mafic granulite MGG = mafic garnet granulite, PG = paragneiss.
Gr: n
i3 2 Q
Q
43
2
s
$ O9
5z
366
J. Ruiz
Okrusch et al., 1979; Kay and Kay, 1983; Rudnick and Taylor, 1987; Loock, 1988; Stolz and Davies, 1989). Accessory phases are apatite, rutile, zircon, sulfides, and sphene. The mafic xenoliths have been alternatively described as cumulates, restitites or mantle-derived magma. In many cases the geochemical characteristics, such as LREE depletions or positive E u anomalies can be used to constrain the origin of the mafic orthogneisses, but it is sometimes difficult to distinguish between cumulates and restites. The intermediate and felsic granulites have a more variable mineralogy. Quartz, alkali feldspar and plagioclase are generally the major phases present. Garnet, pyroxenes, sillimanite, kyanite, Fe-Ti oxides, apatite, zircon, monazite, sulfides, scapolite, corundum, spinel and sapphirine are minor and accessory minerals. There are hydrous minerals in lower crustal xenoliths that indicate the presence of a fluid phase at lower crustal conditions. These minerals include biotite and amphibole and scapolite. Amphibole is the most common phase (e.g. northern Queensland, Australia, Stolz, 1987; Camp Creek, Arizona, Esperanca et al., 1988; Ichinomegata, Japan, Kuno, 1976, Aoki, 1971). Interestingly, none of the amphiboles or biotites found in lower crustal xenoliths seems to be F-rich (Okrusch et al., 1979; Ruiz et al., 1982; Wass and Hollis, 1983; Stolz, 1989). In exposed granulite-facies terranes, i.e. the Adirondacks, New York, the hydroxyl in the hydrous minerals is almost completely replaced by fluorine (Petersen et al., 1982). This substitution increases the stability field of amphiboles and micas into the granulite-facies pressure and temperature conditions. Although some of the amphiboles found in crustal xenoliths may be of primary origin (Wass and Hollis, 1983), most of the amphiboles seems to be a product of retrograde reactions (Okrusch et al., 1979; Stosch et al., 1984). Scapolite has been in xenoliths from central Mexico (Ruiz et al., 1982) but it also seems to have been formed by retrograde reactions. Since plagioclase and calcium carbonate can react to produce scapolite at about 1 atm (Newton and Goldsmith, 1975), scapolite may form in xenoliths that are emplaced on limestone. Peraluminous xenoliths contain garnet f pyroxene, quartz, sillimanite or kyanite, and feldspar (Padovani and Carter, 1977b; Roberts and Ruiz, 1989; Reid et al., 1989). Their mineralogy may represent restite after melt removal (e.g. Vielzeuf and Holloway, 1988; Clemens, 1990), or unmelted aluminous sediments metamorphosed at the base of the crust. The presence of the metasedimentary xenoliths attests to tectonic processes capable of emplacing sediments that where once on the surface of the Earth in the lower crust.
EFFECTS OF TRANSPORT
There is textural evidence that some lower crustal xenoliths interact with the host basalt during entrainment and ascent. Xenoliths may also interact with rocks on the surface of the Earth. Specially reactive rocks like limestone and dolomite (Ruiz et al., 1982). The most important processes capable of altering the original
Xenoliths in Proterozoic crust: evidencefor reworking of the lower crust
367
mineralogy and geochemistry of the xenoliths are heating of the xenolith by the host, infiltration of host magma, hydrothermal alteration by fluids derived from the host magma or from meteoric water, decompression-induced melting, mineral breakdown, and weathering at the Earth’s surface (Padovani and Carter, 1977b; Rogers, 1977; Griffin et al., 1979; Padovani et al., 1982; Rogers and Hawkesworth, 1982; Jones et al., 1983; Garvey and Robinson, 1984; Emery et al., 1985; Rudnick and Bylor, 1987; Arculus et al., 1989). The degree of post-entrainment alteration depends on the host. Kimberlites are fluid-rich and generally cause significant alteration in their entrapped xenoliths (e.g. Rogers, 1977; Griffin et al., 1979). Hydrothermal alteration is commonly seen in thin section in xenoliths from kimberlites (Griffin et al., 1979; Padovani et al., 1982). Elements such as Sr and Ba are greatly enriched in kimberlite-hosted xenoliths. Because these elements are known to be very mobile in most fluids, the high Sr and Ba contents are generally attributed to the precipitation of Sr-rich barite in cracks and in between mineral grains of the xenolith (Rogers, 1977; Griffin et al., 1979; Emery et al., 1985; Rudnick and Thylor, 1987; Arculus et al., 1989). Variations in K20, Rb and R E E have also been attributed to host-xenolith interactions (Rogers and Hawkesworth, 1982). Xenoliths carried by alkali basalts generally show less evidence of post- entrainment chemical modification even though volatiles must also have been present in these eruptions. Host penetration into the xenolith is the most common phenomenon observed in alkalic basalt hosted xenoliths.
GEOCHEMISTRY OF XENOLITHS
The average chemical composition of xenolith suites in Proterozoic crustal blocks is mafic (Rudnick and Presper, 1990). However, S i 0 2 contents range from less than 40 wt.% to almost 70 wt.%. The more siliceous rocks are generally aluminous paragneisses. Magnesium numbers range from almost 80 to 15 (Fig. 2). The mafic xenoliths cover the basalt spectrum from hypersthene to quartz normative basalts. Transition metals, such as Ni and Cr contents range from >lo00 ppm and >2000 ppm, respectively, for rocks with high Mg numbers to <10 pprn for rocks with low Mg numbers. The mafic xenoliths may represent mafic melts or cumulates or restites following crustal anatexis. It is dificult to chemically distinguish restites from cumulates and both types have been documented in xenolith suites (Rogers and Hawkesworth, 1982; Rudnick and Taylor, 1987; Roberts and Ruiz, 1989). Xenolith suites show variable depletions of large ion lithophile elements (LIL). Significant depletions may be caused by extraction of melts (e.g. Vielzeuf et al., 1990) or fluids during granulite-facies metamorphism (e.g. Roberts and Ruiz, 1989). The most commonly analyzed LIL elements are Ba, Rb, Sr and K. However available analyses of Cs show greater depletions than other LIL in xenoliths (e.g. Rudnick and Presper, 1990). Barium, for example, ranges from ca. 5 ppm to
J. Ruiz
368
30
40
50
60
70
80
Si02, Wt % Fig. 2. SiOz vs. Mg# for granulites. Triangles indicate xenoliths in kimberlites, lamprophyres, minettes and other magmatic rocks not associated with large scale magmatism. Open circles indicate xenoliths in alkalic basalts erupted in extensional regimes, where large scale magmatism has occurred. Mg# = 55 denotes the value above which few kimberlite-hosted xenoliths exist.
>lo00 ppm in mafic granulites. The highest Ba and other LIL concentrations are found in kimberlite-hosted xenoliths. This fact may indicate that Ba is introduced into xenoliths by fluids from the kimberlites (Griffin et al., 1979). However, an alternative possibility is that the lower crust sampled by kimberlites is different than the lower crust sampled by alkali basalts (Ruiz et al., in review). Rb contents in mafic xenoliths range from < 1 ppm to >300 ppm. Mafic xenoliths generally have the lowest Rb concentrations and paragneisses the highest values. The Rb variations in the xenoliths can be explained by loss of Rb during granulite-facies metamorphism or partial melting since Rb will removed by melts or fluids. High K/Rb ratios (up to 2500) (Leyleroup et al., 1977; Dostal et al., 1980; Stosch et al., 1986; Rudnick et al., 1986; Rudnick and Bylor, 1987; Roberts and Ruiz, 1989) (Fig. 3 ) are thought to result from the depletion of Rb relative to K (Sighinolfi, 1969; Heier and Thorensen, 1971; Rudnick et al., 1985). Plots of Rb/Sr vs. Sr indicate that the Sr content of xenoliths is independent of Rb/Sr ratio (Fig. 4). Generally, Sr content of upper and lower crustal rocks is not very different; however, Rb content can be greatly depleted in the xenoliths (e.g. Roberts and Ruiz, 1989). This phenomenon is shown by the correlation between Rb/Sr ratios and Rb content of xenoliths. As with Ba, there are systematic differences between Rb-Sr systematics of xenoliths of xenoliths emplaced by alkalic basalts in rifts and kimberlites and other rocks not associated with rifts (Figs. 4 and 5). The xenoliths in rifts have lower Rb/Sr and Rb contents than those from lamprophyres, kimberlites and minettes. These differences have been explained as metasomatism of the xenoliths by fluids in the kimberlitic melts. However, these differences are also seen in the Sm-Nd systematics (Figs. 6, 7). Figure 6 shows
Xenoliths in Proterozoic crust: evidence for reworking of the lower crust
.-
.*
'. . ,.-4.5 . .-. * 1000 I
,:r
-
I.
KlRb = 1000 500
i
' .
-
369
c n t
100
r
10
0.1"
'
I
'
.
"
'
I
.
j 0
0.00
0
500
1000
1500
Sr, PPm Fig. 4. Rb/Sr vs. Sr for xenoliths. Closed circles are samples from rifts and open squares are samples from lamprophyres and similar rocks. Data from work listed in Table 1.
that xenoliths have large variations in Sm/Nd ratios. However, xenoliths in rifts are the only ones with Mg# greater than 55 and Sm/Nd ratios less than 0.3. Generally, xenoliths in rifts also have the greatest Nd concentrations (Fig. 7). These differences have been interpreted by Ruiz et al. (1991) as original features of the xenoliths and not products of metasomatism. Consequently, these may be differences between the lower crust sampled in rifts and magmatically quiescent areas. Thorium and U are important elements because the Th/U ratio may be used to determine the mode of LIL depletion of the lower crust. Most igneous rocks have Th/U ratios between 3.5 and 4 (Rogers and Adams, 1978) (Fig. 8). Sedimentary rocks have similar Th/U ratios, between 2 and 4 for island arc greywackes (Nance and Taylor, 1977) to up to 10 for some shales (McLennan and Taylor, 1980). A compilation of U and Th (Rudnick et al., 1985; Rudnick and Presper, 1990) shows that Th/U ratios in some xenoliths are generally low, between 1 and 5.
J. Ruiz
370 0.15
(188,2.54)3'
(167,0.30)" 0
0
0.10
.
0
L
L c
z
0
0
0
on
0.05
o
0.00 0
20
40
60
Rb, PPm Fig. 5. Rb/Sr vs. R b €or xenoliths. Closed circles are samples from rifts. Open squares are samples from Iamprophyres and similar rocks. Data from references listed in Table 1.
0.400
0
60
70
80
Fig. 6. Sm/Nd vs. Mg# for xenoliths. Closed circles are samples from rifts. Open squares are samples from lamprophyres and related rocks. Mostly samples from rifts have Mg# > 55, and mostly samples found in larnprophyres and similar rocks formed by small magmatic events have Sm/Nd ratios greater than 0.3. Data from references listed in Table 1.
Although the Th/U ratios of many xenoliths are similar to fresh igneous rocks, their elemental abundance is usually quite low (U <0.09; Th <0.4 ppm). The low U and Th concentrations could reflect either primary features of the rocks or the effects of later depletion. U is much more soluble than T h in both water and carbonate-rich fluids (Rose et al., 1979). Therefore fluids should preferentially remove U and increase the Th/U ratio of rocks in the lower crust (Bylor and McLennan, 1985). This apparently has not occurred in all xenolith suites (Fig. 8). Removal of a melt phase would lower the concentration of U and T h without affecting their ratio because the distribution coefficients of both elements between
Xenoliths in Proterozoic crust: evidence for reworking of the lower crust
10
20
30
50
40
60
70
371
R0
Mg # Fig. 7. Nd vs. Mg# for xenoliths. Closed circles are samples from rifts. Open squares are samples from lampropyres and similar rocks. Samples from rifts generally have higher Nd contents. Data from references listed in B b l e 1.
.1
1
10
100
U (ppm) Fig. 8. U vs. Th for xenoliths. Ratios between 2 and 4 are common for unaltered sedimentary rocks (Nance and Taylor, 1977). Higher Th/U ratios would indicate fractionation of Th relative to U. Data from references in Table 1.
mineral and melt are similar (Henderson, 1982 and references therein). Rb and K would also be partitioned into the melt, decreasing their abundance in the xenoliths. Furthermore, if feldspar remained as a residual phase, Rb would be preferentially partitioned into the melt over K (Philpotts and Schnetzler, 1970), increasing the WRb ratio of the material remaining in the lower crust. It therefore appears that many orthogneiss xenoliths have had a melt extracted. Removal of a melt phase, either after formation of cumulates or following anatexis could explain the U/Th and K/Rb ratios of many granulites.
GEOCHRONOLOGYAND ISOTOPIC DATA
Perhaps some of the most interesting data from lower crustal xenoliths erupted through Proterozoic crust are those that constrain their age. The age of xenoliths
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provides information on how they relate to the rocks of the upper crust and to regional tectonics. However, dating xenoliths is one of the most difficult aspects of xenolith studies. A variety of isotopic methods have been employed to date xenoliths, including whole-rock and mineral Rb-Sr, Sm-Nd, and Pb-Pb isochrons, Nd model ages, K-Ar mineral ages and U-Pb methods on zircons. All these methods yield results that are often ambiguous and that attest to the complicated evolution of the lower crust. The most straightforward results have been obtained by ion microprobe U-Pb analysis on zircons (e.g. Rudnick and Williams, 1987). Zircons are particularly useful because they are resilient minerals even during granulite-facies metamorphism. Because of partial Pb loss of zircons, single zircon analysis generally gives discordant ages. If the lower crust has had a particularly complicated history that includes more than one granulite-facies metamorphic event, even high-resolution ion microprobes yield data that are difficult to interpret (xenoliths from L a Olivina, Mexico; Rudnick and Cameron, 1991). Because of the limited access to ion probes, the most commonly tried dating technique has been by Sm-Nd methods (e.g. Ruiz et al., 1988a; Wendlandt, 1991). However, model ages are have significant uncertainties (Amdt and Goldstein, 1987). At best they may give the crust formation age but at worst they may represent a mixed age (Rudnick, 1990). Intracrustal fractionation is particularly problematic since many mafic xenoliths may be cumulates and their Sm/Nd is not representative of mantle-crustal fractionation (Cameron et al., 1989). Model Nd ages from xenoliths from the Colorado Plateau yield remarkably consistent ages regardless of the Sm/Nd of the sample (Fig. 9). Furthermore, the ages are geologically reasonable. However, model Nd ages from xenoliths in rifts (Fig. 10)
E
0
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0
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AGE, Ga
Fig. 9. Model Nd ages for the Colorado Plateau. Data from Wendlandt (1991). Ages obtained for upper crustal rocks and xenoliths are similar suggesting little disturbance of the lower crust since Proterozoic times.
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MEXICAN LOWER CRUST
'.
orthoanriss
..metasedirnent
MEXICAN LOWER CRUST
I 1.5
1.0 AGE
L
I
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0
Go
Fig. 10. Model Nd ages for xenoliths from central Mexico. The Santo Doming0 and Ventura group of maars and the El Tor0 maar are in a rift. The age of the crust is thought to be Proterozoic. However, many mafic xenoliths yield younger model ages. Data from Ruiz et al. (1988a).
yield more complicated age patterns. A possible interpretation of these data is that there is younger material under rifts that is being sampled by the xenoliths. Nd model ages generally work best with felsic granulites since they cannot be juvenile mantle-derived material. Whole-rock Rb-Sr isochrons have been obtained from a variety of localities. However, linear arrays in isotopic systems can always be interpreted as mixing lines. Only one xenolith suite from Lesotho has yielded a poorly defined Sm-Nd linear array (Rogers and Hawkesworth, 1982). A 1.1 Ga Pb-Pb isochron was obtained for xenoliths from Camp Creek, Arizona by Esperanca et a]. (1988). However, there has been debate as to whether this line is a mixing line for young mantle-derived rocks (Johnson et al., 1990). Ages determined for mineral separates date the time of last equilibration. This information has been useful in determining the thermal history of the crust from which the xenoliths were obtained (Van Breemen and Hawkesworth, 1980; Padovani and Hart, 1981; Harte et al., 1981). Mineral ages are unlikely to yield information on the age of the protoliths because the xenoliths have been subjected to temperatures and pressures that would be above the closure temperature of the isotopic systems. Thus this technique is best used in conjunction with U-Pb or Sm-Nd whole-rock geochronology in order to constrain the geologic history of the lower crust in Proterozoic terranes.
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206Pb / 204Pb Fig. 11. s7Sr/86Srvs. 6Nd values for xenoliths. Note that paragneisses are generally the samples with highest s7Sr/86Sr and lowest ENd values. Many mafic orthogneisses have isotopic values close to bulk Earth. Data from references in B b l e 1.
The Sr and Nd isotopic compositions of lower crustal xenoliths in Proterozoic crustal blocks are shown in Fig. 11. The range of isotopic compositions is variable within each suite. Generally, the highest s7Sr/86Srand lowest 143Nd/144Ndvalues correspond to aluminous paragneisses. The xenoliths that plot close to bulk Earth values are generally mafic and intermediate orthogneisses. If the low Rb/Sr ratios were caused by LIL depletions, the event that caused the depletion had to be old in order to retard the growth of s7Sr/86Srratios so extremely. The high 143Nd/ 144Ndratios of the mafic xenoliths are due to the Sm/Nd ratio of the sample. In some cases it can be shown that samples that plot close to bulk Earth values are Proterozoic (e.g. Wendlandt, 1991). In other cases, however, the Nd isotopic values are indicative of young juvenile material added to the base of the crust (e.g. Rudnick and Cameron, 1991). Lead isotope data can also be used to crudely date the lower crust. Plots of 207Pb/204Pb vs. zo6Pb/204Pbclearly show that many mafic xenoliths in Proterozoic crustal blocks have isotopic compositions that are more similar to those of presentday mantle, as represented by MORB (Fig. 12). The only mafic xenoliths that yield Pb isotopic compositions unradiogenic enough to be Proterozoic are those that are found in volcanic fields where there has been extensive young magmatism (regions of extension). Xenolith suites from rifts, where there has been long-lived
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‘NO
Fig. 12. Lead isotopic data for xenoliths worldwide. 4.55 Ga. Geochron for reference. Samples from Highwood Mountains, Montana,paragnekses from Kilbourne Hole and sample from the Lesotho Kimberlite have less radiogenic lead composition than samples from extensional regimes where magmatism is of a greater scale. Data from Highwood Mountains and Mexico (Ruiz et al., 1991). The rest of the data are from Rudnick and Goldstein (1989).
magmatism, have isotopic compositions indicating a recent derivation from the mantle. These data suggest that xenoliths from rifts may not be representative of the lower crust beneath most Proterozoic crustal blocks (Ruiz et al., in review). The lead isotopic data clearly indicate that there is significant interaction between the lower crust and the mantle in rifts and that the lower crust in these tectonic environments must be intruded by significant amounts of mantle-derived mafic magmas of the age of rifting (Amdt and Goldstein, 1989; Rudnick and Goldstein, 1989; Kempton et al., 1990; Downes et al., 1991). COMPARISON OF GRANULITE XENOLITHS AND EXPOSED GRANULITE-FACIES PROTEROZOIC CRUST
There are compositional differences between granulite-facies xenoliths and exposed granulite terranes. Because some of the differences can be caused by metamorphism, only exposed terranes that have been metamorphosed to granulitefacies metamorphic conditions should be directly compared with granulite-facies xenoliths. On average, crustal xenoliths are more mafic (Fig. 2) and contain lower abundances of incompatible elements than exposed granulite-facies terranes (Rudnick and Presper, 1989; Roberts and Ruiz, 1989). The granulite-facies terranes are dominated by supracrustal sequences that have been variably depleted in selected elements during metamorphism. The zonation that is observed in the major element chemistry also extends to trace elements (e.g. Roberts and Ruiz, 1989). Granulite facies-xenoliths are generally more depleted in LIL than exposed granulites.
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The chemical differences between xenoliths and terranes suggest that granulitefacies terranes and xenoliths represent different parts of the crust (Roberts and Ruiz, 1989). This interpretation is supported by differences in the equilibration pressures between terranes and xenoliths (Bohlen and Metzger, 1989). In general, equilibration pressures are higher for xenoliths ( 2 8 kb) than terranes (ca. 8 kb). A chemically stratified crust composed of mafic material at the deepest levels underlying diverse lithologies, more typical of granulite-facies terranes, at higher levels is also supported by refraction profiles (Jackson, 1989). Velocities greater than 7 km/s, which indicate denser, mafic material, only typically occur at depths of 30 km or more.
EVOLUTION O F THE LOWER CRUST I N PROTEROZOIC CRUSTAL BLOCKS
Proterozoic crustal blocks have younger lower crust in areas where there has been extensive volcanism. This observation is evident from xenoliths worldwide that have been brought to the surface in regions of protracted magmatism. The xenolith suites studied thus far are biased toward the study of the lower crust in rifts. This is mainly the result of finding the best preserved and most abundant xenolith suites in volcanic rocks in rifts. Fewer studies have been done on xenoliths found in smaller volume magmatic fields such as lamprophyres and minettes. Interestingly, the main xenolith type found in lamprophyres and minettes is also mafic. However, comparison of the chemical composition of the xenoliths from minettes-lamprophyres-kimberlites and alkalic rocks in rifts shows differences in Mg# with the mafic xenoliths from rifts generally having Mg# less than 55 and xenoliths from lamprophyres and minettes having Mg# up to 79 (Fig. 2; Ruiz, 1991). In addition to the young mafic magmas, the lower crust in Proterozoic crust contains granulites that in many cases can be interpreted to be former sedimentary rocks that were once at the Earth’s surface (Padovani and Carter, 1977; Roberts and Ruiz, 1989). The tectonic regimes that are capable of underthrusting sediments to the base of the crust probably occur at the edge of the continents during continental collisions or subduction. Thus it seems that the chemical and lithological characteristics of the lower crust in Proterozoic terranes is controlled by younger tectonic regimes that affect the whole crust. Thus, the lower crust will become more mafic with time in areas of abundant magmatism. During these events, the lower crust may become an open boundary with the upper mantle. This has also been proposed by Arndt and Goldstein (1989) based on density considerations. Bi-directional transport may also explain some isotopic anomalies of the mantle (Zartman and Haines, 1988; Arndt and Goldstein, 1989). The interaction between lower crust and mantle has also been documented in the Ivrea Zone (Voshage et al., 1990), where the lower crust has been exposed by tectonic processes.
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ACKNOWLEDGEMENTS
Pace element and Pb isotopic work on xenoliths has been supported by NSF grant EAR 9011234. The W.M. Keck Foundation funded the ICP-MS, used for the collection of the data. I thank two anonymous reviewers and K. Condie for their detailed comments on an earlier draft of this manuscript. Their reviews considerably improved this chapter. This chapter was prepared while I was on sabbatical leave at Berkeley, where I had many stimulating discussions with students, specially J. Lassiter, and E. Wendlandt on topics of the lower crust. Excellent reviews of the geochemistry of the lower crust by Rudnick and Presper (1990) and Rudnick (1992) considerably simplified the work for this review.
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Ruiz, J., Essene, E.J. and Ortega, G.E, 1982. Pyroxenites from La Joya Honda Maar, San Luis Potosi, central Mexico. EOS, Trans. Am. Geophys. Union, 6 3 463. Ruiz, J., Wendlandt, E., Lassiter, J. and Rattray, G., 1991. Differences between the lower crust in cratons and rifts. EOS, Trans. Am. Geophys. Union. Schulze, D.J. and Helmstaedt, H., 1979. Garnet pyroxenite and eclogite xenoliths from the Sullivan Buttes latite, Chino Valley, Arizona. In: ER. Boyd and H.O.A. Meyer (Editors), The Mantle Sample: Inclusions, Kimberlites and Other Volcanics. Am. Geophys. Union, Washington, D.C., pp. 318-329. Selverstone, J. and Stern, C.R., 1983. Petrochemistry and recrystallization history of granulite xenoliths from the Pali-Aike volcanic field, Chile. Am. Mineral., 68: 1102-1111. Sighinolfi, G.P., 1969. K-Rb ratio in high grade metamorphism: a confirmation of the hypothesis of a continual crustal evolution. Contrib. Mineral. Petrol., 21: 346-356. Stolz, A.J., 1987. Fluid activity in the lower crust and upper mantle: mineralogical evidence bearing on the origin of amphibole and scapolite in ultramafic and mafic granulite xenoliths. Mineral. Mag., 51: 719-732. Stolz, A.J. and Davies, G.R., 1989. Metasomatized lower crustal and upper mantle xenoliths from North Queensland: chemical and isotopic evidence bearing on the composition and source of the fluid phase. Geochim. Cosmochim. Acta, 5 3 649-660. Stosch, H.G., Lugmair, G.W. and Seck, H.A., 1984. Evolution of the lower continental crust: granulite facies xenoliths from the Eifel, West Germany. Nature, 311: 368-370. Stosch, H.G., Lugmair, G.W. and Seck, H.A., 1986. Geochemistry of granulite facies lower crustal xenoliths: implications for the geological history of the lower continental crust beneath the Eifel, West Germany. In: B. Dawson, D.A. Carswell, H. Hall and K.H. Wedepohl (Editors), The Nature of the Lower Continental Crust. Geol. SOC.London, Spec. Publ., 2 4 331-350. Bylor, S.R. and McLennan, S.M., 1985. The Continental Crust: Its Composition and Evolution. Blackwell Scientific Publishers, Oxford, 312 pp. Toft, P.B., Hills, D.V. and Haggerty, S.E., 1989. Crustal evolution and the granulite to eclogite transition in xenoliths from kimberlites in the west African craton. Tectonophysics, 161: 318-333. Upton, B.G.J., Aspen, P. and Chapman, N.A., 1983. The upper mantle and deep crust beneath the British Isles: evidence from inclusion suites in volcanic rocks. J. Geol. SOC.,140: 105-122. Van Breemen, 0. and Hawkesworth, C.J., 1980. Sm-Nd isotopic study of garnets and their metamorphic host rocks. Bans. R. SOC.Edinburgh, 71: 97-102. Van Calsteren, P.W.C., Harris, N.B.W., Hawkesworth, C.J., Menzies, M.A. and Rogers, N.W., 1986. Xenoliths from southern Africa: a perspective on the lower crust. In: B. Dawson, D.A. Carswell, J. Hall and K.H. Wedepohl (Editors), The Nature of the Lower Continental Crust. Geol. SOC.London, Spec. Publ., 2 5 351-362. Van Calsteren, P.W.C., Kempton, P.D. and Hawkesworth, C.J., 1988. Depletion of U in the lower crust: evidence from granulite xenoliths from southern Africa. Chem. Geol., 70: 74 (abstract). Vielzeuf, D. and Holloway, J.R., 1988. Experimental determination of the fluid-absent melting relations in the pelitic system. Consequences for crustal differentiation. Contrib. Mineral. Petrol., 98: 257-276. Vielzeuf, D., Clemens, J.D., Pin, C . and Moinet, E., 1990. Granites, granulites and crustal differentiation. In: D. Vielzeuf and Ph. Vidal (Editors), Granulites and Crustal Evolution. NATO, AS1 Ser. C, 311: 59-85. Voshage, H., Hoffman, A.W, Mazzucchelli, M., Rivalenti, G., Sinigoi, S., Raczek, I. and Demarchi, G., 1990. Isotopic evidence from the Ivrea Zone for a hybrid lower crust formed by magmatic underplating. Nature, 347: 731-736. Warren, R.G., Kudo, A.M. and Keil, K., 1979. Geochemistry of lithic and single crystal inclusions in basalts and a characterization of the upper mantle-lower crust in the Engle Basin, Rio Grande Rift, New Mexico. In: R.E. Riecker (Editor), Rio Grande Rift: Tectonics and Magmatism. Am. Geophys.
382
J. Ruiz
Union, Washington, D.C., pp. 393-415. Wass, S.Y. and Hollis, J.D., 1983. Crustal growth in south eastern Australia - evidence from lower crustal eclogitic and granulitic xenoliths. J. Metamorph. Geol., 1: 25-45. Wendlandt, E., 1991. Nd Isotopic Study of Lower Crustal Xenoliths from the Colorado Plateau. Ph.D. Dissertation, University of California at Berkeley, 350 pp. (unpublished). Windrim, D.P. and McCulloch, M.T., 1986. Nd and Sr isotopic systematics of central Australian granulites: chronology of crustal development and constraints on the evolution of the lower continental crust. Contrib. Mineral. Petrol., 94: 289-303. Zartman, R.E. and Haines, S.M., 1988. The plumbotectonics model for Pb isotopic systematics among major terrestrial reservoirs - a case for bi-directional transport. Geochim. Cosmochim. Acta, 5 2 1327-1339.
383
Chapter I0
PROTEROZOIC IRON-FORMATIONS * C. KLEIN AND N.J. BEUKES
INTRODUCTION
This chapter presents an overview of major Proterozoic iron-formations ranging in age from 2.5 Ga to about 0.8 Ga (the latter being the approximate age for iron-formation in the Rapitan Group, Yukon and Northwest Territories, Canada). The term iron-formation as used here is similar to that of James (1954) with some modifications as suggested by %endall (1983a). As such, iron-formation is defined as: a chemical sediment, typical& thin-bedded or laminated, whose principal chemical characteristic is an anomalous& high content of iron, common& but not necessarily containing layers of chert. Different bulk chemistries, reflected in major mineralogical variations, are expressed in the nomenclature of four different types: (i) oxide iron-formation; (ii) carbonate iron-formation; (iii) sulfide iron-formation and (iv) silicate iron-formation. In James’s original definition of 1954, a quantitative lower limit of iron content (215 wt.% of iron) was incorporated. This arbitrary lower limit is commonly too restrictive in the evaluation of rock types that reflect a range in iron content (from ferruginous to iron-rich) all of which show the pertinent characteristics of iron-formation, and which are part of iron-rich sedimentary sequences. Although many Proterozoic iron-formations have undergone various degrees of metamorphism, most of the interpretations in this overview are based on data derived from iron-formations that have undergone only low-grade metamorphic (lower greenschist) facies conditions. These relatively unmetamorphosed iron-formations consist mainly of chert, magnetite, various carbonates (siderite, members of the dolomite-ankerite series, and calcite), hematite, and silicates such as greenalite, stilpnomelane, minnesotaite and riebeckite. The mineralogy of iron-formations as a function of diagenesis and metamorphism is reviewed briefly in a subsequent section.
*Parts of this chapter are reprinted with permission from sections 4.2 (C. Klein and N.J. Beukes) and 4.3 (N.J.Beukes and C. Klein) in The Proterozoic Biosphere: A Multidisciplinary Study (J.W. Schopf and C. Klein, Editors), Cambridge University Press, 1992.
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384
DISTRIBUTION OF IRON-FORMATIONS THROUGHOUT THE PRECAMBRIAN
The assessment of the abundance of iron-formations with time in the Precambrian (Fig. 1) is based on compilations by James (1983) and Walker et al. (1983). The age assignments for specific banded iron-formations (BIFs) in these tabulations appear not to have changed significantly since 1983. For example, Walker et al. (1983) report an age for the Labrador Trough sequence as 1.87 Ga (from Fryer, 1972); Hoffman (1988) gives a date for the Wishart-Sokoman-Menihek sequence of the Labrador Trough (the New Quebec Orogen) as 1.88 Ga; and ChevC and Machado (1988) report a date of 1.88 Ga (f2 Ma) for the Sokoman Formation. In short, in an overview of iron-formation abundances throughout Precambrian time, the dates of many BIFs are now probably better known than their absolute abundances. Interpretations of relative abundances of iron-formations versus age have been made by Gole and Klein (1981), James and Trendall (1982), James (1983), and Walker et al. (1983), with results that differ mainly in the estimated abundance of Archean iron-formations. There is general agreement that the most extensive iron-formations appear to be those of the Hamersley Range of Western Australia ( ~ 2 . 5Ga) and the Transvaal Supergroup of South Africa, with an age of 2.43 f0.3 Ga (Trendall et al., 1990). There is additional agreement on a sharp decline in iron-formation abundance in the Precambrian record at or before 1.8 Ga, except for several that show approximate ages between 0.8 and 0.6 Ga (such as ironformations in the Rapitan Group, N.W.T., Canada; in the Urucum-Mutun region, Brazil-Bolivia; and in the Damara Supergroup of Namibia). There is, however, less agreement on the evaluation of the relative abundances of the generally
Hamersley Transvaal Supergroup,S Africa Canadian greenstone belts; Yilgarn Block. W.A.
Lake Superior Region, USA
Zirnbabwe;S.Africa; Ukraine, U.S.S.R; Venezuela; Western Australia
4.5
4.0
Rapitan Group, Canada ; Urucum Region. Brazil ;
3.0
2.5
2.0
1.0
0
Ga Fig. 1. Schematic diagram showing the relative abundance of Precambrian banded iron-formations versus time, with several of the major iron-formations or major iron-formation regions identified. Estimated abundance values are relative to the Hamersley Group banded iron-formation volume taken as a maximum (from Klein and Beukes, 1992). Data for this plot were obtained from James (1983) and Walker et al. (1983).
Proterozoic iron-formations
385
smaller, discontinuous, and commonly tectonically deformed iron-formations in the Archean. Gole and Klein (1981) conclude that the size and extent of Archean iron-formations have commonly been underestimated. The schematic relative abundance curve shown in Fig. 1 is adapted from that given by Gole and Klein (1981) with some minor modifications and with names added for several wellstudied and major iron-formations. This curve is in essence a “best estimate” of volume of iron-formation (not iron ore) for all iron-formations tabulated by James and Pendall (1982), James (1983), and Walker et al. (1983), relative to a maximum represented by the total iron-formation content of the Hamersley Range of Western Australia.
METAMORPHISM OF IRON-FORMATION
Several well-known Proterozoic iron-formations have undergone only lowgrade (greenschist) metamorphism whereas others show a considerable range in metamorphic grade. Very-low-grade metamorphic mineral assemblages make up most of the iron-formation lithologies of the Hamersley Basin (Trendall and Blockley, 1970; Klein and Gole, 1981). Almost all of the extensive ironformations in the llansvaal Supergroup of South Africa (see Figs. 11 and 12) reflect only very low temperature conditions of metamorphism, as concluded from maximum estimated temperatures of 110°C to 170”C, at pressures not more than 2 kbar (Miyano and Beukes, 1984; Klein and Beukes, 1989). Well-preserved low-grade metamorphic regions occur in the Sokoman Iron Formation of the west-central part of the Labrador llough (Dimroth, 1968; Klein, 1974; Klein and Fink, 1976; Lesher, 1978), and parts of the Negaunee Iron Formation in northern Michigan also contain assemblages that reflect only late diagenetic to low-grade metamorphic conditions; the lowest metamorphic grade rocks in northern Michigan have been assigned to the chlorite grade by James (1955) and classified as part of “zone l” by Haase (1979 and 1982). Location maps, generalized geologic maps and distributions of metamorphic zones for several major Proterozoic iron-formations that show a range of metamorphic grade are given in Fig. 2. The assemblages in late diagenetic and very-low-grade metamorphic ironformations consist of various combinations of the following minerals: chert (or quartz, as recrystallized chert), magnetite, hematite, siderite, calcite, members of the dolomite-ankerite series, greenalite, stilpnomelane, minnesotaite, riebeckite (and the fibrous variety crocidolite), ferribiotite, and iron sulfides, mainly pyrite and pyrrhotite. Minor amounts of chamosite, ripidolite (an Fe-rich chlorite) and talc are also found. Most of these minerals occur in mesobands (with an average thickness of about an inch, or about 2.5 cm, made of alternating chert, or quartz and iron-rich minerals), in massive lutitic beds or mesobands, in fine laminations or in microbands (with a band thickness of about 0.3 to 1.7 mm), and in oolitic or granular textures. Examples of some of the very low-grade metamorphic
C.Klein and N.J. Beukes
386
...
,----\
\
1
/
/
\
\
'.
EXPLANATION
lsograds (after James, 1955) Metamorphic zones (after Hoase. 1979)
I
Babbitt
I
Biwabik Iron F
-
' 30 K m
1
WISCONSIN
Fig. 2. Locations, extent and various metamorphic zones in some Proterozoic iron-formations (from Klein, 1983). A. Generalized regional geology of the Marquette Trough which includes the Negaunee Iron Formation. Identified localities are mines or exploration sites (after Haase, 1982). B. Generalized geologic map of part of the Mesabi Range showing the distribution of the Biwabik Iron Formation and metamorphic zones (from French, 1968).
Proterozoic iron-formations 70°
690 I
I
387 66"
67O
68'
I
ARDUA L A K E A R E A
15'
"Schefferville
Edge o Labrador Trough' INDEX MAP SHOWING LOCATION OF CABRQOOR TROUGH
40 Miles
0
0
ARCHEAN BASEMENT
50 Km /
m
Sowbill Lake
Iran formation
.-
I
,
IC
; >by], -l,
-. ' I
/ -Fire
GRENVILLE PROV I NC E Lake Deposit
Fig. 2 (continued). C. Generalized geologicmap of the distribution of iron-formation in the central and southwestern parts of the Labrador Trough (after Klein, 1966 and 1978). Identified localities refer to mining towns and exploration sites.
assemblages are given in Fig. 3, and the compositional extents of several common low-grade metamorphic silicates are given in Fig. 4. Medium-grade metamorphic assemblages of iron-formation are characterized by the common development of amphiboles, mainly members of the cummingtonite-
C.Kkin and N.J. Beukes
388 A120,(20 mole
W)
A1,0,(20
mole
%I
h
t mag t q t z
,6
+mag t qtz (or ch)
y,\
M
B
/ MgO
/
minnesotaite 50
stilpnomelanef
FeO t MnO
CaO
minnesotaite.
LX
50
stilpnornelane
MINERAL ABBREVIATIONS onk - ankerite colc - calcite ch - chert ferrodol - ferrodolomite mag - magnetite minn - minnesotaite py - pyrite qtz - quortz
Fig. 3. Graphical representation of silicate-carbonate assemblages, and carbonate compositions in late diagenetic to very-low-grade metamorphic iron-formation assemblages. Tielines connect coexisting minerals. A and B are from the Sokoman Iron Formation in the Howells River area, west-central Labrador Trough (from Klein, 1983).
grunerite series. For example, the Negaunee Iron Formation in the Lake Superior region of the U.S.A. and the Sokoman Iron Formation, especially in the region of Labrador City, Canada show extensive development of coarse-grained cummingtonite- and grunerite-rich assemblages (see Fig. 5). Such medium-grade conditions are found in the Bansvaal Supergroup iron-formations only in the Penge area, in close proximity to the Bushveld Complex (Beukes and Dreyer, 1986). The iron-formations in the Krivoy Rog Basin of the Ukraine show extensive medium- to high-grade metamorphic assemblages (Belevtsev et al., 1983). Although members of the Fe-Mg- and Ca-amphibole series are common in
Proterozoic iron-formations
389
TALC -MINNESOTAITE SERIES 0
0 0 X
Arduo Lake area (Lesher, 1978) Howells River area (Klein,19741
GREENALITE
Gunflint Iron Formation (Floran 8 Papike.1975) Dales Gorge Member, Hamersley Range (Miyano.1978)
compositional field (Klein. 1974 and Floron 8 Papike. 19751
STILPNOMELANE A A P
Ardua Lake area (Lesher. 1978) Howells River area (Klein,1974 and Klein 8 Fink.1976) Gunflint Iron Formation (Floron 8. Papike, 1975)
60
60
1..
stilpnomelane
/
I.
10 10
XX .
I
20
/’
A
\\
/ //
/’ //
A
A A A ( A (\--_------A \--_-------talc
Mg Mg
%’
/
X X
- rninnesotaite
0
XI.YXXXI
30
40
-.
series I
50 50
I .
60 60
I
70
80
rninnesotaite rninnesotaite’
90
+
F n Fee + M Mn
Fig. 4. Compositional ranges of the talc-minnesotaite series, (Fe,Mg)3 Si4010 (OH);?; greenalite, Fe6 (O,oH);?7.2-4H;?O in late diagenetic to very-low-grade metamorphic iron-formations (from Klein, 1983). Si4O10 (OH)8; and stilpnomelane, K0.6(Mg,Fe2+,Fe3+)6
medium-grade iron-formation assemblages, they are not the exclusive silicates because some minnesotaite and slilpnomelane may persist into biotite and garnet zones, respectively (Klein, 1978, 1983). Members of the pyroxene group may be sporadically present in medium-grade metamorphic iron-formation occurrences, but their abundance is generally very minor compared to that of amphiboles. As such, the most common minerals at medium metamorphic grade are: quartz, magnetite, hematite, members of the cummingtonite-grunerite series (and the fibrous variety amosite), Ca-rich amphiboles such as actinolite and hornblende; less common are hedenbergite, eulite, ferrosalite, and aegirine-augite; fayalite is rare; and garnet is found only in iron-formations with a considerable A203content. Most of the carbonate species, present at lower metamorphic grade, survive during medium-grade conditions, although the overall percentage of the carbonates is reduced on account of the formation of new silicates, as a result of the reaction of original carbonates with chert (or quartz). For a discussion of the temperature and pressure conditions of medium-grade metamorphism of iron-formations see Klein (1983). High-grade metamorphic iron-formations consist of essentially anhydrous assemblages in which variable amounts of ortho- and clinopyroxene predominate (see Fig. 6). Fayalite may be present, as well as carbonates and garnet, and lesser amounts of amphiboles; quartz, magnetite and/or hematite are still the major constituents of oxide-rich iron-formations. Assemblages that reflect the highest
C.k%in and N.J. Beukes
390 LABRADOR TROUGH x
e D
Butler (1969) Kranck (1961) Klein (1966)
Klein (1968) Mueller (1960)
v 0
Ca, Mg,SieO,,(OH)~
Mg7SieOz,(OH),
Ca,Fe5
I:
SieO,,(OH),
Mole % FeO
LAKE SUPERIOR
+ o
Ca, Mg, S isO,(OH),
I
Mg7Sie0,,(OH)z
Bonnichsen (1969) Floran (1975)
I
Q)
Morey, et 01. (1972) Simmons, etal. (1974)
+
I
Mole % FeO
Fe7Sie022(OH)z
NEGAUNEE IRON FORMATION
Fig. 5. Compilations of compositional ranges and major element fractionation data for amphiboles from several medium-grade metamorphic Proterozoic iron-formations (from Klein, 1983).
metamorphic temperatures are the result of contact metamorphism, as in the case of the Gunflint and Biwabik Iron Formations (in the Lake Superior region of the U.S.A.) near the contact with the Duluth Gabbro Complex. Somewhat lower temperatures, but higher pressures are characteristic of regionally metamorphosed iron-formations such as those in the southwestern part of the Labrador Trough (see Fig. 2C; Klein, 1978, 1983). It appears that calcite and members of the dolomite-ankerite series are abundant throughout the complete metamorphic range exhibited, but that siderite becomes less abundant at the highest grades. This indicates that the FeCO3 component is more involved in the production of
391
Proterozoic iron-formations
Ca Mg Si206
Ca Fe Si206
LABRADOR TROUGH x Butler (1969) e Kranck (1961) v Klein (1966) E Mueller (1960)
Mole % FeO
MgSiOJ
FeSiO,
A
o
n o
M~
10
20
30
40
50
60
70
80
90
clinopyroxene fayalite grunerite orthopyroxene pigeonite (inverted)
Fe
Fig. 6. A, B. Compilations of compositional ranges and major element fractionation data of orthopyroxene and clinopyroxene in metamorphozed iron-formations in the Labrador Trough, Canada and in the Lake Superior region of the U.S.A. C. Compositional ranges of orthopyroxene, clinopyroxene and fayalite (and lesser grunerite) in the highest metamorphic zone of the Biwabik Iron Formation (from Klein, 1983).
C. Kkin and N.J. Beukes
392
1
GRADE OF METAMORPHISM
1
LOW
81OT ITE
DIAGE N E T I C Early
yt:k;
GARNET
'ONE
ate
I chert "Fe,O,-
HIGH
MEDIUM
'ONE
liti
SILL I M A N IT E
K Y A N I T E ZONE
'ONE
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quartz
H20"+
magnetite
"Fe (OH)3"-
hematite
greenalite
I
I
stilpnomelbne
--
+
---
I
F i i b n atalc Fe t--chlorite i minnesotaite t e (ripidolitel
-
I I
I
I
I
I
--
dolomite - ankerite
I
I
calcite siderite
--
- magnesite
--- +---- -+--
riebeckite
cummingtonite - grunerite (anthophyl lite) Itremolite - ferroactinolile (hornblende) I
I I I
---
--* ------
;
t
5
------
almandine
I I
I
I I
I I
+---
I I
I
+---
I
I
I
I
or thopyroxene clinopyroxene
--
fayalite
Fig. 7. Relative stabilities of minerals in metamorphosed iron-formations as a function of metamorphic zone (from Klein, 1983).
metamorphic silicates than calcite or dolomite (or ankerite); for example, 8(Fe7Mg)C03 + 8SiO2 siderite
quartz
+ H20 + (Fe7Mg)-/Sig022(OH)2+ 7 C 0 2 grunerite
A schematic diagram of the relative mineral stabilities in iron-formations ranging from very low to the highest metamorphic grade is given in Fig. 7.
MAJOR ELEMENT CHEMISTRY OF SEVERAL MAJOR IRON-FORMATIONS
The averages for the major oxide components for a range of Precambrian ironformations are given in Fig. 8 on an H20- and C02-free basis. All original analyses were recalculated in this way in order to make the analyses of unmetamorphosed and metamorphosed iron-formations comparable, assuming that metamorphic reactions are essentially isochemical except for the loss of H2O and C02. The chemical similarities between, for example, the unmetamorphosed and highly metamorphosed Labrador Pough iron-formations (Klein and Beukes, 1992, table
393
Proterozoic iron-formations
.
=
lS"a
0
=
Yilgarn
X
=Montana
=
Marra Mamba
0
=
Dales Gorge
0
=
BIF
0= S
Bands
0+
Kururnan
OX.
L 3
I
Kuruman Sad.
B
I
BiWabik
GriqUatOwn pel.
MnO
A
~
*
r
Labrador
""met
=
Labrador met
=
Rapitan
AI,O,
K,O
Na,O 1
P,05
-
Fig. 8. Plot of the averages of major chemical components of iron-formations with the analytical results recalculated to 100% on an H20- and CO2-free basis. The original data for this figure are given in Klein and Beukes (1992, table 4.2.1).The shaded area enclosed by thin dashed lines brackets the overall range of all values except those of the Rapitan Group iron-formations which are denoted by solid triangles connected by a thick dashed line. Note that data for the various components are presented relative to three different (vertical) weight percentage scales (from Klein and Beukes, 1992).
4.2.1) suggest that this is a reasonable assumption for the major element data. The general similarity of all of the averages, except those for the Rapitan ironformation, is well shown in Fig. 8. In the selection of these analyses, every effort was made to exclude iron-formation samples that had undergone any type of secondary alteration, such as oxidation or leaching, thus excluding any materials that might be considered iron ore, or in the process of becoming ore. This is reflected in the values for total iron, ranging from a minimum of 23 wt.% (S bands of the Brockman Iron Formation sequence) to a maximum of 34 wt.% (BIF of the Brockman Iron Formation), as well as in the Fe3+/(Fe2+ + Fe3+)values. The latter range from 0.05 (in siderite BIF in the Kuruman Iron Formation) to 0.58 (in metamorphosed Archean iron-formation in Montana), except for one value of 0.97 (in Rapitan BIF). It is instructive to compare the 0.05 to 0.58 range with Fe3+) = 0.671 the values for two of the iron oxides, magnetite [Fe3'/(Fe2' and hematite [Fe3+/(Fe2' + Fe3') = 11. This illustrates that the iron in most iron-formations is in an average oxidation state between that of wiistite (FeO) and that of magnetite (Fe304), reflecting the very common association of magnetite with Fe2' - containing minerals such as carbonates (siderite and ankerite), silicates (such as greenalite, in essentially unmetamorphosed iron-formations; minnesotaite; members of the cummingtonite-grunerite series; and members of the orthopyroxene series in metamorphosed assemblages), and locally pyrite. The
+
C.Kkin and N.J;Beukes
394
a
1
0.5
q/
0.5 l F b
Seawater: Hydrothermal Fluid Mix = 1OOO:l
J \
I.
i
0
cn
.
. -
a z al
al
0.1
Q
s cn
5 cn
0.1
O.O!
K r u m a n mag-sid BIF(4) I I I 1 I 1 / Y b Lu . a c e Nd S m E u Tb I
1
I
Lace
Nd
SmEu Tb
I
Y b LU
C
0.5
:a:
z . g
0.1
E
0
0.05
0.01
Lace
Nd Sm EU Tb
Y b Lu
Fig. 9. R EE (rare earth element) patterns, normalized to the North American Shale Composite (NASC). The original data for this figure are given in Klein and Beukes (1992, table 4.2.2). a. Average patterns for five iron-formation types compared with that for a calculated mixture of 1000 parts of North Atlantic seawater from 100 m depth (Elderfield and Greaves, 1982) and one part of hydrothermal fluid from the East Pacific Rise (Michard et al., 1983). The seawater: hydrothermal fluid mix pattern is multiplied by lo6. Among the iron-formation patterns, there is a clear decrease in overall REE abundance, as well as a decrease in the positive Eu anomaly, with decreasing geologic age. b. Patterns for some deep marine hydrothermal deposits from the Galapagos mounds (Corliss et al., 1978) compared with the average pattern for siderite-rich iron-formation from the Kuruman sequence.
Proterozoic iron-formations
395
+
average Fe3+/(Fe2+ Fe3+) value for all averages shown in Fig. 8, but excluding the Rapitan value is 0.42, which recalculates to an average oxidation state for iron in iron-formations of Fe2.4+.This means that any models for iron-formation precipitation require considerably less oxygen input than would be needed if iron-formations are assumed to consist mainly of hematite, Fe203, as has been commonly done in the literature. The only iron-formation that is radically different from any of the others is the Rapitan iron-formation, which consists of chert and hematite, and is devoid of Fe-carbonates or silicates.
RARE EARTH AND TRACE ELEMENT CHEMISTRY OF SEVERAL MAJOR IRON-FORMATIONS
The average rare earth element (REE) patterns for several iron-formations, normalized to the North American Shale Composite (NASC; Gromet et al., 1984), are given in Fig. 9. Figure 9a shows REE patterns for the Archean Isua ironformations and the very early Proterozoic Brockman and Kuruman microbanded iron-formations. All patterns are essentially the same showing pronounced positive Eu anomalies, negative Ce anomalies, and a depletion in the light REE. Such patterns, also described by Fryer (1977, 1983) for other Archean and Proterozoic iron-formations, are very similar to those of some modern deep sea hydrothermal deposits (Fig. 9b) and the element distribution that is obtained when modern deep sea hydrothermal fluids are mixed with typical North Atlantic Seawater. The pattern of such a mixture, consisting of 1000 parts of North Atlantic seawater from 100 m depth (Elderfield and Greaves, 1982) and one part of hydrothermal fluid from the East Pacific Rise (Michard et al., 1953), is shown in the upper part of Fig. 9a. Such mixing calculations, originally made by Dymek and Klein (1988), produce REE patterns with a very strong similarity to those of the various iron-formations. Derry and Jacobsen (1990) point out that such calculations are in principle correct, but in practice neglect the effects of REE fractionation during the mixing of hydrothermal fluids and seawater, and are critically dependent on the assumption about the size of the Eu anomaly in Precambrian hydrothermal fluids. Nonetheless, the striking similarity between a calculated pattern and that shown by several BIFs leads to the conclusion that Precambrian iron-formations ranging in age from those at Isua (3.8 Ga) to those about 1.9 Ga-old, (e.g. the Sokoman Iron Formation) are the result of chemical precipitation from solutions that represent mixtures of seawater and hydrothermal input. Figure 9b shows the similarity of the iron-formation patterns to those of some Fig. 9 (continued). c. Patterns for an oolitic hematite-rich sample from the Sokoman Iron Formation and two hematite chert-rich samples from the Rapitan Group (Klein and Beukes, in press), a hematite-rich, granular, volcaniclastic iron-formation (field sample no. Y5-168.7) and a very hematite-rich, laminated and nodular lutite (field sample number Y7-142); from Klein and Beukes, 1992. Also shown is a curve for modem ocean water at 100 m depth (from Elderfield and Greaves, 1982) multiplied by lo6.
C. Kkin and N.J. Beukes
396
modern deep sea hydrothermal deposits, and Fig. 9c shows the fairly flat R E E pattern for a sample of oolitic hematite-rich iron-formation of the Sokoman Iron Formation, in which a positive Eu anomaly is still clearly present. For comparison, this same diagram also shows a pattern for modern North Atlantic seawater (at 100 m depth), obtained from Elderfield and Greaves (1982), which in contrast lacks the E u anomaly, and exhibits a considerably stronger negative Ce anomaly. The two iron-formation samples of the Rapitan sequence (about 0.7 to 0.8 Ga in age) are considerably different from any of the patterns in Fig. 9a or that given in Fig. 9c for the Sokoman iron-formation sample. A nodular and laminated hematite-chert sample (number 8) shows a pattern that completely lacks an E u anomaly with a general REE pattern distribution similar to that of the modern seawater pattern at 100 m depth. The other sample (number 7; granular with a volcaniclastic component) has an Eu anomaly, probably reflecting the proximity of this volcaniclastic iron-formation sample to hydrothermal source terranes, as in rift areas of modern oceans (Klein and Beukes, 1992). The patterns of Fig. 9a and c show a continuous decrease in the size of the positive E u anomaly from older to younger iron-formations, with the anomaly completely absent in the nodular iron-formation of Rapitan age. Such secular trends in REE in ironformations were first suggested by Fryer (1977, 1983). These are verified by Moeller and Danielson (1988) who, on the basis of chondrite-normalized REE patterns for iron-formations in the Hamersley Range, conclude that positive E u anomalies start to disappear in iron-formations of middle Early Proterozoic age. Very similar conclusions on the basis of the Nd and Sr isotopic compositions of a range of Proterozoic iron-formations are reported by Derry and Jacobsen (1988). They state that the transition from Archean mantle-dominated systematics to modern seawater behavior was established by approximately 0.7 to 0.8 Ga. It appears, therefore, that over time the hydrothermal component in Archean-Early Proterozoic ocean water decreased, and that at the time of deposition of the Late Proterozoic Rapitan iron-formation it was generally absent from the system. Further support for the notion of hydrothermal input into ocean waters from which iron-formations precipitated, can be obtained from the chemical distinctions between REE-rich deep sea sediments and REE-poor hydrothermal deposits such as studied by Bonnot-Courtois (1981) in the FAMOUS area of the Galapagos mounds. The hydrothermal deposits consist in large part of the Fe-rich layer silicate nontronite, whereas their bulk compositions are remarkably similar to those of typical BIFs. Figure 10 illustrates the nature of this distinction in terms of (Co + Ni Cu) abundances versus total REE content. All the iron-formation data points fall in or near the field of the hydrothermal deposits. In short, the striking similarities in the R E E patterns of iron-formations to seawater-dominated hydrothermal fluid mixtures (Fig. 9), and the trace element distribution plot in Fig. 10, suggest that similar processes with hydrothermal inputs have been responsible for the origin of iron-formations throughout Precambrian time. This conclusion is supported by sulfur (Goode et al., 1983) and Nd (Jacobsen and Pimentel-Klose, 1988) isotopic compositions of Archean and very early
+
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Proterozoic iron-formations, and Srg7/Srg6ratios in early Precambrian carbonates (Veizer et al., 1982). “ 0 depletion in chert and carbonates from this time period (Perry and Thn, 1972; Veizer and Hoefs, 1976; Perry and Ahmad, 1983; Beukes et al., 1990) may also be related to enhanced hydrothermal activity in the ocean system.
STRATIGRAPHY AND SEDIMENTOLOGY OF DEPOSITIONAL BASINS OF IRON-FORMATION
The stratigraphic sequences in which Precambrian iron-formations occur are highly variable (Gole and Klein, 1981) and show a wide diversity in lithological associations. In general, the stratigraphic associations of Archean iron-formations appear somewhat less variable than those of Proterozoic age, because volcanic rocks dominate in many Archean banded iron-formation sequences. In Proterozoic ironformations, volcanic rocks tend to be much less abundant, but materials of volcaniclastic parentage occur in the form of stilpnomelane-rich bands with still clearly recognizable shard structures (LaBerge, 1966a, b). They are described from two
398
C. Klein and MJ: Beukes
major iron-formations (both of which have undergone only very low-grade metamorphism and little tectonic deformation) for which the stratigraphy is extremely well known, namely, those of the Hamersley Range of Western Australia (Tiendall and Blockley, 1970) and of the Kuruman and Griquatown iron-formations in the Transvaal Supergroup of South Africa (see Fig. 12; Beukes, 1983). An evaluation of the sedimentary setting and subsequent interpretation of the original basins in which iron-formations were deposited is generally a much more difficult task than the measurement and compilation of the stratigraphy of an iron-formation sequence. Furthermore, iron-formation sequences commonly have been primarily studied by mineralogists, metamorphic petrologists, and economic geologists with lesser input from sedimentologists. The assessment of the original sedimentary setting of metamorphosed iron-formations, be they of Archean or Proterozoic age, is a difficult task. Major metamorphosed iron-formation sequences that show extensive regions of very low-grade metamorphic conditions are found in the Lake Superior region of the U.S.A. and Canada, and in the Labrador Trough sequence in Canada. Sedimentologic models for the Animikie Basin in the Lake Superior region are found in the work of James (1954) and extensive subsequent analysis by Morey (see Morey, 1983, for a review). Although a broad sedimentologic picture is now available for the Lake Superior region, Morey (1983) notes that “understanding the detailed sedimentologic history of a particular iron-formation in the basin is still an important objective”. The sedimentology of the iron-formations and associated lithologies in the Labrador Trough have been extensively studied by E. Dimroth and colleagues, by Zajac (see Gross and Zajac, 1983, for a review), and more recently by Simonson (1985a). Gross and Zajac (1983) provide a broad picture of chemical precipitation of ironformation in high-energy sedimentary environments of shallow troughs, layered basins and tidal flats. They do not support the replacement origin for these ironformations as proposed by Dimroth and Chauvel (1973). However, many details of the reconstructed sedimentary basins are uncertain. Although the stratigraphy of the Hamersley Range in Western Australia (Trendall and Blockley, 1970, and a review by Tiendall, 1983b), which has undergone only lower greenschist facies (burial) metamorphism, is extremely well documented, the sedimentologic reconstruction of this vast iron-formation basin is only in its infancy (Simonson, 1989). Trendall (1983b) notes that the “lack of detailed information on the lateral extent, thickness variations, and depositional environment of each unit of the Fortescue Group (which underlies the major ironformation sequences) precludes any definite conclusion concerning the tectonic development of the basin during deposition.” In terms of major, well-documented, only slightly metamorphosed and essentially non-deformed iron-formation sequences, the most complete stratigraphic and sedimentologic reconstruction of a basinal sequence is available from the Griqualand West region of the Transvaal Supergroup in South Africa (Figs. 11 and 12). This is almost totally based on the work of N.J. Beukes since 1973 (Beukes, 1983, 1984). This basinal sequence involves a facies transition from underlying lime-
399
Proterozoic iron-formations I
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Fig. 11. Geologic map showing the distribution of the Kuruman Iron Formation and the Korannaberg fold and thrust belt in South Africa. The unpatterned area represents Quaternaq cover (from Klein and Beukes, 1989).
stone, dolomite, and shale into an overlying sequence of siderite-rich to oxide-rich iron-formations. Because it has undergone only a very-low-grade metamorphic overprint (within an estimated temperature range of 110-170°C) and is essentially undeformed, it provides a unique opportunity for evaluation of the geochemistry of the various lithologies as a function of depositional basin constraints; this has led to much new information on chemical deposition processes and Proterozoic
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401
Proterozoic iron-formations SNAKE RIVER AREA m M A S S l V E MlXTlTE WITH I SHARP EROSIVE BASE
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Fig. 13. Sedimentologic setting of Rapitan Iron Formation at the top of the Sayunei Formation@ the Mackenzie Mountains and in the eastern Wernecke Mountains of the northeastern Cordillera of North America. The local unconformity in the lower part of the Shezal Formation is based upon data from Gabrielse et al., (1973) and Young (1976). Illustration after Klein and Beukes (in press).
seawater chemistry (Klein and Beukes, 1989; Beukes and Klein, 1990; Beukes et al., 1990; Kaufman et al., 1990), as discussed in a subsequent section. The very late Proterozoic iron-formations, among them iron-formations in the Rapitan Group of the Yukon and Northwest Territories, Canada (with an approximate age of about 0.7 to 0.8 Ga), those of the Morro do Urucum region of western Brazil and eastern Bolivia (approximately 0.8 Ga; Dorr, 1973a), and the iron-formations in the Damara Supergroup of Namibia (with an approximate age of 0.8 Ga; Kroner, 1981) all occur in stratigraphic sequences associated with deposits of glacial origin (see Fig. 13; Young, 1976; Yeo, 1986; Henry et al., 1986; Klein and Beukes, in press). The sedimentologic setting of these Late Proterozoic deposits is thus very different from all of the other Precambrian iron-formation sequences and their mineralogy tends to be mainly hematite and chert, with locally abundant manganese oxides.
REVIEW OF RECENTLY PUBLISHED IRON-FORMATION MODELS
Over the past 15 years a number of authors have specifically addressed the chemical and/or sedimentological aspects of iron-formation deposition. In
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C, Klein and N.J. Beukes
chronological order, these include: Cloud (1973); Eugster and Chou (1973); Holland (1973); Drever (1974); Button et al. (1982); Ewers (1983); Holland (1984); Garrels (1987); and Morris and Tiendall (1988). In most of these models, the chemical sediment (iron-formation) is precipitated from a relatively open stratified Precambrian ocean in which an upper oxic layer overlies a much larger volume of anoxic water. Garrels’ (1987) model, in contrast, is based on the evaporation of stream water in restricted basins. Garrels’ model explains the microbanded aspects of the iron-formations in the Hamersley region but because it is set in a restricted basin, the validity of the model has been questioned by Morris and Tiendall (1988). In most marine models the necessary concentrations of iron (and, possibly, silica and other oxide components) in solution are stored in a large volume of essentially anoxic deep water, which through upwelling comes into contact with more oxygenated shallower water. Where these two water masses interact is generally considered to be the depositional environment for the various types of iron-formation, with the oxygen content of the upper waters responsible for precipitation of iron hydroxides and/or oxides. In several models (e.g., Cloud, 1973), microorganisms are considered to have played a direct role in precipitation of the iron. Furthermore, Perry et al. (1973), Walker (1984, 1987), and Baur et al. (1985) explain the isotopically light carbonate carbon in iron-formation as resulting from the oxidation of organic matter, by a reaction such as: 6Fe2O3 + C + 4Fe304 + C02, in which primary hematite is converted to magnetite. There are several problems with this reaction mechanism: (i) magnetite appears commonly as an early phase in iron-formations (Klein, 1974; Klein and Bricker, 1977) and is the probable dehydration product of a precursor mixture of Fe(OH)2 and Fc!(OH)~; (ii) productivity of organic matter in the water column during iron-formation deposition may well have been low, as deduced from very low organic carbon values in BIF (see Fig. 15a; Klein and Beukes, 1989); and (iii) the scarcity of well-preserved organic remains (microfossils or stromatolites) in banded (non-peloidal) iron-formations (Walter and Hofman, 1983). In this regard, it should be noted that much of the biogenic interpretation of iron-formation of Laberge (1973), Laberge et al. (1987), and Robbins et al. (1987) is based on certain microstructures in BIF‘s that have been found to be non-biogenic (that is, dubiofossils or non-fossils) (Hofman and Schopf, 1983, p. 329, also photo 9-3; Mendelson and Schopf, 1992, p. 869). The fine-scale alternation of iron-rich and iron-poor microbands, mainly in iron-formations of the Hamersley Range, Western Australia and the Kuruman Iron formation sequence in South Africa, is interpreted as the result of deep water (below wave base) deposition linked with evaporation. The interpretation of microbands as varves (Trendall and Blockley, 1970; Trendall, 1973) led Garrels (1987) to develop a quite realistic chemical precipitation scheme for alternating hematite-chert-carbonate microbands. However, the evaporative origin proposed for such varves is not without problems as pointed out by Kaufman et al. (1990). These authors conclude that mixing of waters of different isotopic compositions
Proterozoic iron-formations
403
is a more likely mechanism for the production of primary isotopic compositions in microbanded Brockman Iron Formation of the Hamersley Range of Western Australia, than is evaporation. The non-microbanded iron-formations exhibiting granules, ooids, and cross bedding are considered to be shallow-water equivalents of the microbanded BIFs of deeper water origin. In general, in the above models for BIF deposition, the predicted paleoenvironment is based upon modern oceanic conditions and present-day shallow water carbonate deposition (Drever, 1974). They also have in common a lack of sedimentologic constraints on basinal configurations, because, in general, detailed sedimentologic interpretations of BIF sequences have not been available. Joint geochemical and sedimentologic studies of the iron-formation sequences in the Transvaal Supergroup by Klein and Beukes (1989), Beukes and Klein (1990), and Beukes et al. (1990) over the last seven years, have provided a much-needed correlation between detailed geochemistry and the well-established sedimentologic setting (see Beukes, 1983, for review). The following section provides a synopsis of these studies which lead to a new interpretation of the depositional setting of iron-formations based on the South African occurrences.
PALEOENVIRONMENTAL INTERPRETATION O F IRON-FORMATION DEPOSITION IN THE TRANSVAAL SUPERGROUP, SOUTH AFRICA
The geochemistry and sedimentology of the transition from underlying interbedded carbonate and shale to banded iron-formation in the Transvaal Supergroup ( ~ 2 . 4 3Ga) of South Africa is documented in detail by Klein and Beukes (1989). It consists of a transgressive iron-formation sequence (the Kuruman and Griquatown Iron Formations) on top of a stromatolitic carbonate platform sequence on the Kaapvaal Craton (see Figs. 11 and 12). The stromatolitic carbonates and associated shale beds represent regressive increments of sedimentation, whereas the iron-formation units were deposited immediately following transgressive events. At such times, siliciclastic input was inhibited and primary organic matter and carbonate production shifted shoreward (Klein and Beukes, 1989). This process of alternating regressive and transgressive sedimentation is well illustrated by geochemical signatures in the sequence. The shallow water stromatolitic limestones and associated carbonaceous shales have geochemical signatures that are distinct from those of the interbedded, deeper water, microbanded iron-formations. This is shown in Fig. 14 with the limestones and shales displaying REE patterns similar to those of modern shallow marine surface waters with terrigenous detrital influx, while the iron-formation patterns are similar to deep marine water (with no terrigenous input) with a pronounced admixed hydrothermal component, as evidenced by the positive Eu anomalies. A cross-plot of A 1 2 0 3 versus organic carbon contents (Fig. 15) further supports the concept of the iron-formation having been deposited from water
C.Klein and N.J. Beukes
404
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Fig. 14. Average REE patterns for various lithologies in the limestone to iron-formation transition in the Transvaal Supergroup, (a) normalized to the North American Shale Composite (NASC), and (b) normalized to chondrites (from KIein and Beukes, 1989).
masses devoid of detrital input, as opposed to that of the associated limestone and shale. The geochemistry and sedimentologic data support the concept of the iron-formations having been precipitated in deep marine environments, far removed from weathered continental sources. The iron, and probably silica as well, were derived from a hydrothermal source (as discussed in an earlier section). A further aspect of the studies by Klein and Beukes (1989), Beukes and Klein (1990), and Beukes et al. (1990) is the insight they provide into the relationship of primary production of organic [carbon] matter and iron-formation deposition. Figure 15a suggests that most of the organic [carbon] matter in the depositional basin was derived from the shallow, near-shore stromatolitic limestone/shale lithofacies with very little organic carbon reaching the deeper and distal parts of the iron-formation depositional system. This notion is further supported by the carbon-sulfur-iron relationship shown in Fig. 15b for the various lithofacies in the limestone to iron-formation transition. The organic carbon/sulfur ratios are very similar to the present-day open marine curve, suggesting that sulfate reduction was taking place at that time (Berner, 1984; Raiswell and Berner, 1985; Leventhal, 1987). The age of the transition zone (from limestone to BIF deposition) studied by Klein and Beukes (1989) conforms to that of a sequence in which Cameron (1982) established positive evidence for sulfate reduction, based on sulfur isotopic studies. Figure 15b also depicts the very low sulfur contents of the iron-formations
405
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which would correlate with low initial organic carbon contents. Other evidence presented by Klein and Beukes (1989) and Beukes and Klein (1990) in favor of low primary organic productivity in the locale of iron-formation deposition are: (i) high estimated rates of sedimentation for microbanded iron-formations which would have been favorable to the preservation of organic matter (Muller and Mangini, 1980) if it had originally been present; (ii) very low phosphorus and low barium contents in most iron-formations; and (iii) little preservation of organic carbon in siderite-rich iron-formations (see Fig. 15a), which must have been
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C.Klein and N.J. Beukes
deposited under reducing conditions. It is therefore concluded that iron-formation deposition and primary organic productivity were decoupled as first hinted at by Tbwe (1983). This is in contrast to models that couple deposition of iron-formation to either aerobic (Cloud, 1968; Button et al., 1982) or anaerobic (Hartman, 1984; Walker, 1987) photosynthesis. Decoupling of iron-formation deposition and primary organic productivity implies transfer of some oxygen from sites of its production (photosynthesis) to sites of the most oxidized (oxide type) iron-formation deposition. In turn, this would imply the presence of some oxygen in the atmosphere and surface ocean waters even in Archean times. If the Archean atmosphere and surface ocean were anoxic as suggested by Walker et al. (1983), Walker (1987), and Kasting (1987), dissolved ferrous iron would have been present and either through biotic (Cloud, 1968; Walker, 1987) or abiotic oxidation (Cairns-Smith, 1978; FranGois, 1986), oxide-rich iron-formations would have been deposited in shallow water environments. However, there is no geological evidence for this because typical Archean iron-formations are very similar in texture, bulk composition (major and trace elements), and mineralogy to the giant microbanded Early Proterozoic iron-formations (Gole and Klein, 1981; Klein and Beukes, 1989; see also Figs. 8 and 9) suggesting similar deep water environments of deposition. Archean shallow water chemical sediments, such as Early Archean stromatolitic cherts (Lowe, 1980, 1983; Walter et al., 1980; Byerly et al., 1986) and stromatolitic Middle (Mason and Von Brunn, 1977; Beukes and Lawe, 1989) and Late (Henderson, 1975; Martin et al., 1980) Archean carbonates, are concluded to have formed in the presence of oxygen-producing cyanobacteria (Schopf and Packer, 1987) and are depleted in iron. Especially in the Middle Archean Pongola Supergroup (Mason and Von Brunn, 1977) there is evidence that stromatolites were deposited in tidal marine environments (Von Brunn and Mason, 1977; Beukes and Lowe, 1989), which implies that Archean ocean surface waters were depleted in dissolved iron (Towe, 1983). This is similar to the Early Proterozoic environment for the deposition of iron-poor stromatolitic limestones in a shallow water platform and that of microbanded iron-formation under deeper basinal conditions (Fig. 12; Beukes, 1987). The interpretation of such conditions does not favor the coupling of iron-formation deposition to any photochemical process, be it biogenic or abiogenic. Possibly a more satisfactory model is one of a more or less permanentIy density-stratified ocean system with surface waters that were somewhat oxic and deeper waters that were anoxic and enriched in dissolved ferrous iron. We suggest a model in which there is interaction between two water masses; one of these is a nearshore water column of which the physical-chemical characteristics are largely determined by primary organic activity, and the other is an open marine system in which the deep water has a hydrothermal input (Fig. 16). In this model we stress the importance of the interaction between oxygen supply (probably from the atmosphere) and carbon supply (probably mainly from benthic organisms because the pelagic contribution may have been small) in the shallower, mixed surface
Proterozoic iron-formations
407 REGRESSIVE STAGE
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Fig. 16. Schematic depositional environment for iron-formation deposition and associated lithofacies in a marine system with a stratified water column in (a) a regressive stage, and (b) a transgressive stage (from Klein and Beukes, 1989). In (a) the photic zone reaches the floor of the deep shelf, allowing for cxyptalgalaminated limestone deposition. In (b) the photic zone is considerably above the floor of the deep shelf, causing the deposition of various iron-formation facies and chert. The thick arrows labeled C (carbon) in (a) represent high carbon productivity and supply, and the narrow arrows in (b) represent much less carbon productivity and supply. The vertical depth scale is based upon the basinal reconstruction of Klein et al. (1987).
water. We regard the deeper water as the main source of supply of iron (in solution as Fe2+) and probably of Si02, both of which originated mainly from a hydrothermal source. In contrast, the surface waters (above a depth of about 100 m) must have been essentially devoid of iron in solution. The rates of supply of oxygen and carbon to the interface between the surface and the deep waters was probably responsible for the various types (facies) of iron-formation. In Fig. 16a we depict a regressive stage with an abundant supply of carbon to the zone of mixing above the chemocline. The depth of the chemocline is arbitrarily set, in both Fig. 16a and b at 100 m, because this is the depth of mixing in the present oceans. In Fig. 16a the photic zone reaches down to the floor of the deep shelf; this is the region for precipitation of cryptalgal limestones for which the depth limit may have been about 45 to 50 m (Klein et al., 1987). The supply of carbon to the deeper waters is large at this stage and resulted mainly from benthic microbial mats that formed stromatolites in the carbonates; this provided the ultimate carbon source for the intraclastic carbonates (shown as limestone turbidities
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in Fig. 16a) and eventually for the carbonaceous shales as well. In this regressive stage, considerable precipitation of pyrite might have been expected. However, if SO4 contents were low in the Early Proterozoic ocean (Cameron, 1982; Melezhik, 1987), and/or if S04-reducing bacteria were not abundantly present, a source of H2S might not have been available for pyrite precipitation. Instead, Fe2+in solution might well have precipitated as part of layer silicates, which were subsequently converted to the Fe-rich chlorites found in many of the shales. In Fig. 16b we illustrate a transgressive stage in which the supply of carbon to the interface is much reduced, such that oxygen becomes available. Here the photic zone does not reach the bottom (and is shifted landward) and the input of siliciclastic material is less than in Fig. 16a. This results in oxic conditions and the deposition of hematite- and magnetite-rich iron-formation because there is little or nothing in the water column to reduce the primary Fe oxides or hydroxides. When, on the other hand, oxygen is in short supply at the water interface (chemocline), and some carbon is available, Fe2+ and PCO, build up to saturation levels, resulting in primary precipitation of siderite-rich iron-formation. This is especially so because conditions of limited carbon availability result in a lack of sulfate reduction, and favor primary precipitation of FeCO3 instead of FeS2. The various types of iron-formation probably resulted from variations in the supply of oxygen to the chemocline and may not reflect changes in water depth. As such, variations in carbon productivity and concomitant oxygen supply may well have resulted in siderite-rich iron-formation in one locale and oxide-rich ironformation at another locale, with both having formed at an essentially similar water depth. Figure 16b also suggests an explanation for the occurrence of chert-rich units that commonly separate shale or limestone from banded iron-formations. With very little iron in solution, and because of a lack of Si02-precipitating organisms, conditions (above the photic zone, or in the shallow waters) may lead to supersaturation with respect to Si02 in the lower surface waters, resulting in direct precipitation of chert-rich units. Another implication of decoupling primary organic productivity from ironformation deposition is that 13C depletion in the carbonates of iron-formations in general (Becker and Clayton, 1972; Perry et al., 1973; Baur et al., 1985), and of siderite-rich iron-formations in particular (Beukes et al., 1990), may not be the result of the degradation of organic matter. Siderite-rich ironformations in the Bansvaal Supergroup are depleted by about 4%0 in I3C with respect to their shallow water stromatolitic limestone equivalents which display a modern marine S13C value of close to zero permil (Fig. 1%; see also Beukes et al., 1990). This translates to a S13C difference of about 9%0 if effects of mineral fractionation (Golyshev et al., 1981) are taken into consideration. Kkin and Beukes (1989) and Beukes and Klein (1990) present petrographic and geochemical evidence indicating that many of the siderite microsparites are primary sedimentary precipitates formed in the water column. This is in contrast to the suggestion by Walker (1984) that siderite in iron-formations is the diagenetic
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product of a ferric oxide precipitate having reacted with organic matter as a reductant. The oxide-rich iron-formations in the Kuruman sequence consistently contain less carbon than do siderite-rich iron-formations (see Fig. lSa), suggesting that the latter could not have been derived from the former by making use of organic carbon as a reductant. This process is also suggested as the cause for I3C depletion in siderites (Walker, 1984), yet carbonates in the oxide-rich iron-formation are generally more depleted in 13C than those in siderite-rich iron-formations (Fig. 1%).Similarly, the isotopic composition of organic carbon in oxide-rich iron-formations is more enriched in 13C than that of siderite-rich units (and not the other way around as would be the case if the siderite had been derived from a reaction between hematite and organic matter) (Beukes et al., 1990). It is therefore concluded that the 13C depletion in siderite microsparites represents a primaIy signature of Early Proterozoic deep ocean water, and it appears that the ocean system at that time was most probably also stratified with regard to 13C composition of total dissolved carbon in contrast with present-day oceans (Schidlowski et al., 1983).
PALEOENVIRONMENTAL INTERPRETATION O F IRON-FORMATION DEPOSITION THROUGHOUT PRECAMBRIAN TIME
Paleoenvironmental models for iron-formation deposition at various stages in the Precambrian are shown in Fig. 17. The older iron-formations, ranging in age from Archean to Early Proterozoic, are thought to have been precipitated under similar conditions as a result of a stratified ocean system (similar to that depicted in Fig. 16) with predominantly deep water deposition of microbanded iron-formation, and a large but continually diminishing (with time) hydrothermal input into the oceans (Fig. 17a). In middle Early Proterozoic time, the stratified ocean system may have started to break down (Fig. 17b) with the development of abundant oolitic iron-formations such as those of the Lake Superior area (Goodwin, 1956; Shegelski, 1982; Simonson, 1985b; Ojakangas, 1988), the Sokoman Iron Formation in the Labrador Trough (Dimroth and Chauvel, 1973; Chauvel and Dimroth, 1974), and of the Nabberu Basin of Western Australia (Goode et al., 1983) in shoal areas. A mechanism for the transport of iron to surface ocean water must have developed. This may have been the result of a declining chemical density stratification due to lesser hydrothermal input as indicated by R E E results (Fig. 9). Following this period of time (ca. 1.9 Ga), the oceans may have become completely mixed, oxygenated, and depleted in iron (Fig. 17c) as indicated by the absence of iron-formations in late Early and Middle Proterozoic times (Fig. 1). In the Late Proterozoic, however, iron-formations are again part of the geologic record. These iron-formations are intimately associated with glaciomarine deposits and may also contain interbedded manganese deposits (Roper, 1956; Dorr, 1973b; Walde et al., 1981). In the Rapitan sequence, the iron-formation
C.Klein and N.J. Beukes
410 a. Archean-Early Proterozoic
Stratified ocean system. Hydrothemal input large. Production of organic matter low i n open marine environment.
b. Middle Early Proterozoic
------
Oolitic BIF
Hydro thermal input
Loss of stratification and diminished hydrothermal input.Deposition of oolitic hematite-rich B l F s on submerged platforms. Low productivity of organic matter.
c. Middle Proterozoic
Well mixed ocean system depleted i n iron Iron-formation absent.
d. Late Proterozoic
Stagnant
F e 2 7
Stage-1
'Snowball earth': Ice age with build UQ of F e z +
Stage-
Interglacial period: Downwelling o f 0 2 and oxidized BIF deposition.
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beds, composed essentially of hematite femicrite, occur immediately above planes of transgression at the base of progradational, shallowing-upward, glaciomarine sequences (Klein and Beukes, in press). A distinctive feature of the glaciations is that they apparently took place in near-equatorial settings at sea level as indicated by paIeomagnetic data (Frakes, 1979; Walter 1979) and abundant glaciomarine deposits (Yeo, 1986). This may call for a “snowball Earth” at certain periods of time (Kirschvink, 1992). A combination of all of these features may provide an explanation for the iron and manganese deposits of this period. In a “snowball-type” Earth situation, sea level is very low and the ocean highly stagnant such that reducing conditions can develop for accumulation of dissolved iron and/or manganese either from hydrothermal sources or from dissolution of material along basin floors. The onset of interglacial or postglacial stages results in transgressions, restoration of ocean circulation, and precipitation of iron and manganese-rich sediments at the base of prograding glacial meltwater deposits (Fig. 17d). Hydrothermal activity, generalIy thought to have been instrumental in the formation of the Late Proterozoic iron and manganese deposits (Yeo, 1986; Breitkopf, 1988), may then not have played such an important role. This is suggested by REE concentrations displaying patterns similar to modern deep marine water without any clear-cut hydrothermal signature in the form of a positive Eu anomaly (Fig. 9c).
ACKNOWLEDGEMENTS
Much of the above reported research by C.K. and N.J.B. was made possible by National Science Foundation grants EAR-8419161, EAR-845681, and EAR-861780.5 to CK N.J.B. acknowledges Rand Afrikaans University and the Foundation for Research Development, for several research allocations and support during a sabbatical leave in 1987-88. We are grateful to J.W. Schopf of the University of California, Los Angeles, and the Center for the Study of Evolution and the Origin of Life (CSEOL) for support during our respective research leaves in 1987-88. This manuscript was typed by Mabel T. Chavez and the illustrations were drafted by Dag Lopez.
Fig. 17. Paleoceanographic models for iron-formation deposition from the Archean to the Late Proterozoic. a. Archean to Early Proterozoic: stratified ocean system with predominantly deep water deposition of microbanded iron-formation (after Klein and Beukes, 1989). b. Middle Early Proterozoic: breakdown of the stratified ocean system and deposition of hematite-rich oolitic iron-formations. c. Middle Proterozoic: iron-depleted, well mixed ocean system with no deposition of iron-formations. d. Late Proterozoic: “snowball Earth”, with build-up of ferrous iron in solution in deeper water during glacial periods (Stage 1) and deposition of iron oxides during interglacial periods (Stage 2). (b, c, and d taken from Beukes and Klein, 1992.)
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Lesher, C.M., 1978. Mineralogy and petrology of the Sokoman Iron Formation near Ardua Lake, Quebec. Can. J. Earth Sci., 15: 480-500. Leventhal, J.S., 1987. Carbon and sulfur relationships in Devonian shales from the Appalachian basin as an indicator of environment of deposition. Am. J. Sci., 287: 33-49. Lowe, D.R., 1980. Stromatolites 3,400-Myr old from the Archean of Western Australia. Nature, 284: 441-443. Lowe, D.R., 1983. Restricted shallow-water sedimentation of early Archean stromatolitic and evaporitic strata of the Strelley Pool chert, Pilbara Block, Western Australia. Precambrian Res., 19: 239-283. Martin, A., Nisbet, E.G. and Bickle, M.J., 1980. Archaean stromatolites of the Belingwe Greenstone belt, Zimbabwe (Rhodesia). Precambrian Res., 13: 337-362. Mason, TR. and Von Brunn, V., 1977. 3-Gyr-old stromatolites from South Africa. Nature, 266: 47-49. Melezhik, V.A., 1987. Composition of waters of Precambrian basins as indicated by geochemical data. Int. Geol. Rev., 29: 1188-1199. Mendelson, C.V. and Schopf, J.W., 1992. Proterozoic and selected early Cambrian microfossils and microfossil-like objects. In: The Proterozoic Biosphere; a Multidisciplinary Study. Cambridge University Press, N.Y., pp. 867-951. Michard, A., AlbarPde, E, Michard, G., Minster, J.E and Charlou, J.L., 1983. Rare-earth elements and uranium in high-temperature solutions from East Pacific Rise hydrothermal vent field (13"N). Nature, 303: 795-797. Miyano, T, 1978. Effects of COz on mineralogical differences in some low-grade metamorphic ironformations. Geochem. J., 1 2 201-211. Miyano, T and Beukes, N.J.,1984. Phase relations of stilpnomelane, fern-annite and riebeckite in very low-grade metamorphosed iron-formations. Geol. SOC.S. Afr., Trans., 87: 111-124 Moeller, P. and Danielson, A,, 1988. Significance of Eu-anomalies in banded iron-formation. Geol. SOC. Am., Abstr. Progr., 2 0 381. Morey, G.B., 1983. Animikie Basin, Lake Superior Region, U.S.A. In: A.E Trendall and R.C. Morris (Editors), Iron-Formations: Facts and Problems. Elsevier, Amsterdam, pp. 13-67. Morey, G.B., Papike, J.J., Smith, R.W. and Weiblen, P.W., 1972. Observations on the contact metamorphism of the Biwabik Iron-Formation, East Mesabi District, Minnesota. Geol. SOC.Am., Mem., 13: 225-264. Morris, R.C. and Trendall, A.F., 1988. A model for the deposition of the microbanded Precambrian iron-formations. Am. J. Sci., 288: 664-669. Mueller, R.E, 1960. Compositional characteristics and equilibrium relations in mineral assemblages of a metamorphosed iron-formation. Am. J. Sci., 258 449-497. Muller, P.J. and Mangini, A., 1980. Organic carbon decomposition rates in sediments of the Pacific manganese nodule belt dated by uOThand 231Pa. Earth Planet. Sci. Lett., 51: 94-114. Ojakangas, R.W., 1988. Environments of deposition for lower Proterozoic Lake Superior type ironformation: Biwabik, Ironwood and Negaunee iron-formations, western Lake Superior region. Geol. SOC.Am., Abstr. Prog., 2 0 383. Perry, E.C. and Ahmad, S.N., 1983. Oxygen isotope geochemistry of Proterozoic chemical sediments. Geol. SOC. Am., Mem., 161: 253-264. Perry, E.C. and ?an, EC., 1972. Significance of oxygen and carbon isotope variations in early Precambrian cherts and carbonate rocks of Southern Africa. Geol. SOC.Am., Bull., 83: 647-664. Perry, E.C., Tan, EC. and Morey, G.B., 1973. Geology and stable isotope geochemistry of the Biwabik iron-formation, northern Minnesota. Econ. Geol., 68: 1110-1125. Raiswell, R. and Berner, R.A., 1985. Pyrite formation in euxinic and semi-euxinic sediments. Am. J. Sci., 285: 710-724.
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Robbins, E.I., LaBerge, G.L. and Schmidt, R.G., 1987. A model for the biological precipitation of Precambrian iron-formations, B. Morphological evidence and modem analogs. In: P.W. Uitterdyk Appel and G.L. LaBerge (Editors), Precambrian Iron-Formation. Theophrastus Publishers, pp. 97141. Roper, H., 1956. The manganese deposits at Otjosondu, South West Africa. 20th Int. Geol. Congr. Symp. Manganese, 2 115-122. Schidlowski, M., Hayes, J.M. and Kaplan, I.R., 1983. Isotopic inferences of ancient biochemistries: Carbon, sulfur, hydrogen and nitrogen. In: J.W. Schopf (Editor), Earth’s Earliest Biosphere: Its Origin and Evolution. Princeton University Press, Princeton, N.J., pp. 149-186. Schopf, J.W.and Packer, B.M., 1987. Early Archean (3.3-billion to 3.5 billion-year-old) microEossils from Warrawoona Group, Australia. Science, 237: 70-73. Shegelski, R.J., 1982. The Gunflint Formation in the Thunder Bay area. In: J.M. Franklin (Editor), Proterozoic Geology of the Northern Lake Superior Area. Field Trip 4. Geol. Assoc. Can., Winnipeg, pp. 15-31. Simmons, E.C., Lindsley, D.H. and Papike, J.J., 1974. Phase relations and crystallization sequence in a contact-metamorphosed rock from the Gunflint Iron Formation, Minnesota. J. Petrol., 15: 539-565. Simonson, B.M., 1985a. Sedimentology of cherts in the early Proterozoic Wishart Formation, QuebecNewfoundland, Canada. Sedimentology, 32: 23-40. Simonson, B.M., 1985b. Sedimentological constraints on the origins of Precambrian iron-formations. Geol. SOC.Am., Bull., 96: 244-252. Simonson, B.M., 1989. First discovery of ferruginous chertarenites in the early Precambrian Hamersley Group of Western Australia. Geology, 17: 269-272. Towe, K.M., 1983. Precambrian atmospheric oxygen and banded iron-formations: a delayed ocean model. Precambrian Res., u): 161-170. Trendall, A.F., 1973. Varve cycles in the Weeli Wolli Formation of the Precambrian Hamersley Group, western Australia. Econ. Geol., 6 8 1089-1097. Trendall, A.E, 1983a. Introduction. In: k E Trendall and R.C. Morris (Editors), Iron-Formations: Facts and Problems. Elsevier, Amsterdam, pp. 1-11. Trendall, A.E, 1983b. The Hamersley Basin. In: A.E Trendall and R.C. Morris (Editors), Iron-Formations: Facts and Problems. Elsevier, Amsterdam, pp. 69-123. Trendall, A.E and Blockley, J.G., 1970. The iron-formations of the Precambrian Hamersley Group, Western Australia, Aust. Geol. Surv., Bull., 119: 366 pp. Trendall, A.F., Compston, W., Williams, IS., Armstrong, R.A., Arndt, N.T, McNaughton, N.J., Nelson, D.R., Barley, M.E., Beukes, N.J., De Laeter, J.R., Retief, E.A. and Thorne, A.M., 1990. Precise zircon U-Pb chronological comparison of the volcano-sedimentaly sequences of the Kaapvaal and Pilbara cratons between 3.1 and 2.4 Ga. 3rd Int. Archean Symp., Perth, W.A., Extended Abstracts Volume, pp. 81-83. Veizer, J. and Hoefs, J., 1976. The nature of 018/016 and C13/C1* secular trends in sedimentary carbonate rocks. Geochim. Cosmochim. Acta, 40: 1387-1395. Veizer, J., Compston, W., Hoefs, J. and Nielson, H., 1982. Mantle buffering of the early oceans. Naturwissenschaften, 69: 173-180. Von Brunn, V. and Mason, T.R., 1977. Siliciclastic-carbonate tidal deposits from the 3000 Ma Pongola Supergroup, South Africa. Sediment. Geol., 18 245-255. Walde, D.H.G., Gierth, E. and Leonardos, O.H., 1981. Stratigraphy and mineralogy of the manganese ores of Urucum, Mato Grosso, Brazil. Geol. Rundsch., 70: 1077-1085. Walker, J.C.G., 1984. Suboxic diagenesis in banded iron-formations. Nature, 309 340-342. Walker, J.C.G., 1987. Was the Archean biosphere upside down? Nature, 329: 710-712.
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Walker, J.C.G., Klein, C., Schidlowski, M., Stevenson, D.J. and Walter, M.R., 1983. Environmental evolution of the Archean-Early Proterozoic Earth. In: J.W Schopf (Editor), The Earth’s Earliest Biosphere: Its Origin and Evolution. Princeton University Press, Princeton, N.J., pp. 260-290. Walter, M., 1979. Precambrian glaciation. Am. Sci., 67: 142. Walter, M.R. and Hofman, H.J., 1983. The paleontology and paleoecology of Precambrian ironformations. In: A.E Bendall and R.C. Morris (Editors), Iron-Formations: Facts and Problems. Elsevier, Amsterdam, pp. 373-400. Walter, M.R., Buick, R. and Dunlop, J.S.R., 1980. Stromatolites 3400-3500 Myr old from the North Pole area, Western Australia. Nature, 284: 443-445. Yeo, G.M., 1986. Iron-formation in the late Proterozoic Rapitan Group, Yukon and Northwest Tenitories. In: J.A. Morin (Editor), Mineral Deposits of the Northern Cordillera, Can. Inst. Min. Metall., Spec. Vol., 37: 142-153. Young, G.M., 1976. Iron-formation and glaciogenic rocks of the Rapitan Group, Northwest Territories, Canada. Precambrian Res., 3 137-158.
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Chapter 11
PROTEROZOIC COLLISIONAL AND ACCRETIONARY OROGENS B. WINDLEY
INTRODUCTION
During the Proterozoic two types of orogenesis were in operation. One involved collision of two or more continental blocks and gave rise, where collision was orthogonal, to orogens with thickened and commonly uplifted and eroded crust and to much reworking (by thrusting and partial melting) of older continental crust. These collisional orogens were associated with little or no crustal growth. The second type of orogenesis involved the growth and amalgamation of many juvenile island arcs and slices of oceanic crust or oceanic plateaus, which were in places mutually sealed by intervening accretionary wedges, and this process typically gave rise to orogens that contained little or no older crustal material. These accretionary orogens are characterised by considerable crustal growth; their crustal thickening, uplift and erosion were variable. This subdivision of Proterozoic orogens into two different types is important for the understanding of crustal evolution. The aim of this chapter is to describe key examples and to use these as a basis for evaluating the principal features of Proterozoic orogens in the light of current ideas of late Phanerozoic tectonics.
COLLISIONAL OROGENS (CO)
The Kola-Karelian orogen
This orogenic belt occupies the northern part of the Baltic Shield (Fig. 1). It contains five Archaean terranes: the Murmansk and Inari terranes consist of highgrade gneisses, whereas the Sorvaranger, Belomorian and Karelian terranes are composite, consisting of both low-grade greenstone belts and high-grade gneisses. In the period 2.0-1.9 Ga, these Archaean terranes collided and were amalgamated to form the Kola-Karelian orogen (Windley, 1991b). Early Proterozoic (2.4-1.9 Ga) rocks and structures added to these terranes include island arcs, Andean-type magmatic arcs, sutures, and remnant shelf successions. The Early Proterozoic structure of this orogen is well constrained by geophysical data of the POLAR Profile of the European Geotraverse (Gail et al., 1989; Marker, 1989).
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Rapakivi Granites Trans-Scandinavian Batholith Arc Volcanics Tanaelv belt (suture) Jatulian Shelf
A-A 6 -B C-C D D
-
Kola-Karelian Orogen Svecofennian Orogen Sveconorwegian Orogen Gothian Orogen
W
Fig. 1. Map of the Baltic Shield showing the four Proterozoic orogens, major Archaean terranes and other important geological features. B = Bergslagen; K = Kuopio; Ke = Keiv; L = Ladoga; LGB = Lapland Granulite-Gneiss Belt; Lu = Luled; N = Norbotten; 0 = Orijawi; R = Raahe; Sa = Savo; S = Skellefte; T = Brnpere. Modified after Windley (1991b).
The southern border of the Murmansk terrane is marked by a thrust zone several kilometres wide and dipping northwards at 60-80". It is marked by a prominent negative linear magnetic anomaly and current topographic depression, and it consists of biotite-bearing mylonitic gneisses within which there are remnants of metasedimentary and metavolcanic rocks, tectonic lenses of anorthosite up to 100 km long and 2.5 km wide, and ultramafic rocks; this is the KeivPorosozero suture, in which thrusts truncate the Early Proterozoic Keiv Group to the south. The Keiv Group in the former U.S.S.R. is a 2.0-1.9 Ga rifted, passive continental margin and overlying foredeep succession that passes upwards from sandstones, conglomerates and minor carbonates, to andesites and basalts overlain by dacites and rhyolites, that are succeeded by arkosic sandstones, greywackes and aluminous schists. The aluminous pelites are metamorphosed to biotite, garnet, staurolite and kyanite grades and thrust in recumbent nappes onto the craton margin. The Kola suture zone (Fig. 1) (the Polmak-Pasvik-Pechenga belt of Gas1 et al., 1989) is a south-dipping thrust zone up to 40 km wide that has placed the Inari terrane against and over the Sorvaranger terrane (Berthelsen and Marker, 1986; Marker, 1989). The borders of the suture zone are marked by mylonites and it contains at least two thrust-bound slices made-up of the 2.4-
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2.0 Ga Pechenga Series that contain sediments ranging from conglomerates, sandstones (rifted continental margin), through stromatolitic dolomites, siltstones and phyllites (shelf-rise transition), to turbidites (trench), and volcanic rocks that pass upwards from andesites through alkalic basalts to tholeiitic basalts (with REE characteristics resembling those of MORB) to 1.99 Ga ferropicrites. On the south side of the suture zone there is a thrust-bound, greenschist-grade Early Proterozoic island arc sequence that consists of weakly deformed abundant andesites, basaltic pillow lavas, minor komatiitic lavas, tuffs, and sulphide-bearing carbonaceous pelites (Marker, 1985; Berthelsen and Marker, 1986; GaAl et al., 1989). A further result of the southward subduction that gave rise to the island arc was emplacement into the northern border of the Inari terrane of an Andean-type magmatic arc represented by 1.95-1.9 Ga calc-alkaline plutons (Barbey et al., 1984). The Lapland Granulite/Gneiss Belt is 80 km wide, and consists largely of paragneisses derived from turbidites. Its tectonically lower part is highly sheared, and in its upper part metasedimentary granulites contain conformable lenses of noritic-enderbitic, calc-alkaline orthogneisses (Hormann et al., 1980; Barbey and Cuney, 1982). Isotopic ages of meta-igneous granulites suggest emplacement and metamorphism at 1.9-2.0 Ga and post-tectonic crustal melt granites are dated at 1.8 Ga (Merilainen, 1976; Bernard-Griffiths et al., 1984). The granulite belt of Finland continues eastwards as amphibolite-facies gneisses on the western side of the White Sea, where it occurs north of rocks characteristic of the Thnaelv belt (Fig. 1). The turbidite precursors of the Lapland Granulites may have been deposited in a back-arc basin adjacent to a magmatic calc-alkaline arc (Berthelsen and Marker, 1986), or on a passive continental margin subsequently intruded by arc magmas (Barbey and Raith, 1990). A further possibility is that the turbidites accumulated in an accretionary wedge. The Tanaelv high-grade ductile thrust belt contains intercalated tholeiitic amphibolites and calc-alkaline gneisses, within which there are many lenses of gabbroic, anorthositic and ultramafic rocks including some eclogitic and high-pressure metamorphic rocks. A combination of reflection seismic, gravity, electromagnetic and magnetic data of the POLAR profile shows that the Granulite and Tanaelv belts form a north-dipping thrusted wedge, the apex of which continues to a depth of 15-20 km (Marker, 1989). This wedge is made up of a stack of high-grade thrust sheets on top of a major basal dkcollement. As a result of 1.9 Ga thrusting of a hot slab of Lapland granulites over a cooler slab of Belomorian gneisses, the Thnaelv belt contains an inverted metamorphic sequence (Krill, 1985). Barbey et al. (1984) and Berthelsen and Marker (1986) interpreted the Thnaelv belt as a suture zone containing relict oceanic crustal material located on the site of a closed back-arc basin, that may have been about 125 km wide. The Sirkka Thrust is a major tectonic boundary along which the high-grade Belomorian terrane has been thrust southwards under the low-grade Karelian terrane. The boundary is marked by mylonites and refraction seismic data show that it dips ca. 40"s (Luosto et al., 1989). There is no positive evidence that
422
B. Windley
this is a suture; more likely it is a post-collisional thrust. The Karelian terrane has considerable evidence of rifting and shelf deposition in the Early Proterozoic when it constituted the foreland or rifted passive continental margin of the Svecofennian ocean to the south (GaA1, 1986). There are 2.44 Ga mafic-ultramafic layered intrusions, many 2.4-2.3 Ga and 2.2-2.0 Ga NW-trending (coast-suture parallel) basic dykes (Gorbatschev and GaB1, 1987), and 2.5-2.3 Ga conglomerates and tholeiitic volcanics (continental rift facies). After erosion and weathering, a continental platform developed with deposition of the Jatulian Group of basal quartzites, shelf carbonates (2.05 Ga, Pb-Pb) and banded iron formations (2.08 Ga, Pb-Pb), followed by mafic volcanics, carbonaceous slates and black schists which are the analogue of the deep water facies starved of terrigenous debris typically formed during the break-up of a modern carbonate platform. In summary, the Early Proterozoic collision tectonics in the Kola-Karelian orogen had markedly different effects on the five Archaean terranes. The Grenville The 4000 km long Grenville orogen in N. America (Moore et al., 1986 and papers therein; Rivers et al., 1989) is a small segment of a 13000 km long collisional orogen that was probably responsible for joining many continental blocks into a Late Proterozoic supercontinent (Hoffman, 1991). The history of the Grenvillian Wilson Cycle in North America started with the prominent 1.48-1.43 Ga anorogenic magmatism in Canada (especially anorthosites) and in the central/southern USA - mostly rhyolitic ashfall tuffs and peraluminous granites (Van Schmus et al., 1987); this magmatism most likely developed on the continental margin of the Grenvillian Ocean and modern analogues border the Atlantic Ocean (Windley, 1989). Closure of the ocean by subduction is indicated by the 1.28-1.25 Ga island arc of the Central Metasedimentary Belt (CMB) of Ontario, and by an island arc associated with an incomplete ophiolite in Texas which was thrust northwestwards onto a foreland and shelf (Garrison, 1981). Collision of the CMB with adjacent continental blocks gave rise to the 1.25-1.22 Ga Elsevirian orogeny and the 1.12-1.03 Ga Ottawan orogeny (Easton, 1986). The result of these orogenies was the formation of a collisional orogen which Rivers et al. (1989) divided into three zones (Fig. 2). The northwestern zone is paraautochthonous and contains rocks of the adjacent foreland reworked by Grenvillian deformation and metamorphism. The central zone contains amphibolite- and granulite-grade allochthonous gneisses developed from earlier continental material. The southeastern zone contains several terranes, some of which are juvenile, like the arc of the CMB. Other zones are older terranes accreted during the Grenvillian orogenies. The boundaries between the three zones are major thrusts, which transported all rocks and zones towards the northwest. The thrusting gave rise to a thickened crust, which underwent both uplift and erosion, and gravitational collapse and formation of 1.01-0.935 Ga mylonitic normal faults (van der Pluijm and Carlson, 1989). The northwest-directed deformation caused by the terminal Ottawan orogeny fractured the foreland giving
Pmterozoic collisional and accretiona y orogens
423
Northwestern Zone
Southeastern Zone and individual Terra
Fig. 2. The Grenville orogen in SE Canadian Shield showing the three internal zones discussed in the text. Based on Rivers et al. (1989) and modified after Hoffman (1989).
rise to the 1.1 Ga Keweenawan rift (Gordon and Hempton, 1986), that in origin is comparable to the Rhine graben caused by the Tertiary Alpine deformation in
Europe (Windley, 1989). The Grenville orogen continues in the SW Baltic Shield as the Svecononvegian orogen (Fig. l), although it is uncertain whether the latter belongs to the northwestern (Gower, 1985) or southeastern part of the former (Hoffman, 1991).
Wopmay and Thelon orogens, NW Canada Wopmay (Fig. 3) is a short-lived, 1.95-1.84 Ga orogen in NW Canada (Hoffman and Bowring, 1984) that developed as a result of the collision between the Archaean Slave Province and an unknown Nahanni continental block to the west; a small island arc, the Hottah (that was built offshore on a 2.3-2.1 Ga crust) was trapped between the colliding blocks (Hoffman, 1988, 1989). The western rifted margin of the Slave Province is overlain by shelf-rise sediments of the westward-facing Coronation Supergroup and succeeded by an eastward-migrating foredeep that formed in a late thin-skinned thrust-fold belt. The shelf began to collapse at 1.97 Ga, and collided at 1.91-1.90 Ga with the 1.95-1.91 Ga Hottah arc as a result of westward subduction below the arc. A new dextral-oblique, east-dipping subduction zone developed on the west side of the accreted arc and led to generation of the 1.88-1.86 Ga Great Bear calc-alkaline batholith, partly on top of the Hottah arc and partly on the deformed continental margin to the east. 1.86-1.84 Ga syenogranites were generated in the metasediments of the deformed shelf-rise. Xrminal collision at about 1.8 Ga of the Nahanni terrane in the west led to formation of the postulated Johnny Hoe suture, indicated by gravity and magnetic highs, and to a set of conjugate transitional faults in the Slave foreland in the east.
B. Windley
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116"
1120
108"
104"
100'
Fig. 3. The northwestern Canadian Shield showing the Archaean Slave Province bordered on the west by the Woprnay orogen (containing the Hottah Arc, the Great Bear Arc and the Coronation Supergroup) and on the east by the Thelon orogen (containing the Thelon-Taltson Arc), and the Tibetan-type plateau of the Queen Maud Uplift. After Hoffman (1988).
Thelon is a 2.02-1.91 Ga orogen (Fig. 3) that developed as a result of collision between the Slave Province and the Rae Province (Hoffman, 1987a, 1988). Eastward transport of the former was facilitated by 600-700 km dextral slip along the Great Slave Lake Shear Zone and an eastward-dipping subduction zone that led to formation of the narrow (80 km) Bltson-Thelon composite batholith (equivalent to the Kangdese batholith of the Himalayas), which is partly precollisional and calc-alkaline and partly post-collisional and anatectic. The resulting Thelon suture is cryptic. Extensive post-collisional shortening gave rise to crustal thickening in a 300 km wide Tibetan-type plateau of the Queen Maud Uplift.
Australian orogens In SW Australia there are four narrow Proterozoic collisional orogens (Fig. 4). The analysis of Myers (1990) is based on an understanding of the distribution and age of major rock units and tectonic belts across the orogens from foreland to hinterland, and evaluation of these in terms of an expected distribution in space and time across a modern collisional orogen. The Capricorn orogen (Fig. 4A) resulted from collision of the Pilbara and Yilgarn cratons between 2.0 Ga and 1.6 Ga. On the northern side there is early sandstone and basalt, comprising a northward-transported foreland fold-and-thrust belt associated with a developing
425
Proterozoic collisional and accretionaiy orogens 20"
25"
. . . . . . . . . . . . . . ..
I
I( 1I! II I{
115"
120"
BRESNAHAN BASIN c.lM)OMa
Molasse
ASHBURTON BASIN Low Melamorphic grade ,2 ~2000-1700Ma
Passive margin Deposits 8 Overlying Foreland basin
[n
Sedimentary rocks
Volcanic, Sedimentary a Plutonic rocks of Island arc, Back arc or Oceanic Orioin
NABBERU BASIN Low Metamorphic grade c.2oOO-1700 Ma
Granite 1800-1600Ma Moderate Metamorphic grade
~
-
ARCHEAN CRATON 2 4 0 0 Ma
I
Normal Fault
'
InterleavedArchean Basement 8 Early Proterozoic sedimentary cover
0
t;;;;
Strike-Slip Fault 115'
120"
125"
Fig. 4. The Proterozoic orogens of southwestern Australia. A. Capricorn orogen, 2.0-1.6 Ga; B. Pinjarra and Albany-Fraser orogens, 1.3-1.0 Ga; C . Patterson orogen, 750-550 Ma. After Myers (1990).
foreland basin. On the south side of an inferred south-dipping suture zone there is a 1.8-1.6 Ga granitic batholith, and a central zone of interthrusted Archaean basement and Early Proterozoic sedimentary cover. On the southern side of the orogen, shelf deposits of the Nabberu basin are overlain by oceanic-type basalts and greywackes representing obducted arc or oceanic crust, and a southwarddirected thrust belt. The Albany-Fraser and Pinjarra orogens (Fig. 4B) along the western and southeastern borders of the West Australian craton respectively formed in the period 1.3-1.0 Ga and the more well known Albany-Fraser orogen has a symmetrical structure. There is an early basic dyke swarm in the craton on the northern side
426
B. Windley
that was probably parallel to the continental margin. Clastic sediments deposited on the craton were deformed in a fold-and-thrust belt, and a marginal zone of tectonic slices of lower crustal gneiss and gabbro was thrust northwestwards towards the craton. A zone of high-grade sediments, granites and gneisses occurs on the southeastern margin of the belt. The Patterson orogen on the northeastern side of this craton (Fig. 4C) formed between 750 Ma and 550 Ma. A dolerite dyke swarm emplaced into the West Australian craton at a high angle to the craton represents the root of a failed aulacogen arm. The orogen has a symmetrical structure, and is dominated by craton-directed thrusts that deformed foreland basins on either side. A granite in the northern part of the orogen has an age of 600 Ma. From Myers’ (1990) study it is most likely that modern-style collision tectonics were in operation in Australia throughout the Proterozoic. There is currently a controversy about the role of plate tectonic models in the formation of Australian Proterozoic orogens. Etheridge et al. (1987) and Wyborn (1988; based on geochemical data) proposed an ensialic, non-actualistic model to explain the evolution of deformed sedimentary basins during the widespread 1.88-1.85 Ga Barramundi orogeny. The following are essential features. Basins, that resulted from extension of preexisting Archaean continental crust, contain three sequences. A lower rift sequence is overlain by a laterally extensive shallowwater sequence ascribed to post-rift thermal subsidence, which is succeeded by thick flysch deposits, heralding the main orogenic phase. The orogeny included intensive thrust-nappe deformation, low pressure (andalusite-sillimanite) isobaric metamorphism, and the intrusion of I-type granitoids and comagmatic felsic volcanics, that formed from mantle-derived 2.3-2.0 Ga sources, and that contain no evidence of Archaean components. These characteristics, together with the lack of evidence of ophiolites, sutures, paired metamorphic belts and ‘other diagnostic features of a modern orogeny suggest that the Barramundi orogeny was essentially ensialic’. The above authors proposed a model in which a polygonal array of upwelling convection zones triggered mantle melting, and consequent crustal underplating. Mantle cooling was responsible for crust-mantle delamination and A-subduction. The delamination provided the driving forces for the compressional orogeny and the enhanced heat flow necessary to produce the felsic magmatic rocks by melting of the underplated layer. Fundamental problems exist with this innovative, but unrealistic model. SmNd isotopic data require a difference of 400 Ma between the underplating and the magmatism, and it is difficult to appeal to a single mantle convection system to account for both (McCulloch, 1987). Sm-Nd data also indicate that crustal growth was voluminous at 2.9 Ga, from which Windrim and McCulloch (1986) suggested that deformation and granulite-facies metamorphism at 1.8 Ga were a consequence of ‘collision between continental blocks’. Furthermore, some sediments are not situated in basins, but in tectonic synforms; tectonic thickening by thrusting in the Harts Range gave rise to crust-mantle melting and intrusion of granites in a Cordilleran-style fold-and-thrust belt on a continental margin (Ding and James, 1985). The question has to be asked: did Etheridge et al.
Proterozoic collisional and accretionaiy orogens
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(1987) and Wyborn (1988) make comparison with the correct equivalents in a modern orogen? If not, their conclusions cannot be sustained. It seems to me that the most likely equivalents are not in modern orogens with sutures, but in the thrusted Jurassic-Palaeogene basins in the Tien Shan range of Central Asia, that is situated about 2000 km away from its deformation front in the Himalayas of Pakistan (Windley et al., 1990; Allen et al., 1991). These basins have a similar tripartite sequence (rifting, thermal subsidence, deposition of coarse clastics) as the Australian basins, and they likewise never opened to become oceans. Postcollisional (Himalayan), Neogene-Quaternary thrusting has highly imbricated the basins with their basement; this thrusting has given rise to a thickened crust, a consequent high (>7000 m) mountain range, erosion, and formation of molasse basins that are in turn being thrusted today, and thus the modern Tien Shan range has no ophiolites, no suture, no Andean-type magmatic arc, and no paired metamorphic belts. Melting of mantle-crust and/or underplated material can be expected beneath such thrust-thickened crust, but obviously it is too young to have given rise yet to granites at the present surface. This model is actualistic and realistic.
ACCRETIONARY OROGENS (AO)
In the Proterozoic numerous orogens formed by the growth and amalgamation of island arcs, oceanic plateaus and accretionary wedges. Prominent is the 1.9-1.6 Ga zone that extends from the Baltic Shield, through South Greenland, Michigan, Colorado to Arizona (see Condie, this volume), that Patchett and Arndt (1986) showed from Nd isotopes consists of >80% newly differentiated, subductionrelated material. This zone contains three belts of A 0 that decrease in age from north to south (Fig. 5): (1) 1.9-1.8 Ga. The Svecofennian of the Baltic Shield (Park, 1991), the Ketilidian of South Greenland (Allaart, 1976), the Makkovik of Labrador (Gower and Ryan, 1986), and the Penokean in the Lake Superior region (Barovich et al., 1989). (2) 1.8-1.7 Ga. The Killarney belt near Lake Huron (van Breemen and Davidson, 1988), the Central Plains orogen (Sims and Peterman, 1986), and the Yapavai Province of Colorado and Arizona (Karlstrom and Bowring, 1988). (3) 1.7-1.6 Ga. The Trans-Scandinavian batholith of Sweden (Park, 1991), the Labradorian orogen (Thomas et al., 1986), and the Mazatzal Province of New Mexico and Arizona (Karlstrom and Bowring, 1988). In the formation of these belts, an aggregate area of new crust up to 1500 km wide and 5000 km long was accreted in about 300 Ma (Hoffman, 1988). For this chapter the Svecofennian, Ketilidian and Penokean orogens are selected for description, as well as the Early Proterozoic Birimian orogen in West Africa. Orogens of SW North America are described in Chapter 12. In the Late Proterozoic the Pan-African of the Arabian-Nubian Shield formed by similar processes.
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B. Windley
Fig. 5. The Precambrian of Laurentia showing the three belts of juvenile arc-accretionary orogens that extend from California to the Baltic Shield and decrease in age from north to south. Modified after Hoffman (1988).
The Svecofennian orogen
There are more data available from this key Early Proterozoic arc-accretionary orogen than most others and therefore it will be described in detail (Fig. 1). It contains no Archaean terranes and little or no Archaean isotopic material (Wilson et al., 1985; Patchett and KOUVO, 1986; Huhma, 1987; Patchett et al., 1987; Romer, 1991). It developed by the growth and collision of 1.9-1.8 Ga juvenile arcs (Park, 1991), and by extensive crustal melting in the period 1.8-1.55 Ga. Extending from Central Finland and Sweden southwards to beyond Estonia as far as the Tornquist Zone in Poland this orogen is at least 1200 km wide. The LuleA-Kuopio suture zone (named after the eponymous towns) is a thrust (Fig. 1) that separates the Kola-Karelian orogen from the Svecofennian orogen. It contains thrust slices of two types of Kalevian turbidites: (a) autochthonous turbidites that contain Archaean and Proterozoic detritus, and are locally interbedded with tholeiitic volcanics - Ward (1987, 1988) suggested deposition in transtensional, intra-cratonic rifts near the craton margin; (b) allochthonous turbidites that Ward (1987, 1988) suggested were deposited from debris flows and turbidity currents in submarine canyons on an accretionary margin. Also in the suture zone is the thrusted Outokumpu nappe that contains serpentinites, gabbros, basaltic pillow lavas, non-detrital quartzites, dolomites, Mg-rich meta-volcanics and Cu-sulphide deposits (Park, 1984). The 1.96 Ga Jormua ophiolite (Kontinen,
Proterozoic collisionaland accretionary orogens
429
1987) was thrust about 30 km onto the continental margin of the Karelian terrane (Fig. 1). It contains an upward sequence of serpentinites; gabbros cut by basic dykes; a sheeted dyke complex with screens of gabbro and serpentinite; basaltic pillow lavas and pillow breccias; tuffites, cherts and carbonate sediments; and turbiditic greywackes and semi-pelites. Thrust nappes have imbricated Karelian basement and its Jatulian cover and they have transported Kalevian turbidites and accretionary prism rocks northwards over the Achaean craton (Park and Bowes, 1983). There are several magmatic arcs within the Svecofennian orogen. In North Sweden the Norrbotten arc (1.90-1.87 Ga) consists of porphyritic intermediate and felsic lavas that resemble modern high-K, calc-alkaline arc lavas (Pharaoh and Brewer, 1990). Along the Svecofennian side of the suture is the NW-SE trending, ore-rich Raahe-Ladoga fault zone (Fig. 1) (Korsman, 1988) within which there are 1.91-1.90 intermediate-acidic lavas associated with base metal mineralisation, 1.9-1.88 Ga gabbros with nickel ores, and 1.89-1.875 Ga granites and gabbros (Vaasjoki and Sakko, 1988). The rocks and ores are comparable to those in modern island arcs and intra-arc rifts. Late faulting is often concentrated along hot, young arcs. The 1.89 Ga Skellefte island arc has mature intra-arc volcanics (Vivallo and Claesson, 1987) and granodiorite-granite intrusions derived from subduction-related melts (Wilson et al., 1987). Magnetotelluric transfer functions along the FENNOLORA profile show that there is an extremely high conductivity and low resistive anomaly across the Skellefte arc-suture (Rasmussen et al., 1987). The BABEL seismic profile in Sweden shows a major N-dipping thrust that may be related to a subduction zone (BABEL, 1990). The arc-suture zone continues eastwards around a 500 km long arcuate loop outlined by a major geoelectric conductivity anomaly up to 20 km wide to join the 1.9 Ga llimpere island arc (Hjelt, 1991), that contains medium-K basalts, medium/high-K andesites, and high-K dacites and rhyolites; these calc-alkaline volcanic rocks resemble those in Recent mature island arcs (Kiihkonen et al., 1989). The structure of the Bmpere arc indicates that the derivative subduction zone dipped southwards (Nironen, 1989b). The Savo schist belt contains several highly metamorphosed and deformed arcs and granite plutons were emplaced at 1.88-1.89 Ga (Nironen, 1989a). In SW Finland the Orijarvi arc comprises submarine alkaline-subalkaline basic-intermediate lavas, felsic pyroclastic rocks associated with massive Cu-PbZn deposits, co-magmatic gabbro-tonalite bodies and a 1.891 Ga (U-Pb on zircon) granodiorite, all possibly formed in a back-arc environment (Colley and Westra, 1987). Along strike to the west are arc-type, mafidintermediate to rhyolitic volcanic rocks (Ehlers and Lindros, 1990). At Bergslagen in southern Sweden (Baker and Hellingwerf, 1988) there is a >lO-thick pile of 1.9-1.8 Ga felsic pyroclastics, ignimbrites and immature sediments that probably developed in major intra-arc or back-arc rifts. U-Pb zircon data indicate that many of the Svecofennian arc lavas were erupted in the short period of 1.92-1.87 Ga contemporaneously with the intrusion of innumerable 1.91-1.86 Ga, subductionderived plutons and batholiths of tonalite, granodiorite and granite (Nurmi and
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Haapala, 1986; Welin, 1987). Many volcanic rocks are associated with massive Cu-Zn-Pb-Ag-Au mineralisation, and the granites with porphyry-type Cu-Mo-Au deposits (GaB1, 1990). Between many of the Svecofennian arcs there are biotite-bearing granitic gneisses and schists that have been widely regarded as metagreywackes and metapelites within which there are many conformable km-size lenses of amphibolite, metagabbro and metaultramafic rocks that contain important Ni-Cu deposits. Rather than intrusions into precursor sediments made conformable to the derived gneisses by deformation (see Papunen and Gorbunov, 1985 and references therein), they are more likely original slices of ocean floor that were thrust into the pelites and greywackes of accretionary wedges. Following amalgamation of the arcs, syn- and post-collisional deformation took place. Thrusting and folding was associated with high amphibolite facies metamorphism that locally reached granulite grade. Rb-Sr data suggest that late kinematic activity ceased by about 1.77 Ga (Welin and Stalhos, 1986). Crustal thickening led to the formation of abundant crustal melt granites, of which there are three main types (Nurmi and Haapala, 1986). (1) The peak of regional metamorphism gave rise to partial melting of paragneisses and formation of 1.85-1.80 Ga, water-saturated, near-minimum granites as in situ plutons and dyke networks. These are comparable to the Miocene leucogranites of the High Himalayas (Crawford and Windley, 1990). (2) Deep-crustal dehydration melting promoted by mantle input led to nonminimum melt, 1.80 Ga granites in far-travelled plutons associated with ultrapotassic, mantle-derived lamprophyre dykes. Potassium in the mantle was derived from the breakdown of phlogopite that originally formed in the mantle wedge above the earlier subduction zone. These granites are similar to those in the Miocene Karakoram batholith in the upthrust western end of the Tibetan plateau (Crawford and Windley, 1990). (3) By 50-200 Ma after the last tectonic activity, internal slow heating of the thickened crust had led to its final extension and collapse, and thus to decompression melting of the mantle and melting of depleted granulitic lower crust left over by removal of type 2 magmas, and the result was the formation of 1.7-1.55 Ga rapakivi granites and coeval gabbros, anorthosites and basic dykes (Haapala and Ramo, 1990). The last event in the history of the Svecofennian orogen was the deposition (maximum age, 1.5 Ga) of Jotnian sandstones in a major elongate basin below the Gulf of Bothnia (Fig. 6; 1-2), which locally extends into exposed basement as Jotnian rifts associated with mafic dykes. This basin was the final result of the extension and thinning of the Svecofennian crust; incipient modern analogues are the Quaternary rifts of Tibet.
The Ketilidian This orogen in South Greenland (Bridgwater et al., 1973; Allaart, 1976) is an incomplete segment of an Early Proterozoic accretionary orogen with an Andean/
43 1
Proterozoic collisional and accretionary orogens I
I
I
I
I
I
48"
46"
440
Fig. 6. Map of the Ketilidian orogen in southwest Greenland. A-B is line of section shown in Fig. 7. After Windley (1991).
B
A KETlLlDlAN OROGEN NW
Kobberminebugt Suture Zone
_---__
BaCk.~rC Shear Belt
1800 Ma Thrust- NaPPe stack
SE
*+++++++++++++++++
i+++++++++++++++++ i+++++++++++++++++
Andean-Type Batholith 1850-1800 Ma Granites with significant proportions of Crustal -Melts
Fig. 7. Schematic NW-SE section across the Ketilidian orogen (Fig. 6). See text for discussion. After Windley (1991).
Himalayan-type collisional margin (Windley, 1991a). In a N to S section there are the following zones (Figs. 6 and 7). (1) A northern foreland of Archaean gneisses is overlain unconformably by a shelf-foredeep succession, several kilometres thick. The basal Vallen Group consists of stromatolitic dolomites and quartzites laid down on a stable shelf, overlain by flysch-like greywackes and siltstones deposited by turbidity currents
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into basins on the deepening shelf, and a 30 m thick sulphide-facies iron formation (chert-pyrite-shale), similar to that which commonly occurs on the outer ramp of several Early Proterozoic foredeeps (Hoffman, 1987b). The overlying Sortis Group contains mostly tholeiitic pillow lavas and basic-felsic pyroclastics (Bondeson, 1970); the axial zones of several Early Proterozoic foredeeps are unusual in containing volcanic rocks (Hoffman, 1987b). The above succession is spectacularly deformed and metamorphosed progressively (to sillimanite grade) southwards over 50 km towards the suture zone of the orogen (Windley et al., 1966; Henriksen, 1969). It has been thrust northwards over the foreland and back-thrust near the suture where it and the basement thrusted gneisses are intruded by several 1.7751.675 Ga granites that contain appreciable crustal-melt components (Kalsbeek and lhylor, 1985). Most of the above relations are comparable to those that occur in the deformed foreland of modern collisional orogens such as the Himalayas (Searle et al., 1987). In these respects the continental margins of an A 0 are similar to those of a CO. (2) The Kobberminebugt suture is a 15 km wide vertical shear zone that contains relict greenschist-grade pillow lavas and gabbros, copper and gold mineralisation, and late mylonite zones 100 m thick (Harry and Oen, 1964; Ghisler, 1968). The Julianehaab batholith is a 80-100 km wide Andean-type batholith that consists of tonalites, granodiorites and granites (with a low initial Sr isotope ratio of 0.7032) that contain relicts of pillow lavas, pyroclastic rocks and extensive noritic gabbros (Allaart, 1976) that probably belong to an early island arc into which the major calc-alkaline batholith was intruded (Windley, 1991a). The arc rocks are similar to those in the Kohistan arc in the Himalayas of North Pakistan, the lower part (magma chamber) of which is occupied by the Chilas Complex of noritic gabbros (Khan et al. 1989; Petterson and Windley, 1991a). (3) The southernmost part of the Ketilidian orogen consists largely of metamorphosed supracrustal rocks; Nd isotope ratios indicate that they were derived from new mantle-derived material and not from Archaean continental crust (Patchett and Bridgwater, 1984). The rocks occur in three sub-horizontal thrust nappes (Escher, 1967). The uppermost contains a 2.5 km thick succession of pillow-bearing meta-volcanics and arkosic quartzites, some of which are misidentified meta-acidic volcanics; these rocks probably belong to a mature island arc. The central and lower nappes contain semi-pelites, pelitic schists, para-gneissic schists and granulites, the higher grade parts of which have undergone extensive partial melting. It is likely that many of these metasediments were deposited in an accretionary wedge adjacent to an arc. The thrust slab is intruded by a post-tectonic rapakivi granite suite, that includes monzonite, monzogranite, and granite (Allaart, 1976), that have U-Pb and Rb-Sr isochron ages in the range 1.755-1.74 Ga (Gulson and Krogh, 1975). High-grade metamorphism has an isotopic age of 1.8 Ga. Structural evidence shows that these rapakivi granites were intruded into extensional faults (Bridgwater et al., 1974) or extensional shear zones (Hutton et al., 1990). The emplacement of such granites within 60 Ma of the peak of regional metamorphism and associated thrusting is consistent with the time lag, caused by slow thermal
Proterozoic collisional and accretionaly orogens
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relaxation heating (Sonder et al., 1987; Dewey, 1988), between thrusting/crustal thickening and intrusion of crust-mantle melts in extensional zones in a collapsing crust (Windley, 1991a).
The Penokean The 1.9-1.8 Ga Penokean orogen lies along the southern boundary of the Archaean Superior Province in the Lake Superior region (Fig. 8). The main suture is now occupied by the Niagara Fault which is a vertical mylonitic shear zone, and bordered by two minor para-autochthonous terranes (Ueng and Larue, 1988). To the north of the suture is a shelf-foredeep on the Archaean continental margin, and to the south there are two magmatic arc terranes, the Pembine-Wausau and the Marshfield, themselves joined by the Eau Pleine suture (Sims et al., 1989). On the northern continental margin, rift rocks comprise terrigenous sediments, basalts and gabbroic sills (Ueng et al., 1988). The passive margin sediments (Marquette Range Supergroup), deposited between 1.95 Ga and 1.85 Ga, comprise I
A I
I
Florence-Niagara
Suture Zone
//
: :. . :. .:. . . .:. . .:..\: .
Superior Continental Margin
:
n b ine-Wausau rarre
Marshfield Terrane
02.1-13 5 G a Marquette Range Spg, etc. 1.87-1.84; 1.76 Ga Granitoid Arc Rocks
01.88-1 84 G a Volcanic Arc Rocks
a
1.86 G a Volcanic Arc Rocks
1 89-1.84 Ga Granitoid Arc Rocks 3 0-2.8G a Gneisses: 1.89 Ga Tonalite
0
100 kms
u
I
I
90"
88"
Fig. 8. Geological map of the Penokean orogen in the Lake Superior region of North America. Four terranes and two sutures are indicated south of the Superior Continental Margin. Modified after Hoffman (1989).
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lower fanconglomerates, and shelf-type quartzites and algal dolomites overlain by (the Animikie Group) littoral quartzites, iron formation, mafic tholeiites and greywackes. Nd isotopes suggest that northern lower greywackes were derived from an Archaean source, whereas later greywackes to the south have an Early Proterozoic source; following Hoffman’s (1987b) idea that the Animikie infilled a foredeep, Barovich et al. (1989) suggested that the southern greywackes were deposited in a foredeep basin that migrated northwards in front of the cratondirected thrusts and thus were incorporated into the foreland fold-and-thrust belt. Within the 1.85 Ga Niagara suture zone there is serpentinite, a sheeted dyke complex, plagiorhyolite and tholeiitic basalt. Bordering this zone there are two minor fault-bound terranes (Ueng and Larue, 1988); (1) the Crystal Falls Terrane contains starved ocean-basin sediments (chert, carbonaceous slate, sideritic micrite, thin-bedded turbidites, and iron formation) and tholeiitic basalts; there is no evidence of crystalline basement, and (2) the Florence-Niagara Terrane consists of several thrust slices of highly strained dolomite, quartzite, pelite, iron formation, ferruginous slate and tholeiitic volcanics, that belong either to an accretionary wedge (Larue and Ueng, 1985) or to a portion of the passive margin, deformed during accretion tectonics (Ueng and Larue, 1988). The Pembine-Wausau terrane contains an evolved island arc (with no evidence of a continental basement) in which 1.88-1.86 Ga tholeiites are overlain by a 1.85-1.84 Ga calc-alkaline suite of andesite-dacite-rhyolite (Greenberg and Brown, 1983), together with a bimodal suite of high-alumina basalt and dacite-rhyolite, probably formed in an intra-arc rift (Hoffman, 1988). The arc volcanics were intruded by 1.87-1.76 Ga calc-alkaline granitic plutons. In Wisconsin, this terrane is widely interpreted as an island arc that developed above a southward-dipping subduction zone. The Marshfield terrane contains Archaean gneisses overlain by 1.86 Ga mafic-felsic volcanics and intruded by 1.89-1.82 Ga calc-alkaline tonalites and trondhjemites that are chemically similar to modern high-K arc andesites and dacites (Anderson and Cullers, 1987). The steep Eau Pleine suture zone has an age of 1.86-1.835 Ga; Sims et al. (1989) suggested that it developed from a north-dipping subduction zone that gave rise to the calc-alkaline rocks of the Pembine-Wausau terrane. In the southern Penokean orogen there are 1.76 Ga anorogenic metaluminous and peraluminous rhyolites and granites (Smith, 1978), that were generated by melting of continental crust tectonically thickened by the Penokean orogeny (Anderson et al., 1980). These are high-level equivalents of the rapakivi granites of the Ketilidian and Svecofennian orogens. The Birinzinn orogen of WestAfiica The 2.1 Ga Birimian orogen in West Africa (Fig. 9) extends for about 1600 km across strike through E. Guinea, S . Mali, Ivory Coast, W. Ghana, Burkina Faso and W Niger, and it continues along strike in Mauritania in NW
435
Proterozoic collisional and accretionary orogens
Freetown
0 In
Archean & Phanerozoic
ATLANTIC OCEAN
Granites Birimian Volcanic 8 Sedimentary Rocks
200 krns
Fig. 9. Map of the 2.1 Ga accretionary Birimian orogen in West Africa showing the distribution of volcanic and sedimentary rocks and associated granites. After Abouchami e t al. (1990).
Africa. The orogen consists predominantly of greenschist-grade mafic bimodal volcanic rocks including pillow-bearing basalts, crystal and lithic tuffs, volcanodetrital argillites, turbiditic wackes and chemical sediments. These rocks were intruded by muscovite granites (now deformed) with rare garnets and postorogenic leucogranites. Sm-Nd isotopic data by Abouchami et al. (1990) indicate that the sediments are free of any Archaean or older recycled components, suggesting that they formed in ocean basins far from any continental influence, and they confirm contemporaneity of the Birimian sediments and volcanics. Abouchami et al. (1990) found that the trace element signatures of the volcanic rocks are most comparable to those of flood basalts in modern oceanic plateaus and thus proposed that this is a very extensive accretionary orogen that formed in a short time around 2.1 Ga from juvenile, mantle-derived material. In Ghana the wackes contain fragments of chert, graphitic schist, quartz, some lava, but no granite, chemical sediments include chert, Mn-rich rocks, Fe-Mg carbonates, carbon-rich and sulphide-rich rocks, and appreciable gold occurs in a disseminated sulphide facies related to the Mn and in quartz veins and lenses which are highly deformed. Leube et al. (1990) concluded that these sediments were deposited in intracontinental graben. However, the present writer suggests that they more likely accumulated in an accretionary wedge, the auriferous chemical sediments having originally been deposited on the ocean floor. This idea is consistent with the details of local geology (N. Laffoley, personal communication) and with the common occurrence of slices of oceanic plateau basalts (Abouchami et al., 1990).
B. Windley
436
Pan-Afiican of the Arabian-Nubian Shield The Arabian-Nubian shield has an area of about 2600 km by 1000 km and formed in about 310 Ma (Vail, 1985). It has long been known that the shield is an assemblage of accreted island arcs (Al-Shanti and Mitchell, 1976; Cooper et al., 1979; Engel et al., 1980; Camp, 1984). The earliest accretion processes began at 950-900 Ma and the final post-orogenic intrusions, uplift and cooling occurred at 570-500 Ma. Ophiolites have an age range of 870-700 Ma, arc rocks of 840-650 Ma, and post-tectonic crustal melt granites of 610-510 Ma. Apart from isolated continental lead isotope characteristics and zircon ages of 1.6-2.0 Ga (Stacey and Hedge, 1984), the main rocks associations are shelf sediments, ophiolites, volcanic arc rocks, and intrusive granitoids (Kroner, 1985). Figure 10 shows the position in the west of the continental margin of the northern extension of the Mozambique belt, the distribution of arc terranes, and of suture zones with ophiolites.
38"
Fig. 10. Sketch map of the Pan-African geology of the Arabian-Nubian shield after closure of the Red Sea. The accreted terranes, arcs and suture zones are situated east of the continental margin of the Mozambique Belt. Inset shows general occurrence of the two parts of the shield after opening of the Red Sea. Modified after Kroner et al. (1987).
Proterozoic collisional and accretionary orogens
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In Saudi Arabia the new arc material consists of the following: (1)A lower succession of predominant metabasalt, chert, and turbiditic greywackes, that formed on the ocean floor and in immature oceanic island arcs. After deformation (accretion) these rocks were intruded by plutons of diorite and trondhjemite at 910 Ma. (2) Unconformable on the above rocks are volcanic and sedimentary rocks that formed in mature intra-oceanic or marginal island arcs;basal clastics are overlain by andesite-dacite flows and pyroclastic rocks, and upper clastics. Welded ashflows and waterlain tuffs suggest deposition on a partly emergent ridge of an island arc. These rocks were intruded by granitic batholiths dated at 816 Ma and 743 Ma. (3) An upper succession of shallow water clastic sediments and 650-700 Ma rhyolites and ignimbrites, that are the products of syn- and post-accretion crustal melting. Disrupted ophiolite complexes are widespread and their linear extent for up to 900 km outlines the boundaries of the arc terranes (Pallister et al., 1988). The Nabitah terrane consists of medium- to high-grade, strongly deformed supracrustal rocks, and deformed and/or migmatized orthogneisses and granitoids; Kroner (1985) suggested that this terrane underwent thrust-nappe stacking, consequent crustal thickening, and subsequent uplift and erosion, when the entire Afif terrane was thrust westwards over the Hijaz-Air terrane.
DISCUSSION
The above examples provide a means of highlighting the key features of Proterozoic orogens, which can be compared with modern analogues. During the Cenozoic, two principal types of orogens have been and are still being formed. The collisional type, exemplified by the Himalayas, involved little or no accretion of juvenile, oceanic-derived material; the mountain range and orogen grew largely by the deformation of older continental crust. The accretionary type, illustrated actively by Indonesia and Japan, grows largely by the amalgamation of juvenile, oceanic-derived, arc-domina ted material and associated accretionary wedges. Similar types of orogens formed throughout the Proterozoic, as indicated above. Between these two ends of the spectrum, there are of course intermediate types. In the western Himalayas, the Kohistan island arc was trapped between India and Asia at about 50 Ma; Proterozoic comparable examples could be the 2.085 Ga Karasjok greenstone belt/arc in the Belomorian terrane of the Kola-Karelian orogen, the Central Metasedimentary Belt/arc/terrane within the Grenville, and the La Ronge-Lynne Lake and Flin Flon-Snow Lake island arcs in the Trans-Hudson orogen in Canada (Lewry and Stauffer, 1991). During the amalgamation of an accretionary orogen, suspect terranes of buoyant extinct arcs, ridges, oceanic plateaus and microcontinental blocks may be accreted, as for example, in the Cordillera of western North America (Coney et al., 1980); the Archaean Marshfield terrane was thus accreted in the Penokean orogen, there may be extinct ridges and oceanic plateaus in the Nubian-Arabian Shield (Kroner, 1985), and many oceanic plateaus may have contributed to the accretion
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of the Birimian orogen (Abouchami et al., 1990); in other words these orogens provide examples of major crustal growth. It would be misleading to equate these with collisional orogens like the Alps, Himalayas, the Kola-Karelian, Grenville, Wopmay-Thelon, and Capricorn-Albany-Fraser-Pinjarraorogens, which contributed comparatively little to crustal growth. Murphy and Nance (1991) proposed a model to account for these two contrasting types of Late Proterozoic orogens; continent-continent collision gave rise to many “interior” orogens such as the Rokelides and Mozambique belts during amalgamation of the Pan-African/ Brazilian0 supercontinent, whereas “peripheral” orogens such as the ArabianNubian and the Avalonian-Cadomian formed by accretion around the margins of that supercontinent. Significantly, the Cadomian belts, like those in NW France, contain no evidence of continent-continent collision (D’Lemos e t al., 1990). The width of orogens depends on many factors. But it is important to recognise that several Proterozoic accretionary orogens are very wide and extensive. The Svecofennian is 1200 km wide, the Nubian-Arabian shield is 1000 km wide, the 1.7-1.9 Ga area of juvenile crust of the Yavapai-Mazatzal-Central Plains-Penokean orogens is about 600 km by 3000 km (Hoffman, 1989), and the 1600 km wide Birimian of southern West Africa probably continues along strike for 3000 km to NW Africa and westwards across strike into Guyana. As Abouchami et al. (1990) emphasized, the size of these orogens has important implications for calculations of the growth rate of the continental crust. Pelites and greywackes, the typical sediments of accretionary wedges, are inevitably common in A 0 and rare in CO. This difference has a major effect on the relative abundance of crustal melt granitic material in these orogens. The experimental work of Vielzeuf and Holloway (1988) showed that under fluid-absent conditions in the pelitic system a large proportion (ca. 40%) of S-type granitic liquid is produced within the narrow temperature range of 850875°C. Patiiio Douce and Johnston (1991) showed experimentally that greywackes provide an even better fertile protolith for peraluminous granites. This provides an explanation for the abundance of crustal melt granites in the meta-sediments of the Svecofennian and Birimian. But these data are even more important for explaining the fundamental fact that the pre-orogenic (so-called anorogenic) rhyolites and peraluminous granites of central USA (Anderson, 1983) and the post-orogenic rhyolites and rapakivi granites of the Penokean (Smith, 1978), Ketilidian (Windley, 1991a) and Svecofennian orogens (Haapala and Ramo, 1990) occur almost exclusively in the three belts of 1.9-1.8 Ga, 1.8-1.7 Ga and 1.7-1.6 Ga juvenile crust of the North Atlantic region (Fig. 5). Following accretionsubduction, high temperature metamorphism and modest thrust-controlled crustal thickening provided the final recipe for the generation of this long-problematic Proterozoic anorogenic granitic magmatism. This contrast between orogens with abundant crustal melt granites and fertile sedimentary protoliths, and orogens with few granites and more reworked old basement material was suggested by Vielzeuf e t al. (1990). These differences are highly relevant for Proterozoic crustal evolution, as documented in this paper.
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The forelands of both CO and A 0 are characterized by passive-margin, rift-shelf-rise sediments, foredeep basins, and associated foreland fold-and-thrust belts. Likewise they can both contain Andean-type granitic-volcanic batholiths; in accretionary orogens whenever an island arc is accreted to a continental margin, to another arc or amalgam of arcs, further subduction beneath it may create an Andean-type batholith, examples being the Julianehaab batholith in the Ketilidian and the Central Finland batholith in the Svecofennian. Sutures are not easy to define in modern orogens, and therefore not surprisingly in the Proterozoic. In the Himalayas the Indus Suture in Pakistan, that contains no ophiolites, separates high-grade amphibolites on one side (derived from Cretaceous arc/oceanic volcanic rocks) from conformable high-grade gneisses (derived from Precambrian granites and sediments) on the other (Coward et al., 1982). Further east the suture is commonly unexposed, being covered by molasse clastics, because a suture inevitably forms a topographic low between a highstanding thrusted shelf and a high-standing magmatic arc. Ophiolites are likewise uncommon on or bordering modern sutures. In the 3000 km-length of the Indus Suture between Afghanistan and Burma, ophiolites occupy only about 600 km, and there are only three ophiolites exposed, two at high altitudes, that have been thrust onto the adjacent shelf along that length of orogen. Moreover, accretionary wedges in accretionary orogens may contain lenses of serpentinites, meta-gabbros and meta-volcanics, that were thrust directly into the wedge, but these typically isolated lenses do not constitute ophiolites and do not define sutures; e.g. in the Late Proterozoic accretionary prism of the Damara orogen in Namibia (Kukla and Stanistreet (1991). Therefore, the days have passed when it could be argued that the absence or paucity of Proterozoic ophiolites suggests that an orogen is ensialic, or that this is particularly unusual (e.g. Etheridge et al., 1987), especially when an orogen is deeply eroded. The LuleA-Kuopio suture in the Svecofennian is complex but excellent, and the 1.96 Ga ophiolite with its sheeted dykes was thrusted onto the adjacent shelf (Kontinen, 1987). In Canada, the 1.9 Ga Purtuniq ophiolite with its sheeted dykes was obducted onto the foreland of the Cape Smith belt (St-Onge et al., 1989). In Arizona, the 1.73 Ga Payson ophiolite, that contains sheeted dykes, erupted upon a 1.76-1.75 Ga magmatic arc (Dann, 1991). So many data are available from Proterozoic orogens that are similar to equivalents in modern orogens that it can no longer be in doubt that modern-style plate tectonic processes were in operation in the middle part of Earth history. However, there are some minor but important differences, like the presence of magmatic rocks in Early Proterozoic foredeeps (Hoffman, 1987b).
ACKNOWLEDGEMENTS
I am grateful for useful discussions with Celal Sengor, Nick Laffoley, and John Brney, helpful reviews by Robert Shackleton and Paul Hoffman and editorial comments by Kent Condie.
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1-10. Windley, B.F., 1991b. Precambrian Europe. In: DJ. Blundell, R. Freeman and St. Mueller (Editors), Tectonic Evolution of a Continent: the European Geotraverse. Cambridge University Press, New York, N.Y. (in press). Windley, B.E, Henriksen, N., Higgins, A K . , Bondeson, E. and Jensen, S.B., 1966. Some border relations between supracrustal and infracrustal rocks in Southwest Greenland. Gr@nlandseGeol. Unders., Raw., 9,43 PP. Windley, B E , Allen, M.B., Zhang, C., Zhao, Z H . and Wang, G-R., 1990. Paleozoic accretion and Cenozoic redeformation of the Chinese Tien Shan range, Central Asia. Geology, 18: 128-131. Windrim, D.P. and McCulloch, M.T., 1986. Nd and S r isotopic systematics of central Australian granulites: chronology of crustal development and constraints on the evolution of lower continental crust. Contrib. Mineral. Petrol., 94: 289-303. Wyborn, L.A.I., 1988. Petrology, geochemistry and origin of a major Australian 1880-1840 Ma felsic volcano-plutonic suite: a model for intracontinental felsic magma generation. Precambrian Res., 40/ 41: 37-60.
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Chapter 12 PROTEROZOIC TERRANES AND CONTINENTAL ACCRETION IN SOUTHWESTERN NORTH AMERICA K.C. CONDIE
INTRODUCTION
It is generally agreed that significant portions of the Appalachian and Cordilleran orogenic belts in North America are composed of accreted terranes. Despite limited exposure, geological and geochemical studies of Proterozoic supracrustal successions in southwestern North America suggest that much of this region also comprises accreted terranes that were added to North America during the Early and Middle Proterozoic (Condie, 1982, 1986; Reed et al., 1987; Karlstrom and Bowring, 1988). Similar terranes were accreted to Archean crust in Scandinavia at about the same time (Park, 1991). Several terms have been ' used for segments of continental crust associated with continental accretion, and because they may be used in different ways, it is important to define terms used in this study. Terrane will follow the conventional usage of Jones et al. (1983) and Coney (1989) and refer to a fault-bounded segment of continental crust with a distinctive assemblage of rocks and a tectonic history different from surrounding terranes. A supertewane (or composite terrane) results from amalgamation of two or more terranes prior to final accretion with a continent, and an overlap assemblage is an assemblage of supracrustal rocks that overlaps older terrane boundaries. Finally, a crustal province is a large segment of continental crust (?lo7 km2), composed of several to many terranes that accreted during a limited time interval. Crustal provinces commonly contain anatectic granites younger than the province accretion age. Although terranes in a given province may have different crustal formation ages (i.e., ages of mantle extraction), they all accrete during the same time interval to form the province. For instance, the Cordilleran Province in western North America is composed of numerous terranes of different ages that were accreted to North America during the last 350 Ma (mostly between Late Jurassic and Eocene). With these definitions, four crustal provinces, eight terranes and five overlap assemblages are recognized in the Proterozoic of southwestern North America (Fig. 1; Bible 1). Each of the crustal provinces has distinct Pb isotopic characteristics (Anderson et al., 1991).
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THE MOJAVE PROVINCE
The westernmost Proterozoic province in southwestern North America is the Mojave Province in Nevada and southern California (Fig. 1). Most of the data from this province come from southeastern California and adjacent areas
Fig. 1. Distribution of Early and Middle Proterozoic crustal provinces and terranes in the southwestern United States. Bold dashed lines are proposed crustal province boundaries and patterned areas are terranes and overlap assemblages. Crustal province ages are given along their southern boundaries. Numbers (1-26) refer to locations mentioned in the text. A = Ash Creek; B = Big Bug; and G = Green Gulch blocks (Karlstrom and Bowring, 1988). The Dos Cabezas and Pinal domains appear to be part of the same terrane.
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TABLE 1 Summaly of Proterozoic provinces, terranes, and overlap assemblages in southwestern North America Province
Terrane (T) or overlap assemblage (OA)
1. Mojave
2. Yavapai
Tectonic setting
2300-1700 Green Mountain (T) Dubois (T) Ash Greek (T> Hualapai (T) Idaho Springs-Black Canyon (OA) Wet Mountains (OA) Cochetopa-Salida (OA)
3. Mazatzal
Pecos (T) Dos Cabezas-Pinal (T) Alder (T)
4. Grenville
Carrizo Mountains (T) Franklin Mountains (OA)
a
Age (Ma) 1790-1780 1780-1750 1750-1740 (1780-1750) 1740-1700 1740-1720 1720-1670 1740-1730 1720 1700-1670 -1700 (1760-1740)
1350-1270 1150-1135
CratonKMA
IA IA IA a
CMA CMBAB CMBAB CMA IA CMA CMBAB a
CMBAB CREB
Numbers in parenthesis represent minor components.
CMA = continental-margin arc; IA = island arc; CMBAB = continental-margin backarc basin; CR = continental rift; TB = transpressional basin. Based on published U-Pb zircon dates from igneous rocks; major references: Silver (1968, 1978); Silver and Barker (1968); Anderson and Silver (1981); Bowring and Condie (1982); Bowring et al. (1983); Bickford and Boardman (1984); Bowring et al. (1984); DeWitt et al. (1984); Bryant and Wooden (1986); Silver et al. (1986); Karlstrom et al. (1987); Reed et al. (1987); Chamberlain et al. (1988); Copeland and Bowring (1988); Wooden et al. (1988a, b); Bickford et al. (1989); Robertson and Condie (1989); Chamberlain and Bowring (1990); Walker (1990); Walker et al., (1990); Karlstrom and Bowring (1991).
in Nevada and the proposed extension into northeastern Nevada and Utah is uncertain. Most widespread in the province are layered gneisses, migmatites, and foliated plutons. Little is known of supracrustal rocks in the Mojave Province and terranes as yet have not been described, but available data suggest that supracrustal rocks are diverse. Along the eastern part of the province adjacent to the Hualapai terrane, geochemical results from highly metamorphosed gneisses and schists suggest graywacke and volcanic protoliths with arc-like affinities (Wooden and Miller, 1991). Farther west in the Halloran Hills (18, Fig. 1) and Death Valley areas, quartzites with locally preserved cross bedding and metapelites become more important, suggestive of a more cratonic assemblage (Warnke, 1969; Wright, 1974; Labotka et al., 1980; Howard et al., 1982; Wooden et al., 1986). Metamorphic grade ranges from amphibolite throughout most of the western part of the province to granulite in the east. Foliations are generally north trending with steep dips, usually to the west. Metamorphic studies in southern Nevada and southeastern California indicate widespread low? granulite-facies metamorphism
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generally in the range of 500-700°C and pressures of 2-4 kbar (Thomas et al., 1988; chamberlain et al., 1988; Young et al., 1989; Chamberlain and Bowring, 1990). U-Pb zircon ages constrain the timing of this metamorphism and coeval deformation to about 1700 Ma (Wooden et al., 1988; Wooden and Miller, 1990; Anderson et al., 1991). U-Pb zircon ages from the Mojave Province indicate a long, complex history. In southeastern California and adjacent Arizona, deformed plutons with ages from 1760 to 1630 Ma intrude supracrustal rocks (DeWitt et al., 1984; Wooden et al., 1988; Chamberlain and Bowring, 1990; Wooden and Miller, 1990). Pb-Pb ages from paragneisses range up to 2300 Ma, similar to Nd model ages (TDM) of some granitic rocks from the region (Bennett and DePaolo, 1987). U-Pb zircon chronology from southeastern California and southern Nevada suggests: deposition of sediments at 2000-1800 Ma; intrusion of pretectonic plutons at 1760-1730 Ma; strong compressional deformation, l o w 2 granulite metamorphism and emplacement of syntectonic plutons at 1710-1700 Ma causing the Ivanpah orogeny; post-tectonic intrusion of plutons at 1690-1630 Ma; development of a major mylonite zone at 1650-1600 Ma; emplacement of anorogenic granites at 1430-1400 Ma; and intrusion of diabase dike swarms at about 1200 Ma (Wooden et al., 1988; Wooden and Miller, 1990; Anderson et al., 1991). A great deal of attention has focused on the nature of the eastern boundary of the Mojave Province in western Arizona. Nd and Pb isotopic data imply the presence of a crustal boundary in this region, and also indicate an enriched, probable Archean component in igneous sources, whereas data from igneous rocks in the Yavapai Province to the east reflect relatively depleted mantle sources (Bennett and DePaolo, 1987; Bennett et al., 1988; Wooden et al., 1988). Karlstrom and Bowring (1991) have described the Gneiss Canyon shear zone, a several km-wide, NE-trending shear zone in the lower Grand Canyon (13, Fig. l), which continues southwest across the Grand Wash Cliffs and northern Peacock Mountains (Fig. 2). The shear zone has stretching lineations plunging steeply to the west and kinematic indicators suggest NW side up, dextral movement. U/Pb zircon ages from pre- and post-tectonic granitoids in the Gneiss Canyon shear zone constrain major shearing to 1.73-1.71 Ga and peak metamorphism to 1.68-1.62 Ma (Albin et al., 1991). U-Pb ages from sphene and apatite indicate different cooling and uplift histories on both sides of the shear zone (Chamberlain and Bowring, 1990). Also, metamorphic grade rapidly decreases from granulite to amphibolite grade and 1.4 Ga A-type granites change from metaluminous to peraluminous from west to east across the boundary zone (Anderson et al., 1991). Although the Gneiss Canyon shear zone may be the eastern boundaq of the Mojave Province at shallow crustal levels, Pb isotope data from granitoids suggest that the Mojave Province could underlie much of the Hualapai terrane to the east (Wooden and Aleinikoff, 1991).
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451
Granitic Rocks ( 2 1 7 0 0 M a Gneisstc Complex Ouartrite- Pelite Successior Volcanic Sediment Succersloi Pelite Succession
350
114-
Fig. 2. Generalized geologic map of Proterozoic rocks in the vicinity of Kingman, Arizona. Dashed lines labelled TDM are Nd model age boundaries suggested by Bennett and DePaolo (1987). Location of the Gneiss Canyon shear zone from Karlstrom and Bowring (1991).
THE YAVAPAI PROVINCE
The Yavapai Province, which extends from southwestern Arizona into the midcontinent region, includes the Colorado Province and northern part of the Central Plains Province as previously described by Bickford et al. (1986). In the Southwest, four terranes and three overlap assemblages are exposed in this province (Fig. 1; a b l e 1). The terranes include chiefly submarine volcanic and volcaniclastic rocks with geochemical and isotopic affinities to rocks from modern oceanic arcs (Condie, 1986). They show no evidence for forming on or near significantly older continental crust either in terms of Nd isotopes or trace element characteristics (DePaolo, 1981; Nelson and DePaolo, 1984; Condie, 1986; Wortman et a]., 1990).
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K.C. Condie
Green Mountain terrane The Green Mountain terrane includes a bimodal volcanic suite in which pillow basalts, hyaloclastic breccias and agglomerates, and volcaniclastic sediments dominate (Swift, 1982; Shadel, 1982; Condie and Shadel, 1984; Snyder et al., 1988). Although metamorphosed to greenschist or amphibolite facies, primary textures and structures are commonly preserved and indicate a subaqueous, probable submarine origin for most of the rocks. Gabbros exposed in the northern Park Range near the Wyoming-Colorado state line may represent deeper crustal levels of the mafic volcanic rocks (Pallister and Aleinikoff, 1987). Metapelites and quartzites occur locally in the northern Park Range and appear to have been deposited in a near-shore marine environment (White and Foster, 1987; Snyder et al., 1988). Both the lithologic assemblage and the geochemical characteristics suggest the Green Mountain terrane is a remnant of an oceanic arc (Condie and Shadel, 1984; Condie, 1986). U-Pb zircon ages from volcanic and plutonic rocks in this terrane indicate an age of 1790-1780 Ma (Premo and Van Schmus, 1989; Premo, 1991). The Cheyenne shear zone (Cheyenne belt), which is the northern boundary of the Green Mountain terrane, is the only well documented exposure of a Proterozoic suture in the southwestern United States (Karlstrom and Houston, 1984). This shear zone, which ranges from about 0.5 to 7 km wide, is exposed in the Sierra Madre and Medicine Bow Mountains in southeastern Wyoming ( I , Fig. 1). It consists of strongly deformed, lithologically distinct fragments bounded by mylonite zones ranging in width from 50 to 200 m (Duebendorfer and Houston, 1986). Kinematic and metamorphic studies show that the Cheyenne shear zone is a major low-angle, northward-directed thrust and is a suture between the Green Mountain oceanic arc and the Wyoming Craton (Duebendorfer and Houston, 1986, 1987; Duebendorfer, 1988). U-Pb ages from plutons emplaced before and after deformation constrain major deformation in the shear zone to between 1780 and 1750 Ma (Premo and Van Schmus, 1989; Premo, 1991). Dubois terrane The Dubois terrane includes the Dubois greenstone south of Gunnison, Colorado (2, Fig. l), the Irving greenstone in the Needle Mountains in southwestern Colorado ( 1 9 , the Moppin Series in the Tusas Mountains in northern New Mexico (6), and the Gold Hill and related successions near n o s , New Mexico (16). These greenstone successions are composed chiefly of bimodal volcanic rocks with mafic volcanics generally exceeding felsic volcanics. Mafic volcanics include pillow basalts, hyaloclastic breccias, and associated volcaniclastic sediments (Barker, 1969; Kent, 1980; Hedlund and Olson, 1981; McCrink, 1982; Condie and McCrink, 1982; Reed, 1984; Shonk, 1984; Gabelman, 1988; Knoper and Condie, 1988; Knoper, 1991). Mafic volcaniclastic sediments dominate in the northern New Mexico successions (Kent, 1980; Gabelman, 1988). Felsic volcanic rocks are gener-
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453
ally hyaloclastic tuffs and associated epiclastic sediments. Minor rock types include chert and ironstone, probably deposited on the seafloor by local hydrothermal vents. As with the Green Mountain terrane, geochemical characteristics of mafic volcanic rocks from the Dubois terrane are similar to those of basalts from modern oceanic arcs (Condie and McCrink, 1982; McCrink, 1982; Condie, 1986; Knoper and Condie, 1988; Knoper, 1992). U-Pb zircon ages from the Dubois terrane are somewhat younger than the Green Mountain terrane. Most fall in the range of 1780-1750 Ma (Silver and Barker, 1968; Bickford and Boardman, 1984; Bowring et al., 1984; Reed et al., 1987; Bickford et al., 1989). Although now separated from each other by the Idaho Springs-Black Canyon overlap assemblage, the Dubois and Green Mountain terranes may have been part of the same superterrane at the time of collision with the Wyoming Craton. Ash Creek terrane
The Ash Creek terrane in west-central Arizona (Fig. 1) includes the Yavapai Supergroup and related rocks. Published U-Pb zircon ages from Ash Creek igneous rocks indicate an age for the terrane between 1780 and about 1740 Ma with most rocks falling between 1760 and 1740 Ma (Karlstrom et al., 1987; Bowring et al., 1991; Karlstrom and Bowring, 1992). Karlstrom and Bowring (1988) have identified three “blocks” in the Ash Creek terrane separated by the Shylock and Chaparral faults northwest of Phoenix, Arizona. These are the Ash Creek, Big Bug, and Green Gulch blocks (A, B, and G in Fig. 1). The bounding fault zones are intensely foliated with steep-plunging lineations and appear to record simple shear displacement. Before being brought together by thrust and strike-slip faulting at about 1700 Ma during the Yavapai orogeny, each of these blocks experienced different deformational histories. For instance, the Ash Creek block appears to have collided and been uplifted in a short period of time (< 20 Ma) at 1700 Ma, whereas the Big Bug block records a long cooling and uplift history of 100-200 Ma (Bowring and Karlstrom, 1990; Bowring et al., 1991). Geochemical characteristics and proportions of volcanic rocks indicate that each block has affinities to oceanic arcs (Vance, 1989). It is not yet clear, however, whether the three blocks are remnants of different oceanic arcs or if they represent different levels of exposure or/and lateral segments of the same arc. It is also not known if the Ash Creek terrane extends beneath the Four Corners area and connects with the Dubois terrane in Colorado. Two major deformations are recorded in the Ash Creek terrane. The earliest, as reflected by NW-trending folds, occurred between 1740 and 1735 Ma as constrained by zircon dates from pre- and post-deformational plutons (Karlstrom and Bowring, 1991). The second deformation, evidence for which is found in all terranes in central and western Arizona is the Yavapai orogeny that occurred at about 1700 Ma. In the Ash Creek terrane, it is characterized by a prominent NE-trending fabric.
454
K.C. Condie
The Yavapai Supergroup is composed chiefly of submarine mafic to felsic volcanic rocks and associated volcaniclastic sediments (Anderson and Creasey, 1958; Anderson et al., 1971; O’Hara, 1980; Vance and Condie, 1986; Vance, 1989). Iron formation and small massive sulfide deposits are locally important at some locations (Anderson and Nash, 1972; Slatt et al., 1978; Anderson and Guilbert, 1979; DeWitt, 1979; Lindberg, 1986; Vance and Condie, 1987). The Ash Creek terrane is the only Proterozoic terrane in southwestern North America that contains relatively large volumes of andesite (Vance and Condie, 1986; Vance, 1989). Pace element distributions in Yavapai mafic and andesitic volcanics favor an oceanic arc setting and a relatively depleted mantle source (Vance, 1989).
Hualapai terrane The Hualapai terrane in western Arizona is tectonically sandwiched between the Ash Creek terrane and the Mojave Province (Fig. 1). It is characterized by large volumes of deformed granitoids ranging in age chiefly from 1740 to 1700 Ma, but with some as old 2.2-1.8 Ga (Karlstrom and Bowring, 1991). U-Pb zircon ages from felsic volcanic rocks in the Bagdad area generally fall in the range of 1740 to 1715 Ma with some as low as 1700 Ma (Silver, 1966; Bryant and Wooden, 1986; Chamberlain and Bowring, 1990). Similar ages are recorded by granitoids in the bottom of the Grand Canyon (Karlstrom and Bowring, 1991). The western boundary of the Hualapai terrane is the Gneiss Canyon shear zone (13, Fig. 1) and the eastern boundary is probably the Mesa Butte shear zone in Chino Valley. In the Bagdad area, supracrustal rocks include large volumes of mafic and felsic tuffs, pillow basalts, and the Dick Rhyolite, a thick homogeneous sill or massive flow (Anderson et al., 1955; Conway et al., 1986). Similar rhyolite tuffs and flows occur in Peach Springs Canyon, a tributary to the Colorado River in the lower Grand Canyon. In the Cottonwood Cliffs area northeast of Kingman, Arizona (Fig. 2), two fault-bounded successions of supracrustal rocks are preserved (Beard, 1986). One includes pillow basalts, felsic ash flow tuffs and volcaniclastic sediments, and the other pelites, sandstones, and conglomerates. Zircons from the ash flow tuffs yield U/Pb ages of about 1740 Ma and the assemblages are intruded by the Vallentine granite with a U-Pb zircon age of about 1713 Ma (Bowring et al., 1991). Also, an apparently thick succession of graywacke turbidites with minor pillowed basalt flows occurs in the northern Hualapai Mountains (Conway et al., 1986). The Vishnu Complex in the bottom of the Grand Canyon (17, Fig. 1) includes metasedimentary and metavolcanic rocks and is part of the Hualapai terrane (Clark, 1979). Deformation in the NW Hualapai terrane is similar to that of the adjacent Mojave Province, where NW-striking foliation that developed at about 1.74 Ga is transposed within later NE-trending shear zones (Karlstrom and Bowring, 1991). Timing of the later deformation is bracketed by U/Pb zircon ages to between 1.72 and 1.69 Ga. Similar styles and ages of deformation in the Hualapai terrane and the Mojave Province suggest they collided at approximately 1.7 Ga, during the Yavapai-Ivanpah orogeny.
Proterozoic terranes and continental accretion in Southwestern North America
455
It is likely that the Hualapai terrane is a remnant of a continental-margin arc, which terminated activity upon collision with the Mojave Province.
THE MAZATZAL PROVINCE
The Mazatzal Province includes three terranes in Arizona and New Mexico and the Manzano overlap assemblage in New Mexico (Fig. 1; Table 1).
Pecos terrane The Pecos terrane is exposed in the southern Sangre de Cristo Mountains near Santa Fe, New Mexico (7, Fig. l), in the Manzanita Mountains south of Albuquerque (Woodward et al., 1979; Condie 1980) (9, Fig. l), and in the Pedernal Hills east of Albuquerque (McKee, 1988) (8, Fig. 1). It includes a bimodal maficfelsic volcanic suite in which mafic end members dominate (Robertson and Moench, 1979; Robertson and Condie, 1989). Volcaniclastic sediments and ash flow tuffs are also important in the Pedernal Hills (Armstrong and Holcombe, 1982; McKee and Condie, 1985; McKee, 1988). Well-preserved textures and structures indicate largely or entirely subaqueous eruption. Ultramafic rocks and associated high-Mg basalts that occur near Pecos, New Mexico are cut by mafic dikes and also by trondhjemitic dikes, suggestive of an ophiolite origin for some of the Pecos terrane (Wyman, 1980; Robertson and Condie, 1989). However, layered gabbros, sheared harzburgites, and sheeted diabase dikes have not been found. U-Pb zircon dates from felsic volcanics and associated intrusive rocks in the Pecos terrane indicate igneous ages of about 1720 Ma (Bowring and Condie, 1982; Robertson and Condie, 1989). Geochemical characteristics of the mafic volcanics suggest a relatively depleted mantle source and high ENd values (+4 to +7) support this conclusion (Nelson and DePaolo, 1984). These data are consistent with an oceanic arc.
Alder terrane The Alder terrane in central Arizona comprises a thick succession of sediments and volcanic rocks intruded by granites, and includes the Mazatzal and Sunflower blocks of Karlstrom and Bowring (1991, 1992). The oldest recognized unit in the Alder terrane is a 1.76 Ga old granite in the Tonto Creek area, which is basement to overlying supracrustal rocks (Karlstrom et al., 1990; Dann, 1991). This basement is intruded by mafic to felsic dikes that are part of the overlying Payson ophiolite (N1.73 Ga) which includes an intrusive gabbro-diorite phase, the Gibson Creek Complex (Dann, 1991). This complex is overlain by the East Verde River Formation ( ~ 1 . 7 2Ga), which includes in ascending order, pillow basalts, a marker unit of dacitic breccia and jasper, and a thick succession of graywacke turbidites, with two felsic ash flow tuffs. Overlying the East Verde River
456
K. C. Condie
Formation is the Tonto Basin Supergroup, which includes from bottom to top, the Alder Group, composed chiefly of submarine volcanic rocks and associated volcaniclastic sediments with smaller amounts of quartzite; the Red Rock Group, chiefly ash flow tuffs and related volcanic rocks; and the Mazatzal Group, which includes massive cross-bedded quartzites and associated pelites (Wilson, 1939; Gastil, 1958; Ludwig, 1974; Conway, 1976; Anderson, 1986; Noll, 1988; Conway and Silver, 1989; Conway et al., 1991). Primary textures and structures are well preserved in most supracrustal rocks in the Alder terrane and indicate dominantly submarine deposition. All volcanic rocks have continental-margin arclike geochemical affinities (Noll, 1988; Condie et al., 1992). The rocks have been deformed by a NW-SE compressional event, the Mazatzal orogeny, perhaps in a foreland thrust belt (Karlstrom et al., 1987; Karlstrom and Bowring, 1991). U-Pb zircon ages from felsic tuffs and associated granites suggest deposition of the Tonto Basin Supergroup occurred over a short period of time at about 1700 f 10 Ma (Ludwig, 1974; Silver et al., 1986; Karlstrom et al., 1987; Conway and Silver, 1989). Cross-bedding in interbedded quartzites and turbidites in the Alder Group clearly show quartzite sources to the north and turbidite sources to the south or southeast (Noll, 1988). Geochemical characteristics of these sediments, furthermore, indicate that the northern source was a craton and the southern one an arc system - probably a para-autochthonous continental-margin arc (Condie et al., 1992). The Alder basin may have been a small submarine back-arc basin where sediments from the two sources interfingered during deposition. With time, craton-derived sediments filled the basin as the Mazatzal Group was deposited in shallow marine and fluvial environments (llevena, 1979). The Payson ophiolite has been described in a tectonic block within the Alder terrane south and west of Payson, Arizona (Dann, 1991). The pseudostratigraphic elements incIude in ascending order, layered gabbro, gabbro-tonalitequartz diorite, a sheeted dike complex with granitoid screens, and submarine volcanics. U-Pb zircon ages from the ophiolite are about 1730 Ma (Dann, 1991) significantly older than the Alder Group. The ophiolitic gabbro intrudes the 1760 Ma granitic basement, and thus the ophiolite may reflect the initial stages of rifting of continental crust that produced a back-arc basin. Whether such a back-arc basin evolved into the Alder basin at 1700 Ma is not clear from the existing data base. The northwestern boundary of the Alder terrane is exposed as the Moore Gulch fault north of Phoenix (10,Fig. 1). This fault is a relatively young fault and is not the original contact of the terrane. The original boundary may have been a low-angle decollement where the Alder terrane is thrust northwest over the Ash Creek terrane, or it may have been an uncomformable contact (Karlstrom et al., 1987). Structural and geochronologic data favor the decollement interpretation (Karlstrom and Bowring, 1988). Just how far south the Ash Creek terrane extended prior to formation of the Alder terrane at 1700 Ma is unknown. However, the 1760 Ma plutonic rocks in the Tonto Basin area indicate that it may underlie at least part of the Alder terrane. A north-trending isostatic gravity anomaly in eastern Arizona and adjacent New Mexico (Simpson et al., 1986) is on
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451
line with the Moore Gulch fault and may represent the northern extent of the Ash Creek terrane in this region (25, Fig. 1). An alternate and less preferred extension of this boundary zone is the Holbrook lineament in northeastern Arizona, which trends northeast towards the Four Corners area (Karlstrom and Bowring, 1988).
Dos Cabezas-Pinal terrane The Dos Cabezas-Pinal terrane in southeastern Arizona (Fig. 1) can be divided into two domains based on lithologic association. The Dos Cabezas domain includes bimodal mafic and felsic volcanics, quartzites, arkoses, conglomerates, and a variety of volcaniclastic sediments (Erickson, 1968; Silver, 1978; Condie et al., 1985; Copeland, 1986; Copeland and Condie, 1986; Bowling, 1987). Trace element distributions in volcanics and sediments and detrital modes of sediments indicate mixed arc and cratonic sources, much like the Alder terrane. Massive cross-bedded quartzites and associated pelites similar to the Mazatzal Group occur in the Klondyke area, the southern Dos Cabezas Mountains, and near Morenci, Arizona. U-Pb zircon dates of 1670 Ma from felsic volcanic rocks in the western Dos Cabezas Mountains east of Tucson (Erickson and Bowring, 1990) indicate an age for the Dos Cabezas-domain only slightly younger than the 1700 Ma Alder terrane. The age of major compressive deformation in the Dos Cabezas-Pinal terrane appears to be about 1650 Ma from a zircon date in the syntectonic Sommer pluton near Willcox, Arizona (Erickson and Bowring, 1990) (near 12, Fig. 1). Reconnaissance field studies by the author northwest of Superior, Arizona have revealed the presence of a shear zone herein referred to as the Hewitt Canyon shear zone (11, Fig. 1). Although the northern side of the Hewitt Canyon shear zone is not preserved due to intrusion by the 1450 Ma Ruin granite, this shear zone could be the boundary between the Dos Cabezas-Pinal and Alder terranes. Remnants of the Alder terrane are preserved as xenoliths and roof pendants in the granite north of Salt River Canyon in the southern Mazatzal Mountains. A U-Pb zircon date of 1660 Ma (Karlstrom and Bowring, 1991) from felsic volcanics of the Redmond Formation in the White Ledges area north of Globe, Arizona indicates the Dos Cabezas-Pinal terrane extends at least as far north as the Salt River in this area. This age also shows that the Hess Canyon Group that overlies the Redmond Formation (Conway and Silver, 1989) is not correlative with the 1.7 Ga Mazatzal Group to the north, but is a younger quartzite-pelite succession. The Pinal domain is composed largely of a homogeneous sequence of quartz wacke turbidites with sedimentological and geochemical characteristics of submarine fan deposits derived from a recycled orogen (Condie and DeMalas, 1985; Copeland and Condie, 1986). The quartz wackes are widespread and are recognizable in south-central Arizona at higher metamorphic grades. Graded bedding and other turbidite textures are well preserved in these rocks at many localities (Ransome, 1904; Peterson, 1962; Cooper and Silver, 1964; Copeland, 1986). Bimodal mafic and felsic volcanic rocks are of local importance in some areas, as
458
K. C. Condie
for instance in the vicinity of Mammouth, Arizona. U-Pb detrital zircon ages from Pinal quartz wackes from the Mule Mountains near Bisbee, Arizona show that the sources of these rocks were not much older than 1700 Ma (S.A. Bowring, personal communication, 1985). A felsic sill in the wacke succession in the Johnny Lyon Hills provides an upper limit on depositional age of about 1690 Ma (Silver, 1963, 1978). Provenance studies of Pinal quartz wackes show they are derived chiefly from mixed felsic volcanic and granite sources. The sources for Pinal sediments seem to have been from the north and northeast (Copeland and Condie, 1986; Copeland 1986), probably from the Dos Cabezas domain. These or more distant sources must have been elevated sufficiently to expose granitoids to account for the large volume of monocrystalline quartz sand and granite rock fragments in Pinal sediments. The tectonic setting of the Pinal domain continues to be problematic. Earlier interpretations as a continental rift or aulacogen (Condie and DeMalas, 1985; Copeland and Condie, 1986) no longer seem feasible. The Pinal sediments could represent a deep-water facies associated with an arc environment in the Dos Cabezas domain. Such a model implies a continental-margin platform, a continental slope where turbidity currents are generated, and a nearshore continental-rise basin where Pinal turbidites accumulate. The eastern boundary of the Pinal domain is not exposed, but if this depositional model is correct, it could be an unconformable contact with underlying arc rocks of the Dos Cabezas domain.
THE GRENVILLE PROVINCE
Most subsurface data including Nd isotopic results from crustal xenoliths in Mexico suggest that the Grenville Province extends across Rxas and into central and southern Mexico (Flawn, 1956; Wasserburg et al., 1962; Patchett and Ruiz, 1987; Ruiz et al., 1988, 1990). In west Texas it is represented by the Carrizo Mountains terrane, which is exposed in several mountain ranges in the vicinity of Van Horn, Texas (Fig. 1). This terrane includes the Carrizo Mountains Group, a deformed and metamorphosed succession of chiefly felsic ash flow tuffs and basalts with U/Pb zircon ages of about 1350 Ma and the younger Allamore and Hazel Formations comprising detrital sediments and carbonates locally intruded with diabase (King and Flawn, 1953; Soegaard et al., 1991). Zircon ages from tuffs in the Allamore Formation suggest deposition at about 1280 Ma (Soegaard et al., 1991). The Carrizo Mountains Group is thrust northward over the Allamore Formation along the Steeruwitz thrust. Thrusting led to development of a foreland basin north of the fold and thrust belt in which >2500 m of syn-orogenic alluvial fan and eolian sediments of the Hazel Formation accumulated. A U/Pb zircon age of 1194 Ma from a granite boulder in the Hazel Formation gives a maximum age for this deformation (Callaham and Soegaard, 1991). Volcanic rocks from the Carrizo Mountains Group have trace element distributions similar to continental-margin arc volcanic rocks (Rudnick, 1983), and both rock association
Proterozoic terranes and continental accretion in southwestern North America
459
and geochemistry suggest that this group was deposited in a continental-margin back-arc basin. U-Pb zircon ages from syntectonic plutonic rocks in the Sierra Del Cuervo area of Chihuahua, Mexico and from ash flow tuffs in the Carrizo Mountains indicate an age for the Carrizo Mountains terrane of 1350 to 1270 Ma (Dennison and Hetherington, 1969; Blount et al., 1988). Similar zircon ages are reported from the Llano uplift in central Texas (Walker, 1988; Walker et al., 1990; Mosher, 1991) consistent with an eastward continuation of the Carrizo Mountains terrane.
OVERLAP ASSEMBLAGES
Idaho Springs-Black Canyon assemblage
The Idaho Springs-Black Canyon overlap assemblage (IBA) is located in westcentral and north-central Colorado (Figs. 1 and 3). It comprises a complexly deformed suite of metasediments with associated syntectonic migmatites and granites (Braddock, 1970; Nesse, 1984; Tweto, 1987; Reed et al., 1987). Included also are variable proportions of felsic gneisses and amphibolites that were probably volcanic rocks. Petrographic and geochemical studies of the metasediments reveal recycled-orogen characteristics and favor deposition in either a continental backarc or foreland basin (Condie and Martell, 1983). Remnants of massive, crossbedded quartzites occur in some successions such as south of Boulder, Colorado (Wells, 1967) and seem to reflect input from continental sources. However, U-Pb ages of detrital zircons from IBA metasediments (1850-1800 Ma) indicate that the Archean Wyoming Craton did not serve as a source for these sediments (Aleinikoff et al., 1985). Although the last (and perhaps only) major deformation, metamorphism and syntectonic plutonism in the IBA occurred at 1700-1670 Ma (Reed et al., 1987), the depositional age of the sediments is not well constrained. In the Black Canyon area (4, Fig. l), IBA sediments appear to grade into sediments and volcanic rocks of the Cochetopa-Salida assemblage (Knoper, 1992) which is dated at 1740-1730 Ma (Bickford and Boardman, 1984), suggesting a similar age for IBA sedimentation. With exception of this gradational contact, none of the IBA boundaries is exposed. From pendants and xenolith populations in plutons, however, it appears that the northern boundary must lie in the vicinity of the Colorado River in the northern Gore Range (19, Fig. 1). A pronounced change in xenolith population in the northern part of the Rawah batholith (M.E. McCallum, personal communication, 1988) also may reflect the northern boundary of the IBA (20, Fig. 1). Although the western extent of the IBA is unknown, outcrops along the Utah-Colorado state line indicate that it extends into the basement of southeastern Utah. There is no evidence that it joins with the Hualapai terrane in Arizona, but it may extend as far west as the Mineral Mountains in west-central Utah (24, Fig. l), where Proterozoic rocks of similar lithology and age are exposed
460
K.C. Condie
(Aleinikoff et al., 1986). Drillcore data (Tweto, 1987) also indicate that the IBA extends into the basement of northeastern Colorado. Considering all results, a continental back-arc basin is favored for the IBA (Condie, 1986; Knoper, 1991). The Cochetopa-Salida assemblage south of the IBA may have been part of the associated arc system. Although there is no evidence for oceanic crust forming in the IBA back-arc basin (Knoper and Condie, 1988; Knoper, 1991), a considerable thickness of volcaniclastic sediments must have accumulated in the basin between 1740 and 1720 Ma, perhaps extending to 1700 Ma. Burial of these sediments and steepened geotherms associated with thinning of the lithosphere facilitated partial melting as reflected by widespread migmatites in the IBA. Wet Mountains assemblage The Wet Mountains assemblage underlies the Wet Mountains and central Sangre de Cristo Mountains in south-central Colorado (Figs. 1and 3). It includes a complexly deformed and partially melted suite of metasediments with syntectonic, anatectic granites (Brock and Singewald, 1968; Noblett, 1987; Noblett et al., 1987; Lanzirotti, 1988; Hallett, 1990; Hallett and Condie, 1990). Amphibolites or other rocks that might represent metavolcanics are minor in importance. Geochemical studies of the metasediments indicate most are quartz wackes and pelites with affinities to sediments deposited in continental-margin back-arc basins (Lanzirotti, 1988). Massive cross-bedded quartzites and associated pelites are locally preserved (Reuss, 1974), supporting the existence of continental sources for some of the sediments. Metamorphic grade ranges from upper amphibolite to lower granulite facies, and few primary textures or structures have been preserved in the rocks. Structural studies indicate three periods of isoclinal folding with the dominant deformations (D1 and D2) reflecting NW-SE compression (Lanzirotti, 1988). U-Pb zircon ages from late syntectonic plutons in the northern Wet Mountains suggest that the first two deformations occurred at about 1700 Ma, the age of syntectonic plutons (Bickford, 1986; Bickford et al., 1989). A zircon date of 1692 f 5 Ma from granulites in the central Wet Mountains (Bickford et al., 1989) may approximate the age of these deformations and of regional metamorphism. The third deformation (D3) may be coincident with the ages of later deformed plutons (1650-1615 Ma). Granulites in the Wet Mountains are similar to those in southern India formed at 5-6 kbar burial pressure in the presence of a C02-rich fluid phase (Lanzirotti, 1988). Associated syntectonic leucogranites are S-type granites similar in occurrence and chemical composition to syncollisional S-type granites from the Hercynian and Himalayan orogenic belts, consistent with a similar origin for the Wet Mountains granites. Although poorly known, the Wet Mountains assemblage appears to underlie the Proterozoic core of the central Sangre de Cristo Mountains as far south as northern New Mexico (Figs. 1 and 3). A sheared contact between this assemblage and the Dubois terrane is poorly exposed in a region just south of
Proterozoic terranes and continental accretion in southwestern North America
461
Fig. 3. Distribution of Early Proterozoic terranes and overlap assemblages in south-central Colorado. Dashed lines are inferred terrane or overlap assemblage boundaries.
the Colorado-New Mexico state line (16, Fig. 1) (Grambling e t a]., 1988). The Garell Peak pluton, emplaced a t about 1670 Ma (Bickford e t al., 1989), may have been intruded along the boundary of the Wet Mountains and the CochetopaSalida assemblages (5, Fig. 1 and Fig. 3). Distinct xenolith populations of each assemblage can be traced inwards to the central part of the pluton (Thacker, 1988). The nature of the pre-intrusive boundary, however, remains problematic. The boundary between the Wet Mountains and Idaho Springs-Black Canyon assemblages is probably not preserved due to intrusion with granites, and xenolith populations are not distinctive enough to identify its position. It is possible that the Wet Mountains and Idaho Springs-Black Canyon depositories were part of the same continental back-arc basin system.
462
K.C. Condie
Cochetopa-Salida assemblage The Cochetopa-Salida assemblage is exposed in the Gunnison and Salida areas in west-central Colorado (Figs. 1 and 3). Although the contacts of this assemblage are sheared or intruded by granitoids, it is likely that it was deposited at least in part on the underlying Dubois terrane (Knoper, 1992). Most Cochetopa-Salida supracrustal successions are composed of a bimodal mafic and felsic volcanic assemblage including up to 65% volcaniclastic sediments (Afifi,1981; Condie and Nuter, 1981; Shonk, 1984; Boardman, 1986; Knoper, 1992). Primary textures and structures, which are remarkably well preserved, include pillows, hyaloclastic breccias, graded and cross beds, lapilli, and many small delicate bedding features (Boardman, 1986). In addition, diabase and gabbro sills and dikes are important in some sections, and volcanics with komatiitic compositions (up to 24% MgO) have been reported at one locality (Sauer and Boardman, 1988). In the Gunnison area (2 and 3, Fig. l),thick, well-preserved submarine ash flow tuffs are important in the succession (Hedlund and Olson, 1981; Condie and Nuter, 1981; Afifi, 1981). In a few areas, such as the northern Sange de Cristo Mountains south of Salida, Colorado (Fig. 3), andesites are also important (Thacker and Condie, 1986; Thacker, 1988). Geological and geochemical data are consistent with an origin for the Cochetopa-Salida assemblage in a continental-margin arc system (Boardman and Condie, 1986; Condie, 1986; Knoper and Condie, 1988; Wortman et al., 1990; Knoper, 1992). U-Pb zircon ages from felsic volcanics and associated granites from the Cochetopa-Salida assemblage fall in the range of 1740 to 1730 Ma, although a few as low as 1700 Ma have been reported (Bickford and Boardman, 1984; Bickford et al., 1989). The boundary of the Cochetopa-Salida assemblage with the Dubois terrane is exposed at two locations, where it is a major shear zone (Knoper et al., 1991; Knoper, 1992) (Fig. 3). The Dubois shear zone and the Gold Creek shear zone are up to 500 m wide. Because amphibolite-grade metamorphic rocks occur on both sides of these shear zones, it is unlikely that vertical offset is significant. The Dubois shear zone is folded by F2 (see below) and cut by 1720 Ma granitoids indicating tectonic juxtaposition of the two terranes at 1730-1720 Ma. The Dubois terrane is multiply deformed, whereas only two major deformations are recorded in the Cochetopa terrane (Knoper, 1992). The oldest Cochetopa deformation is characterized by tight, isoclinal recumbent folds (F1), which verged to the NW prior to rotation by F2. F1 folding can be bracketed between 1730 f 7 Ma, the age of the youngest folded volcanic rocks, and 1721 k 7 Ma, the age of oldest post-F1 plutons (Knoper et al., 1991). F2 folding is broadly synchronous with syntectonic plutons at 1720-1713 Ma (Wortman et al., 1990; Knoper et al., 1991). It also seems to roughly coincide with widespread syntectonic plutonism and regional metamorphism in west-central Colorado. F1 may have resulted from the closing of a NW-trending continental-margin back arc basin in central Colorado, and F2 may reflect foreland deformation resulting from the collision of the Pecos terrane in northern New Mexico at 1720-1700 Ma.
Proterozoic tewanes and continental accretion in southwestern North America
463
Manzano assemblage The Manzano assemblage includes numerous remnants of largely metasediment successions in central New Mexico (Fig. 1). It also appears to unconformably overlie the Pecos and Dubois terranes in northern New Mexico and southwestern Colorado (Manzano Extension, Fig. 1). Metasediments are chiefly pelites and thick, cross-bedded quartzites such as the Ortega Quartzite in northern New Mexico. Also found are felsic ash flow tuffs, amphibolites (chiefly mafic dikes and sills), conglomerates, quartz wackes and volcaniclastic sediments (Condie and Budding, 1979; Kent, 1980; Cavin et al., 1982; Grambling and Codding, 1982; Condie, 1986; Soegaard and Eriksson, 1986; Alford, 1987; Gabelman, 1988; Robertson et al., 1991). The Manzano assemblage has been subjected to at least three periods of deformation and associated metamorphism (Bauer, 1982, 1984; Grambling, 1986; Williams and Grambling, 1987; Grambling et al., 1988; Daniel et al., 1990). Peak metamorphic conditions of 475-550°C and 12-16 km burial depth appear to have been reached after N to NW-directed ductile thrusting, and structural studies indicate a north to northwest compression with some thrusts moving as much as 150 km (Grambling et al., 1988). Detailed sedimentological studies in northern New Mexico indicate the 1700 Ma Ortega Group was deposited in a shallow marine shelf environment influenced by tidal, storm and wave processes (Soegaard and Eriksson, 1985, 1986). The Vallecito Conglomerate in southwestern Colorado (15, Fig. 1) appears to represent an alluvial fan complex built by high-gradient braided streams and debris flows (Ethridge et al., 1984). U-Pb zircon ages from ash flow tuffs in the Manzano assemblage are 1700-1650 Ma with most falling at about 1680-1650 Ma (Bowring and Condie, 1982; Bowring et al., 1983). Syn-tectonic granitoids range in age from about 1700 to 1650 Ma and 1600-1350 Ma post-tectonic granites are widespread. The oldest rocks known from southern New Mexico are 1750 Ma gneisses of limited extent in the southern San Andres Mountains (21, Fig. 1) (Roths, 1991). They may represent basement to the Manzano assemblage in this area. The tectonic setting of the Manzano assemblage has been elusive because of variable and mixed lithologic packages (Condie and Budding, 1979; Condie, 1986). Studies of sediments in the Hembrillo succession in the San Andres Mountains of southern New Mexico (21, Fig. 1) indicate a dual provenance (Alford, 1987). Massive quartzites, arkoses, and pelites appear to be derived from cratonic sources to the north or northwest, whereas quartz wacke turbidites and other volcaniclastic sediments interbedded with the quartzites and pelites, reflect southern, largely volcanic arc sources. Such relationships are compatible with a continental back-arc basin with cratonic and arc sediments derived from opposite margins of the basin. Similar, although more structurally complex successions occur in the Manzano Mountains (9, Fig. 1). No evidence of oceanic crust in the proposed back-arc basin is found in the Manzano assemblage. Manzano sediments in northern New Mexico and southwestern Colorado lack the volcaniclastic rocks and seem to record marginal cratonic-basin deposition.
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K.C. Condie
Most or all of the boundaries of the Manzano assemblage are tectonic contacts (Grambling et al., 1988; Robertson et al., 1991; note that the Manzano assemblage as used herein includes both the Manzano and 2uchas terranes of these authors). When exposed, such in the northern Pedernal Hills (8, Fig. l ) , southern Sangre de Cristo Mountains (7, Fig. l), the Tusas Range (6, Fig. l ) , and Cimarron Hills (Grambling and Dallmeyer, 1990), contacts are shear zones. Studies of the Uncompahgre Formation in southern Colorado (15, Fig. l), which is tentatively assigned to the Manzano assemblage, conclude that it is separated from the underlying Dubois terrane by a low-angle thrust (Tewksbury, 1985; Harris et al., 1987). Franklin Mountains assemblage The Franklin Mountains assemblage, exposed in and near the Franklin Mountains near El Paso, Texas (Fig. l), is composed of the Castner Marble, Mundy Breccia, Lanoria Quartzite, and Thunderbird Group, in ascending stratigraphic order, and related intrusive granites (Harbour, 1972; Thomann, 1980). Sedimentological studies of the Lanoria Quartzite indicate a southernly source (Seeley, 1991). The Franklin’s section has escaped significant deformation and metamorphism, and may have been deposited in a submerged continental rift or transpressional basin that developed north of the foreland basin in which the Hazel Formation in the Van Horn area was deposited. A subduction zone geochemical component in the volcanic rocks and associated granites (Norman et al., 1987) is probably inherited from a subcontinental lithospheric source. U-Pb zircon ages from felsic igneous rocks in the Franklin Mountains indicate an age of 1150-1135 Ma (Copeland and Bowring, 19SS), somewhat younger than suggested by previous zircon data (Wasserburg et al., 1962). Although the lateral extent of the Franklin Mountains assemblage is unknown, drillcore data suggest a limited geographic area (Dennison and Hetherington, 1969) (Fig. 1).
DISCUSSION
Province boundaries Of the recognized province boundaries, only the Cheyenne shear zone (1, Fig. 1) and the Gneiss Canyon shear zone (13, Fig. 1) have been described in any detail (Karlstrom and Houston, 1984; Duebendorfer and Houston, 1986; Duebendorfer, 1988; Chamberlain and Bowring, 1990; Karlstrom and Bowring, 1991, 1992). The general features of these shear zones were summarized in previous sections. The western boundary of Precambrian basement in Nevada is the edge of the Cordilleran accretional terranes, and the southwestern extent of the Mojave Province is complicated by younger fault systems. The Mojave-Sonora megashear
Proterozoic terranes and continental accretion in Southwestern North America
465
(Fig. 1) separates Proterozoic rocks with zircon ages of 1800-1725 Ma on the southwest side of the fault from those with ages 1700-1650 Ma on the northeast side (Silver and Anderson, 1974; Anderson and Silver, 1981). Similarities between Late Proterozoic and Paleozoic stratigraphic columns in the Death Valley area of California (18,Fig. 1) with those near Caborca in Mexico (23, Fig. 1) suggest a minimum of 700-800 km of left lateral offset along this fault (occurring chiefly in the Jurassic). Restoration of this offset brings cratonic metasediments in Sonora into the western Mojave Province. It is likely that the Mojave Province extended at least as far west as the San Andreas fault. The boundary between the Yavapai and Mazatzal Provinces appears to cross Arizona and northern New Mexico in a northeasterly direction (Silver, 1968), possibly extending into Colorado. Although the boundary is exposed as the Moore Gulch fault in central Arizona (10, Fig. l), this is a young fault and does not represent the original contact (Karlstrom et al., 1987). As previously mentioned, the boundary may be defined by an isostatic gravity anomaly extending from eastern Arizona into New Mexico (25, Fig. 1). A similar and even better defined gravity anomaly in southeastern Colorado (26, Fig. 1) also may be the southern boundary of the Yavapai Province (Simpson et al., 1986; Lanzirotti, 1988). This anomaly can be traced into northwestern Kansas with an arm extending into southern Nebraska as far east as the Mid-Continent Rift System. Although it enters northern New Mexico near Raton, it is not clear how or if it crosses the Rio Grande Rift, nor if it connects with the isostatic gravity anomaly in northwestern New Mexico. Proterozoic supracrustal rocks in the Tusas Mountains (6, Fig. l), %os Range (16) and Cimarron Hills in northern New Mexico (Grambling and Dallmeyer 1990) are known or probable portions of the Dubois terrane, and hence if the boundary crosses the Rio Grande Rift, it must pass south of these exposures as suggested by the E-W dashed line in Fig. 1. The Yavapai-Mazatzal boundary in northern New Mexico and southern Colorado is probably a suture between the Yavapai Province and the Pecos and related terranes southeast of the Yavapai Province. Although the lateral extent of these terranes is problematic, if the gravity anomaly extending into Kansas reflects a suture, the Pecos terrane must extend at least into central Kansas. The southern boundary of the Mazatzal Province must extend at least as far south as the New Mexico-Mexico border as dictated by the presence of -1650 Ma crustal xenoliths in young volcanics west of El Paso, lkxas (Reid et al., 1985). This boundary may be responsible for the Abilene gravity anomaly in West Texas (22, Fig. l), a feature that is caused by a pronounced change in rock type in the Precambrian basement (Nicholas and Rozendal, 1975). This anomaly passes through the southeast corner of New Mexico and south of El Paso, Texas, and rocks 21500 Ma in age are not recognized south of the anomaly. The actual suture between the Mazatzal and Grenville Provinces is not exposed in West Exas and its location remains problematic. Although it could be reflected by the Abilene gravity anomaly, it may also lie farther south, perhaps buried by the thick foreland basin sediments of the Hazel Formation in the Van Horn area of West
466
K.C. Condie
Xxas. The Streeruwitz thrust exposed near Van Horn, Texas (King and Flawn, 1953) is a foreland thrust, and not the actual suture between the two provinces.
Summary of tectonic settings Four Proterozoic terranes in the southwestern United States have clear oceanic arc affinities and are allochthonous relative to the Archean Wyoming Craton (Xible 1). Of these, the Green Mountain and Dubois terranes were accreted to the craton between about 1780 and 1750 Ma and the Ash Creek and Pecos terranes at about 1700 Ma (Condie, 1986; Karlstrom and Bowring, 1992). It is not known if the Green Mountain and Dubois terranes were part of the same arc system or remnants of different arcs. The Mojave Province is also allochthonous and may represent an accreted microcraton. Lithologic and geochemical characteristics of five of the Proterozoic terranes or overlap assemblages suggest they are remnants of back-arc basins associated with continental-margin arc systems that range in age from 1740 and 1270 Ma (Dble 1). The Manzano assemblage and the Alder and Dos Cabezas-Pinal terranes show evidence of significant cratonic sources for sediments in the back-arc basins, and the Franklin Mountains assemblage may be a remnant of an aborted continental rift or transpressional basin related to the Grenville collision.
Collisionnl timing At least five major Proterozoic collisions are required to explain the relative timing of compressional deformation and plutonism in the Southwest (Bble 2; Fig. 4). The earliest and best documented collision is that of the Green Mountain terrane with the Wyoming Craton causing the Cheyenne orogeny. U-Pb zircon dates from pre- to post-collisional granites in the Cheyenne shear zone constrain this collision to between 1780 and 1765 Ma with the last major motion occurring at about 1750 Ma (Premo and Van Schmus, 1989). The Dubois terrane collided with the Green Mountain terrane either before or after its collision with the Wyoming Craton (Fig. 4). The Ivanpah (Yavapai) orogeny at about 1700 Ma reflects the collision of the Mojave and Yavapai Provinces. Also, at this time the various terranes (and blocks) TABLE 2 Timing of major Proterozoic collisions in southwestern North America Collision 1. 2. 3. 4. 5.
Green Mountain Mojave Pecos Oklahoma Carrizo Mountains
Orogeny
Age (Ma)
Cheyenne Ivanpah (Yavapai) Pecos Mazatzal Greenville
1780-1765 1700 1700 1650 1200
Proterozoic teiranes and continental accretion in southwestern North America 1 .c
-
T
m
0
467
1.5
0 W
Q
I.3
T
I65
I .70
I80
I a5
LI G-Ai
11-
\
I
I \
190
WYOMING CRATON
I
Fig. 4. Cladogram for Proterozoic terrane accretion in southwestern North America. Constructed after the methods of Young (1986) and Hoffman (1989). Boxes indicate terrane U-Pb zircon ages and line intersections with the Wyoming Craton are times of terrane collision.
in the Yavapai and Mazatzal Provinces in central and western Arizona collided, amalgamated and were sutured to the North American continent (Karlstrom and Bowring, 1991) (Fig. 4). The Pecos collision, which resulted in the Pecos orogeny, must be younger than Pecos volcanism (1720 Ma) yet older than widespread late syntectonic to post-tectonic granites (1700-1670 Ma) in central and southern Colorado and northern New Mexico. Variably foliated 1700-1670 Ma plutons in Colorado (Bickford et al., 1989; Hallett, 1990; Hallett and Condie, 1990) and in the Nacimiento Mountains in New Mexico (Woodward, 1987) may represent late syntectonic plutons associated with this collision. U-Pb zircon ages from late syntectonic plutons and a probable metamorphic zircon age of 1692 f 5 Ma from the Wet Mountains in Colorado suggest the collision occurred about 1700 Ma ago. The collision of the Yavapai and Mojave Provinces must be younger than pre-deformational plutons (1725-1710 Ma) and older than post-tectonic plutons in southeastern California (1690-1660 Ma) (Wooden et al., 1988; Wooden and
468
K.C. Condie
Miller, 1990). High-grade metamorphism and cratonic sediments in the western Mojave Province and arc rocks in the Yavapai Province are consistent with a microcraton-arc collision, with the Mojave Province on the descending plate and the Hualapai terrane caught between the Mojave Province and the Ash Creek terrane. Collisions on both sides of the Yavapai Province at about 1700 Ma might explain the widespread syn- to early post-tectonic plutonism of this age throughout the province. One or more post-1700 Ma collisional events in southwestern North America is required by NW-verging thrust faults and folds in the Alder and Dos CabezasPinal terranes and in Manzano assemblage (Karlstrom et al., 1987; Grambling and Ward, 1987; Grambling et al., 1988). In central Arizona, one event is younger than the 1700 Ma Alder and Mazatzal Groups yet older than the post-tectonic Sunflower granite (1640 Ma) (Karlstrom et al., 1990). A 1650 Ma zircon date from the syntectonic Sommer granite in southeastern Arizona (Erickson and Bowring, 1990) may correspond to the age of this collisional event. This collision is tightly constrained in central New Mexico where it must follow eruption of felsic volcanics with U-Pb zircon dates of 1664 Ma, yet precede emplacement of granite at 1654 Ma (Bowring et al., 1983; Bauer et al., 1991). It may be possible to accommodate the results from Arizona and New Mexico by a single collisional event at about 1650 Ma, generally thought to be responsible for the Mazatzal orogeny (Xible 2; Fig. 4). The size of the colliding landmass, herein referred to as the Oklahoma Province, and the location of the suture are unknown. However, the distribution of 2.0-1.8 Ga Nd model ages for granite sources in the southern midcontinent area (Nelson and DePaolo, 1985) suggests that a small continent (perhaps as large as Greenland) may have accreted at this time. The final collisional event in West Texas must have occurred after 1280 Ma, the age of the Allamore Formation, probably at about 1200 Ma. Some northward thrusting must have continued until after 1194 Ma, the age of a granite boulder in the Hazel Formation (Soegaard et al., 1991; Callaham and Soegaard, 1991). This event reflects the collision and accretion of the southern Grenville Province to North America. Although most of North America was assembled between about 2.0 and 1.8 Ga (Hoffman, 1988, 1989), as shown by the cladogram in Fig. 4, southwestern North America was largely assembled between 1.8 and 1.65 Ga. The Yavapai Province in Colorado accreted to the Wyoming Craton at 1780-1750 Ma. The Yavapai Province in Arizona, the Mojave Province, and the Mazatzal Province accreted to the southwestern margin of North America at about 1700 Ma and the Oklahoma Province at 1650 Ma. Lastly, the Grenville Province collided and accreted to the continent at about 1200 Ma. Nd isotopic data suggest that crustal formation ages are not more than 100-200 Ma older than zircon ages (Nelson and DePaolo, 1984, 1985; Bennett and DePaolo, 1987; Wortman et al., 1990), and thus that crustal additions during the Proterozoic in southwestern North America were largely juvenile.
Proterozoic terranes and continental accretion in southwestern North America
469
Cratonization and cooling history Cratonization involves intracrustal melting, uplift, and erosion following a collisional event. Thermal models as well as isotopic age constraints indicate that cratonization should begin within 30 Ma of cessation of collision (Thompson and Ridley, 1987). This certainly seems to be the case in the southern Yavapai Province where northward-derived cratonic sediments were deposited within 20 Ma of the 1700 Ma collisions in some areas, such as in the Alder terrane (Condie et al., 1991). Uplift and erosion continued in the southern Yavapai Province until at least 1650 Ma and introduced large volumes of first-cycle quartz into back-arc basins that extended from central and southeastern Arizona through New Mexico and probably into the Mid-Continent region. Cratonization was probably well advanced before collision of the Oklahoma Province at about 1650 Ma. 40Ar/39Arcooling ages from metamorphic muscovites and hornblendes from the Manzano assemblage in New Mexico reflect a complex uplift and cooling history. The oldest cooling event following peak metamorphic conditions in northern New Mexico is 1440-1400 Ma followed by cooling events at 1360-1340, 1325-1300, and 1000-960 Ma (Dallmeyer et al., 1990; Thompson et al., 1991). Decompressive P-T-t paths characterize the first cooling event (Daniel et al., 1990) consistent with metamorphism in an extensional tectonic regime. The latter three events may coincide with minor ductile thrusting. It is noteworthy that the older two cooling ages (1440-1400 and 1360-1340 Ma) correspond to widespread anorogenic plutonism in the southwestern and south-central United States and may reflect widespread crustal heating by plutons at these times. In Arizona, peak metamorphic conditions associated with terrane accretion at 1700 Ma outlasted deformation by more than 200 Ma (Bowring and Karlstrom, 1990). Terrane boundaries in this area were repeatedly reactivated from the Proterozoic through Tertiary time. The apparent absence of Archean detrital zircons in Proterozoic quartzites of Colorado and New Mexico (Aleinikoff et al., 1985) is an important constraint on river drainage systems during the Early Proterozoic of the southwestern United States. It would appear that a highland must have existed at approximately 40"N latitude (present-day reference) for more than 100 Ma (1780 Ma to 1650 Ma). Archean sources must have drained to the north and only Early Proterozoic sources to the south and southeast during this period of time. Comparison with Cordilleran terranes Although only small remnants of Proterozoic terranes in southwestern North America are preserved, it is of interest in terms of continental accretion to compare these to Phanerozoic terranes in North America. The most obvious comparison is with the Mesozoic-Tertiary Cordilleran terranes in western North America. More than one tectonic setting is commonly represented in Cordilleran terranes, and within a given terrane tectonic setting may change with time. For
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instance, in the Alexander terrane, tectonic setting changed from an arc in the Paleozoic to a continental rift in the Piassic, and back to an arc in the Cretaceous (Plafker et al., 1989). In most of the Southwest Proterozoic terranes, on the other hand, only one tectonic setting is represented in the succession of rocks preserved. Also, most of the Cordilleran terranes have affinities with ophiolites and oceanic arcs, and submarine plateaus and seamounts are well represented (Ernst, 1988; Plafker et al., 1989; Coney, 1989). With the exception of one ophiolite remnant in the Alder terrane, ophiolites and remnants of oceanic plateaus and seamounts are not recognized in the Proterozoic terranes of the Southwest. Also missing in the Proterozoic terranes are pelagic sediments. In contrast to most Cordilleran terranes, most of the SW Proterozoic terranes appear to represent evolved island arcs or continental-margin arcs. Another difference is that the earliest accreted Cordilleran terranes (such as the Roberts Mountains allochthon) represent continental rise/slope assemblages (Tbrner et al., 1989), whereas the earliest accreted Proterozoic terranes are oceanic arcs. Early Proterozoic terranes of the Southwest and Meso-Cenozoic Cordilleran terranes are similar in that they are composed mostly (290%) of juvenile crust as suggested by Nd isotopic data (Nelson and DePaolo, 1985; Samson et al., 1989; Wortman et al., 1990). However, the range in crustal formation ages (i.e., age of extraction from the mantle) is 250-300 Ma in many Cordilleran terranes (mid-Paleozoic to Cretaceous), whereas Southwest Proterozoic crustal formation ages within given terranes are generally 550 Ma. The combined Yavapai and Mazatzal Provinces were formed and accreted to North America in 5150 My. It would appear that rates at which arcs were accreted to the North American craton were considerably greater during the Early Proterozoic than during the Phanerozoic. The accreted Cordilleran Province amounts to about 0.75 x 10' km3 (2.5 x lo6 km2 x 30 km thick) and was accreted chiefly over 120 My (Late Jurassic to Eocene) (Coney, 1989) indicating an accretion rate of about 0.63 km3/a. The corresponding rate for the SW Early Proterozoic provinces (YavapailMazatzal provinces with eastern extensions to the Grenville Front) is 1.4 km3/a (4 x lo6 km x 35 km thick = 1.4 x 10' km3) for a time interval of 100 My (1750-1650 Ma). Thus, it would appear that the rate of Proterozoic continental growth in the southwestern and southern United States was over twice the rate of Phanerozoic growth in the American Cordillera.
ACKNOWLEDGEMENTS
The author's research has been supported by National Science Foundation grants EAR4313735 and EAR-8915232. The manuscript was reviewed and improved from suggestions by J.L. Wooden, T.H. Anderson, W.R. Van Schmus, and Paul Hoffman.
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Premo, W. and Van Schmus, W.R., 1989. Zircon geochronology of Precambrian rocks in southeastern w o m i n g and northern Wyoming and northern Colorado. Geol. SOC.Am., Spec. Pap., 235: 13-32. Ransome, EL., 1904. Geology and ore deposits of the Bisbee quadrangle, Arizona. U.S. Geol. Surv., Prof. Pap., 21,168 pp. Reed, J.C., Jr., 1984. Proterozoic rocks of the Taos Range, Sangre de Cristo Mountains, New Mexico. N.M. Geol. SOC.Field Conf., Guidebook, 35: 179-185. Reed, J.C., Jr., Bickford, M.E., Premo, W.R., Aleinikoff, J.N. and Pallister, J.S., 1987. Evolution of the early Proterozoic Colorado province: constraints from U-Pb geochronology. Geology, 15: 861-845. Reid, M., Hart, S.R. and Padovani, E., 1985. Importance of sedimentary protoliths to the lower crust exemplified by the Kilbourne Hole paragneisses - Sr, Nd and Pb isotope geochemistry. EOS, Trans. Am.Geophys. Union, 66: 1110. Reuss, R.L., 1974. Precambrian quartzite-schist sequence in Wilson Park, Fremont County, Colorado. Mod. Geol., 11: 45-58. Robertson, J.M. and Condie, K.C., 1989. Geology and geochemistry of early Proterozoic volcanic and subvolcanic rocks of the Pews greenstone belt, Sangre de Cristo Mountains, New Mexico. Geol. SOC.Am., Spec. Pap., 235: 119-146. Robertson, J.M. and Moench, R.H., 1979. The P e a s greenstone belt: a Proterozoic volcano-sedimentary sequence in the southern Sangre de Cristo Mountains, New Mexico. N.M. Geol. SOC.Field Conf., Guidebook, 30: 165-173. Robertson, J.M., Grambling, J.A., Mawer, C.K., Bowring, S.A., Williams, M.L., Bauer, P.W. and Silver, L.T., 1991. Precambrian geology of New Mexico. Geol. SOC.Am., DNAG. Precambrian of the United States (in press). Roths, P., 1991. Preliminary results of investigation of Proterozoic outcrops, southern San Andres Mountains, New Mexico. Geol. SOC.Am., Abstr. Progr., 23 (4): 88. Rudnick, R.L., 1983. Geochemistry and tectonic affinities of a Proterozoic bimodal igneous suite, West Texas. Geology, 11: 352-355. Ruiz, J., Patchett, P.J. and Ortega-Gutierrez, E, 1988. Proterozoic and Phanerozoic basement terranes of Mexico from Nd isotopic studies. Geol. SOC.Am., Bull., 100 274-281. Ruiz, J., Patchett, P.J. and Ortega-Gutierrez, E, 1990. Proterozoic and Phanerozoic terranes of Mexico based on Nd, Sr and Pb isotopes. Geol. SOC.Am., Abstr. Progr., 22 (7): A113. Samson, S.D.,McClelland, W.C., Patchett, P.J., Gehrels, G.E. and Anderson, R.G., 1989. Evidence from Nd isotopes for mantle contributions to Phanerozoic crustal genesis in the Canadian Cordillera. Nature, 337: 705-709. Sauer, P.E. and Boardman, S.J., 1988. Komatiitic trends in Early Proterozoic volcanic rocks in central Colorado. Geol. SOC. Am., Abstr. Progr., 2Q: 467. Seeley, J.M., 1991. Middle Proterozoic siliciclastic shelf sediments of the Franklin Mountains, El Paso County, Texas. Geol. SOC. Am., Abstr. Progr., 23 (4): 92. Shadel, C.A., 1982. Geology and Geochemistry of the Proterozoic Metavolcanic and Volcaniclastic Rocks of the Green Mountain Formation, Sierra Madre Range, Wyoming. M.S. Thesis, N.M. Inst. Min. Technol., 164 pp. Shonk, K.N., 1984. Stratigraphy, Structure, Tectonic Setting and Economic Geology of an Early Proterozoic Metasedimentary and Metavolcanic Sequence, South Beaver Creek Area, Gunnison and Saguache Counties, Colorado. M.S. Thesis, Colorado School of Mines, Golden, Colo., 327 pp. Silver, L.T., 1963. The use of cogenetic uranium-lead isotope systems in geochronology, In: Radioactive Dating 1962. Int. Atomic Energy Agency, Vienna, pp. 279-285. Silver, L.T., 1966. U-Pb isotope relations and their historical implications in Precambrian zircons from Bagdad, Arizona. Geol. SOC.Am., Spec. Pap., 101: 420. Silver, L.T, 1968. Precambrian batholiths of Arizona. Geol. Soc. Am., Spec. Pap., 121: 558-559.
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Silver, LT, 1978. Precambrian formations and Precambrian history in Cochise County, southeastern Arizona. N.M. Geol. SOC.Field Conf., Guidebook, 29 157-163. Silver, L.T and Anderson, TH., 1974. Possible left-lateral early to middle Mesozoic disruption of the southwestern North American craton margin. Geol. SOC.Am., Abstr. Progr., 6 955-956. Silver, L.T and Barker, E, 1968. Geochronology of Precambrian rocks of the Needle Mountains, southwestern Colorado. Geol. SOC.Am., Spec. Pap., 115: 204-205. Silver, L.T., Conway, C.M. and Ludwig, K.R., 1986. Duplications pof a precise chronology for Early Proterozoic crustal evolution and caldera formation in the Tonto Basin-Mazatzal Mountains region, Arizona. Geol. SOC.Am., Abstr. Progr., 18: 413. Simpson, R.W., Jachems, R.C. and Blakely, R.J., 1986. New isostatic residual gravity map of the conterminous United States. J. Geophys. Res., 91: 8348-8372. Slatt, R.M., Heintz, G.M., Lowry, P.H. and O’Hara, P.F., 1978. Precambrian Pike’s Peak iron formation, central Arizona. Ariz. Bur. Geol. Miner. Tech., Spec. Pap., 2: 73-82. Snyder, G.L., Brandt, E.L. and Smith, VC., 1988. Precambrian petrochemistry of the northern Park Range, Colorado and its implications for studies of crustal derivation. U.S. Geol. Surv., Prof. Pap., 1343,116 pp. Soegaard, K. and Eriksson, K.A., 1985. Evidence of tide, storm and wave interaction on a Precambrian siliciclastic shelf: the 1700 Ma Ortega Group, New Mexico. J. Sediment. Petrol., 55: 672-684. Soegaard, K. and Eriksson, K.A., 1986. ’Ikansition from arc volcanism to stable-shelf and subsequent convergent margin sedimentation in northern New Mexico from 1.76 Ga. J. Geol., 94: 47-66. Soegaard, K., Callaham, D.M., Nielsen, K.C. and Roths, P.J., 1991. Sedimentary and tectonic history of middle to Late Proterozoic successions near Van Horn, West Texas. Geol. SOC.Am., Abstr. Progr., 23 (4): 96. Swift, P.N., 1982. Precambrian Metavolcanic Rocks and Associated Volcanogenic Mineral Deposits of the Southwestern Sierra Madre, Wyoming. M.A. Thesis, University of Wyoming, Laramie, Wyo., 61 PP. Tewksbury, B.J., 1985. Revised interpretation of the age of allochthonous rocks of the Uncompahgre Formation, Needle Mountains, Colorado. Geol. SOC.Am., Bull., 96: 224-232. Thacker, M.S., 1988. Geology and Geochemistry of Early Proterozoic Supracrustal Rocks from the Northern Sangre de Cristo Mountains, Central Colorado. M.S. Thesis, N.M. Inst. Min. Technol., 129 PP. Thacker, M. and Condie, K.C., 1986. Early Proterozoic supracrustal rocks from the northern Sangre de Cristo Mountains and adjacent areas, Colorado. Int. Field Conf., Proterozoic Geology and Geochemistry, Central Colorado, July 1986, Abstract Volume, p. 122. Thomann, W.E, 1980. Ignimbrites, trachytes and sedimentary rocks of the Precambrian Thunderbird Group, Franklin Mountains, El Paso, Texas. Geol. SOC.Am., Bull., 9 2 94-100. Thomas, W.M., Clarke, H.S., Young, E.D., Orrell, S.E. and Anderson, J.L., 1988. Proterozoic high-grade metamorphism in the Colorado River region, Nevada, Arizona and California. In: W.G. Ernst (Editor), Metamorphism and Crustal Evolution of the Western United States. Prentice-Hall, Englewood Cliffs, N.J., pp. 526-537. Thompson, A.B. and Ridley, J.R., 1987. Pressure-temperature-time histories of orgenic belts. Philos. Trans. R. SOC.London, Ser. A, 321: 27-45. Thompson, A.G., Grambling, J.A., Dallmeyer, R.D., Mawer, C.K. and Daniel, C.G., 1991. A polyphase Middle Proterozoic tectonometamorphic history in ccntral New Mexico: structural, petrologic and 40Ar/39Arisotopic evidence. Geol. SOC.Am., Bull. (in press). Trevena, A.S., 1979. Studies in Sandstone Petrology: Origin of the Precambrian Mazatzal Quartzite and Provenance of Detrital Feldspar. Ph.D. Dissert., University of Utah, Salt Lake City, Utah, 390 pp.
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Turner, R.J.W., Madrid, R.J. and Miller, E.L., 1989. Roberts Mountains allochthon: stratigraphic comparison with lower Paleozoic outer continental margin strata of the northern Canadian Cordillera. Geology, 17: 341-344. 'heto, O., 1987. Rock units of the Precambrian basement in Colorado. U.S. Geol. Sum., Prof. Pap., 1321-A, 54 pp. Vance, R.K., 1989. Geochemistry and Tectonic Setting of the Yavapai Supergroup, West Central Arizona. Ph.D. Dissertation, N.M. Inst. Min. Technol.. 461 pp. Vance, R.K. and Condie, K.C., 1986. Geochemistry and tectonic setting of the early Proterozoic Ash Creek Group, Jerome, Arizona. Geol. SOC.Am., Abstr. Progr., 18, pp. 419. Vance, R.K. and Condie, K.C., 1987. Geochemistry and tectonic setting of volcanic rocks from the early Proterozoic Big Bug Group, Bradshaw Mountains, Arizona. Geol. SOC.Am., Abstr. Progr., 19: 340. Walker, N.W., 1988. U-Pb zircon evidence for 1305-1231 Ma crust in the Llano uplift, central Texas. Geol. SOC.Am., Abstr. Progr., 2 0 A205. Walker, N., Mosher, S. and Carlson, W.D., 1990. Proterozoic evolution of the Llano uplift, central Texas. Geol. SOC.Am., Abstr. Progr., 22 (7): A113. Warnke, D.A., 1969. A geologic investigation of the Halloran Hills, central Mojave Desert, California. Geol. Rundsch., 5 8 998-1047. Wasserburg, G.J., Wetherill, G.W., Silver, LT and Flawn, P.T., 1962. A study of the ages of the Precambrian of Texas. J. Geophys. Res., 6 7 4021-4047. Wells, J.D., 1967. Geology of the Eldorado Springs quadrangle, Boulder and Jefferson Counties, Colorado. U.S. Geol. Surv., Bull., 1221-D. White, C.A. and Foster, C.T., Jr., 1987. Proterozoic metasediments near Lester Mountain, northern Park Range, Colorado. Geol. SOC.Am., Abstr. Progr., 19: 342 Williams, M.L. and Grambling, J.A., 1987. Mid-crustal exposure of a Proterozoic orogenic belt. Geol. SOC.Am., Abstr. Progr., 1 9 890-891. Wilson, E.D., 1939. Precambrian Mazatzal revolution in central Arizona. Geol. SOC.Am., Bull., 50: 1113-1164. Wooden, J.L. and Aleinikoff, J.N., 1991. Early Proterozoic isotopic provinces in the southwestern U.S. Geol. SOC.Am., Abstr. Progr., 23 (4): 107. Wooden, J.L. and Miller, D.M., 1990. Chronologic and isotopic framework for Early Proterozoic crustal evolution in the eastern Mojave desert region, SE California. J. Geophys. Res., 95: 20,133-20,146. Wooden, J.L. and Miller, D.M., 1991. Early Proterozoic geologic history of the Mojave crustal province Geol. SOC.Am., Abstr. Progr., 23 (4): 108. Wooden, J., Miller, D. and Elliott, G., 1986. Early Proterozoic geology of the northern New York mountains, SE California. Geol. SOC.Am., Abstr. Progr., 18 424. Wooden, J.L., Miller, D.M. and Howard, K.A., 1988a. Early Proterozoic chronology of the easterr Mojave Desert. Geol. SOC.Am., Abstr. Progr., 20: 243. Wooden, J.L., Stacey, J.S., Howard, K . k , Doe, B.R. and Miller, D.M., 1988b. Pb isotopic evidencc for the formation of Proterozoic crust in the southwestern United States. In: W.G. Ernst (Editor) Metamorphism and Crustal Evolution of the Western United States, Rubey Volume VII. Prentice Hall, Englewood Cliffs, N.J., pp. 69-86. Woodward, L.A., 1987. Geology and mineral resources of Sierra Nacimiento and vicinity, New Mexico N.M. Bur. Mines Miner. Resour., Mem., 42, 84 pp. Woodward, L.A., Parchman, MA., Edwards, D.L. and Husler, J.W, 1979. Stratigraphy and mineraliza tion of Hell Canyon greenstone belt, New Mexico. N.M. Geol. SOC.Field Conf., Guidebook, 30, PF 189-195. Wortman, G.L., Coleman, D.S. and Bickford, M.E., 1990. Timing of arc accretion and deformation i early Proterozoic volcanogenic rocks, central Colorado. Geol. SOC.Am., Abstr. Progr., 22 (7): A26;
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Wright, L.A., 1974. Geology of the southeast quarter of Tecopa quadrangle, Inyo County, California. Calif. Div. Mines, Map Sheet 20. Wyman, W.F., 1980. Precambrian Geology of the Cow Creek Ultramafic Complex, San Miguel County, New Mexico. M.S. Thesis, N.M. Inst. Min. Technol., 125 pp. Young, E.D., Anderson, J.L., Clarke, H.S. and Thomas, W.M., 1989. Petrology of biotite-cordieritegarnet gneiss of the McCullough Range, Nevada, I. J. Petrol., 30 39-60. Young, G.C., 1986. Cladistic methods in Paleozoic continental reconstruction. J. Geol., 94: 523-537.
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Chapter 13
ISOTOPIC STUDIES OF PROTEROZOIC CRUSTAL GROWTH AND EVOLUTION P.J. PATCHETT
INTRODUCTION
This section of the volume is intended to summarize the state of knowledge of the origins and evolutionary processes of Proterozoic crust as reflected by radiogenic isotopic studies. Recent years have seen a very substantial increase in the number of geochemical and isotope geochemical studies carried out on Proterozoic rocks. In particular, isotopic studies, whether for the purposes of dating or to determine genetic aspects of crustal growth and evolution, have changed in the past 15 years from a somewhat exotic approach applied only occasionally, to a widespread tool of Precambrian research. Because of this, it would be a formidable task to review in detail all the contributions made to the field in recent years. Many new studies of radiogenic isotopes in Proterozoic crust appear every year, but the field is nowhere remotely near to characterization of the origins of Proterozoic crust over the whole globe. For these reasons, a detailed region-by-region review would seem to have a reduced purpose as well as a short useful lifetime. Consequently, this chapter summarizes the state of knowledge of Proterozoic crust origins and development only in general terms, although a large number of specific references are given. Much of the chapter concentrates on the methodology by which large rock terranes can be characterized isotopically. In this context, the advantages of the Sm-Nd isotopic approach, but also pitfalls and problems of methodology and data interpretation are discussed in some detail. XI some extent, the same rationale, advantages and problems occur also in Sr, Pb, Hf or 0 s isotopic study of Precambrian rocks. However, it is the Sm-Nd isotopic system that has been most widely applied, and whose employment is increasing enormously. Therefore, this review concentrates on Nd isotopes both in terms of the analytical and interpretive framework and the state of knowledge of Proterozoic crustal development worldwide.
GLOBAL COVERAGE OF Nd ISOTOPIC DATA
Large rock terranes, or even small areas, cannot be characterized isotopically in terms of rock origins and evolution without rock samples being taken on some
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rational basis from the available outcrops. Samples for Nd isotopic characterization could be collected on a simple grid pattern, say one at every kilometer intersection. Such an approach could lead to a reasonable averaging over a vast craton if hundreds of such samples were analyzed. In general, though, such blind sampling is fraught with hazards, and it is regarded as more meaningful to collect samples on the basis of available understanding of field relations, tectonics and terrane divisions in the region to be studied. Additionally, it is then often the case that large regions can be characterized over their areal extent and geologic history using far fewer samples than would be needed for a grid-pattern characterization. For the above reasons, Nd or other isotopic studies can only be productively performed in contexts where sufficient study has been carried out to allow collection of representative and meaningful samples. The requirements are (1) reconnaissance mapping of wide areas; (2) detailed mapping of critical areas that will be more densely sampled; ( 3 ) existence of some level of tectonic synthesis and/or terrane divisions for the region, and (4) geochronologic constraints giving at least the approximate timing of events that stabilized the crust. In the Proterozoic, meaningful chronologic constraints are mainly provided by U-Pb zircon geochronology. Usually, it takes a number of years to accumulate the level of field mapping and tectonic/chronologic understanding implied by (1)-(4) above. Consequently, studies of Nd isotopic characterization were first performed during 1978-1990 in regions where the requisite background already existed. Many of such regions have been covered now. They lie mainly in North America, Greenland, Western Europe and Australasia. Current and future efforts at isotopic characterization of Proterozoic terranes must concentrate on more poorly known areas where the background information on field relations, tectonics and geochronology is still being gathered. Figure 1 shows an approximate summary of the current state of knowledge of the origins of Proterozoic crust globally. The map has an “age-of-last-orogenicevent” basis, so that the Proterozoic regions actually contain significant Archean crust, and the Phanerozoic regions contain even more Proterozoic/Archean material. The continents are divided into regions last affected by Archean, Proterozoic and Phanerozoic events. The Proterozoic regions are in turn divided into four categories (numbers 1,2 and 3 correspond to Fig. 1): (1) Regions where enough reconnaissance Nd isotopic work has been carried out to approximately characterize the origin and evolution of the whole exposed continental crust. (2) Regions where reconnaissance Nd isotopic work is less complete, either in aerial coverage or in terms of the geologic history or range of rock types of the crustal terranes. This category includes regions where more comprehensive Nd isotopic study is, or may be, in progress. There may also exist smaller regions that are better known within the large domains under this category. (3) Regions where exposed terranes are accessible, but Nd isotopic characterization has not yet been carried out. In some cases this is simply because
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Fig. 1. World map showing approximate level of Nd and other isotopic study to characterize origins of Proterozoic crust. Dark: Archean crust; unshaded with numbers: Proterozoic crust; line shading: Phanerozoic cover or ice of subcontinental scale preventing access to Precambrian rocks; dot pattern: Phanerozoic orogenic belts. Proterozoic terranes are divided into three categories. 1 = Enough reconnaissance Nd isotopic work has been done to approximately characterize the origin of the exposed crust. 2 = Reconnaissance Nd isotopic work is less complete or in progress. 3 = crust is exposed but as yet unstudied for Nd isotopes.
the geochemical approaches have not yet been used, but often there is a lack of sufficient background understanding of field relations and tectonics to enable isotopic approaches to be profitably applied. (4) Regions where Phanerozoic sediments of various ages or ice cover prevent access to the main crust-forming Proterozoic terranes. Antarctica, not on the map, belongs to this category. This covered terrain is line-shaded in Fig. 1. It must be emphasized that this is a very simplified summary of the state of knowledge as of mid-1991. The reference list shows a wide variety of papers which give the basic Nd or other isotopic data. The cited papers contain reference to both earlier isotopic endeavors, and to geologic/tectonic contributions on which the work is based. The reader will note that no regions are described as well known in terms of isotopic characterization. The best that can be said of the better-characterized regions is that the make up of exposed crust is reasonably well understood on the basis of widespread reconnaissance studies, and is unlikely to be severely modified by additional work. Even the better-known regions are subject to the uncertainty that the middle and lower crust has hardly been sampled.
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Evident from Fig. 1is the greater development of geologic, tectonic and isotopic research on North American, European and Australasian terranes, mentioned earlier. Thus there are major terranes available to be studied in the Proterozoic of South America, Africa and Asia. On the other hand, Antarctica, much of Greenland, parts of the large river basins of South America, large regions of the African-Arabian deserts and wide areas of Siberia will probably never be accessible to tectonic and isotopic study. In some regions, such as the midcontinent U.S.A. and the European parts of the former U.S.S.R., the problem of later cover is alleviated by relatively abundant boreholes penetrating to Proterozoic basement. These are a rather poor substitute for exposed terranes however, as field and tectonic relations between lithologic types are hard to determine.
Crustal growth curves Where Nd isotopic characterization has been carried out on a reconnaissance level over major continental areas, it is possible to draw cumulative crustal growth curves. These have been approximated for Australia (McCulloch, 1987) and North America-Europe (Nelson and DePaolo, 1985; Patchett and Arndt, 1986; Condie, 1990). Most of these curves are shown in Fig. 2. It is to be noted that the curves describe the age distribution of presently existing continental crust, and do not allow for any crust that may have been removed to the mantle by any means over geologic time. The differences between the three attempts to describe the North American or North America-Europe age distribution are a measure of the uncertainties in constructing such curves when many regions are poorly known. Significant differences between the Australian and North American-European curves are (1) more abundant pre-3.0 Ga crust in Australia, (2) a more rapid increase of crustal mass 3.0-2.6 Ga in North America-Europe, (3) slightly greater Proterozoic growth in Australia and (4) greater late Proterozoic and Phanerozoic
CONTINENTAL CRUST GROWTH
c
6 v)
1.0
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Age ( Ma 1 Fig. 2. Some published cumulative growth cuwes for regions of continental crust. Slightly different assumptions were made to draw each curve. Note that the SW U.S.A. results are based on a much smaller region than the other curves.
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growth in North America-Europe. Comparison at any greater level of detail is not possible because of assumptions made in drawing the curves. McCulloch (1987) used the older protolith model to explain non-DM initial Nd isotopic ratios, whereas Patchett and Arndt (1986), for example, used a mixing model. These assumptions are explained later in the chapter. Also shown in Fig. 2 is a curve for the southwestern quarter of the U.S.A., a region better known for Nd crustal age than most others (DePaolo et al., 1991). It is a region dominated by crust of mid-Proterozoic age. The major difference between this curve and all the others is a function of the dominant Proterozoic crust in the region. As such, it is clear that the region is geographically too small to be a representative continental mass. This could be true of Australia also, but in any case the different curves for large geographic regions show clearly that curves for growth of the entire continental crust cannot be inferred from partial data sets, e.g. the North American one. Insofar as cumulative crustal growth curves are useful in understanding and modelling differentiation of the Earth, they must be compiled from hard data derived from all available continental regions. Isotopic system stability and reliability of initial Nd isotopic parameters
The Sm-Nd systematics of rocks from Proterozoic and other orogens have been used at a range of levels of detail from the simple identification of rocks of Archean origin in Proterozoic belts to comparison of initial ENd values that differ by only one or two units. To the present reviewer it seems worthwhile drawing a distinction between the use of Nd isotopes to assess large differences in crustal residence time (say >300 Ma), and their application to much more subtle differences in model age or initial ENd value. Later sections will show how the interpretation of Nd model ages, and differences between model ages, are subject to numerous uncertainties. The effect of these uncertainties is to render detailed interpretation of results, and detailed comparison between results, difficult, while leaving the gross distinctions, such as Archean vs. Proterozoic vs. Phanerozoic crust, quite robust. However, another question must be answered before any Sm-Nd results on old rock samples can be interpreted, which is: were the isotope systematics a closed system from the time of interest to the present day? This would seem to be a trivial question in the case of totally pristine igneous rocks, but few old rocks are totally pristine igneous mineral assemblages. There is almost always alteration of original assemblages, most probably involving mobile fluid phases. Additionally, many studied rocks must by definition be from medium to high metamorphic grades, with complete adjustment to new mineral assemblages, achieved at least in part under open-system chemical conditions. Metamorphism of rocks raises the additional possibility that still earlier phases of mineral alteration or metamorphism may be undetectable in the rocks today, meaning that there would be potential chemical disturbances of entirely unknown age. These could mislead the interpretation of Sm-Nd or other isotopic evolutionary results. It is evident that the more detailed the use of Sm-Nd systematics, the more minor are
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the disturbances to rare-earth-element chemistry that have to be considered as potential sources of error. The rare-earth elements (REE) are not very soluble in natural aqueous solutions (e.g. Brookins, 1989 reviews), but they have been observed to have been hydrothermally mobile in a range of instances. Much of the literature concerning mobility or non-mobility of REE was summarized by Grauch (1989). He concluded that authors were about equally divided into those who found mobility and those who found non-mobility. In fact, however, the geochemical environments of all the studies cited by Grauch (1989), and the aims and preconceptions of the authors, vary so much that it is of little value to poll papers for “mobility” or “non-mobility”, as stated already by Grauch (1989). What is evident is that REE can be mobile under certain chemical conditions, but what these conditions are is not yet absolutely clear. Again, it is useful to draw distinction between major REE mobility, causing a significantly changed R E E pattern and grossly disturbed Sm-Nd isotopic systematics, and more subtle disturbances. The major disturbances seem to require very high waterhock ratios whatever the chemical conditions (Brookins, 1989), and thus certainly produce major changes in rock mineralogy and major element composition. These alterations would lead to such samples not being employed for geochemical or isotopic studies directed a t characterizing igneous or sedimentary protoliths of metamorphic rocks. The resilience of the Sm-Nd isotopic system to processes of major hydrothermal alteration in rocks has been investigated to some extent. Farmer and DePaolo (1987) investigated Rb-Sr and Sm-Nd isotopic systematics in a porphyry-copper hydrothermal system, and found very limited susceptibility of Sm-Nd systematics compared to widespread open-system behavior of Rb-Sr systematics. Barovich and Patchett (1991) extended these studies to include Lu/Hf isotope systematics and a range of rock alteration environments. The conclusion is that Rb-Sr is susceptible to disturbance during K-metasomatism, C1-rich fluid alteration and F-bearing fluid alteration. Sm-Nd and Lu-Hf isotopic parameters are only seriously affected by the fluorine-bearing fluids, resulting in gross removal of 75% of the REE and Hf from the rocks. While these studies are important in documenting what kinds of fluids cause chemical and isotopic disturbance, they are of limited direct applicability to petrogenetic studies of old rocks simply because such severely altered rock systems are normally avoided in sampling. It would be much more important, for example, in the light of the alteration studies, to know the concentration of fluorine in a whole range of metamorphic fluids, up to granulite-facies conditions. This type of study of metamorphic fluid compositions is only in its infancy, however. Thus it is difficult to assess the susceptibility of Sm-Nd systematics to metamorphism in a generally meaningful way. One type of chemical mobility that seems very clear is susceptibility of R E E to transport by carbonate-rich fluids a t temperatures of a few hundred degrees. In the best two published examples (Hynes, 1980; Tourpin et al., 1991), alteration occurred under conditions of greenschist-facies metamorphism. Hynes (1980) documented Ti, Y and Zr mobility in mafic rocks, while Tourpin et al. (1991)
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showed large changes of REE abundance patterns in a single komatiite flow, coupled with Nd and Sr isotopic disturbance. The type of alteration which results in rock matrix carbonate or carbonate veins is very common in old mafic and ultramafic rocks. These rocks have also figured prominently in defining Archean and Proterozoic mantle Nd isotopic values, so that considerable caution should be applied to those values. Another approach to isotopic disturbance under metamorphic conditions that has been employed is study of the Sr, Nd and Hf isotopes in a progressively mylonitized granite (Barovich and Patchett, 1992). This was a severe deformation of an initially reasonably homogeneous granitoid, but one that involved only limited fluids. The conclusion is that Sm-Nd and Lu-Hf are extremely resilient to the mylonitization. Clearly, this study does not apply to long-lived metamorphic events where separation of the major and minor elements into gneissic bands may occur. Here the question is what is the scale of element movement, and what is its relationship (larger or smaller) to the scale of the collected rock samples? This and related issues have been addressed by Bridgwater et al. (1989) and Rosing (1990). Rosing (1990) discussed secondary chemical disturbances of Sm/Nd isotopic systematics in a semi-quantitative way. Actually, however, the effects of fluids on REE abundances and hence Sm/Nd ratios are poorly constrained, so that quantitative treatment is not possible. The chief importance of the papers by Bridgwater et al. (1989) and Rosing (1990) is to point out the fact that relatively minor changes of Sm/Nd ratio (
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to positive ENd in 3.8-2.6 Ga Archean rock units (e.g. Shirey and Hanson, 1986) and the modelling of crust-mantle evolution based upon it (Chase and Patchett, 1988; Galer and Goldstein, 1991) are suspect. Probably in this case so many rock units, mafic and felsic, are involved that the general positive ENd must be real. In conclusion then, Sm/Nd (and Lu/Hf) isotope systematics are almost certainly robust enough that gross distinctions of crustal age are possible. At the level of detail where the significance of deviations of 1 or 2 €-units from model growth curves are interpreted from single samples, considerable caution must be exercised. This is particularly the case where the rocks have a complicated, multi-stage metamorphic history spread over several hundred million years or more. One or two epsilon units difference in initial Nd isotopic ratio translates to approximately 100-250 Ma difference in Nd model age of any type; therefore clearly, the uncertainty applies to model ages as well.
Meaning of Nd model ages In addition to the questions of closed-system isotopic evolution discussed in the preceding section, there are additional substantial ambiguities in interpretation of Sm-Nd systematics. These result not from chemical disturbances and consequent failure of the Sm-Nd systematics to yield true information, but from mixing of materials and/or fractionation of Sm from Nd that may have taken place at the time of rock formation. These processes include, for example, mixing of sediments of different provenance at the time of deposition of a sediment, or fractionation of the Sm/Nd ratio during melting of lower crust to yield a granitoid magma. The types of Nd model ages currently used are illustrated in Fig. 3. The TDM of DePaolo (1981) and TCR of Goldstein et al. (1984) are essentially equivalent. They are widely taken to describe the crustal residence age of a rock sample relative to the most plausible isotope chemistry of the upper mantle. This is a mantle depleted by extraction of large-ion lithophile elements (LILE) over geologic time. As a result it is depleted in LREE compared to HREE, hence in Nd compared to Sm, and due to its long-standing depletion, has also a higher 143Nd/144Ndthan that of undifferentiated mantle (Fig. 3). The existence of the depleted mantle is not an issue, and its being the main chemical type of mantle that is involved in oceanic volcanism, island-arc volcanism and orogenic processes is documented repeatedly in the global geochemical data for recent and older ~ ages (McCulloch and rocks. These model ages are distinct from T C H Umodel Wasserburg, 1978), which describe the model age of a crustal rock relative to an undifferentiated mantle reservoir. Such primitive reservoirs do not appear to have taken part in orogenic processes in the Proterozoic and Phanerozoic, although this is by no means certain. It also appears that the Archean upper mantle was at least partly differentiated, so that LILE-depleted sources were common (Shirey and Hanson, 1986; Chase and Patchett, 1988). Consequently, while the CHUR evolution (Fig. 3) is certainly useful as a planetary reference, it is clear that for
Isotopic studies of Proterozoic crustal growth and evolution +I0 t5
‘Nd
Nd ISOTOPE SYSTEMATICS . ’1 I
*
crysfallization age 1.1 Ga ENd = -1.2
/
, .Y
1
0
489
-5 -10
-15
2.0
/1.5 TDM
/
1.0
0.5
0
TCHUR AGE Ga
Fig. 3. Initial 6Nd values and Nd model ages. For samples undisturbed in Sm/Nd systematics, the present day Nd isotopic composition and Sm/Nd ratio are used to calculate a rigorous initial €Nd value, if the age of crystallization is independently known. Model ages for separation from an undifferentiated mantle (TCHUR) or an upper mantle that was already depleted in LILE elements ( T D Mdepend ) upon assuming that some material or combination of materials existed before the dated geologic event that had the same Sm/Nd as the studied rock. This may be true in some cases, but mostly is not true; text and Fig. 4 explain the reasons for this.
major volumes of rocks it may have been a geologically probable source for only the very earliest rocks in the Archean. As such, its usefulness is certainly limited. The limitations on interpretation of Nd model ages have been discussed before, notably by Arndt and Goldstein (1987) and Patchett (1989). The reason why cautionary statements need to be repeated in chapters like this one is that there continue to be misconceptions and oversimplifications regarding the use of Nd model ages. In this context, it is important to distinguish very clearly two different approaches to the use of Sm-Nd results from old rocks. It must be noted that both approaches assume that the Sm-Nd systematics have been undisturbed since the last major geologic event, such as magmatism, metamorphism or sedimentation/ diagenesis. Thus the samples employed must not have suffered any chemical disturbances sufficient to affect Sm-Nd systematics, and collected rock mass must be adequate to average out, for instance, any minor mineral-scale isotopic disturbance. There must be a degree of geological and geochronological control existing on sampled units that is at least appropriate to the goals of the isotopic study. Processing of the rock sample into rock powder must also be such as to guarantee a powder representative of the collected rock material. If one or more of these conditions are not fulfilled, then the isotopic results will be suspect at a much more fundamental level than that being discussed in this section. The two approaches to Sm-Nd studies on old rocks are: (1) To use samples with well-determined ages for the latest major geologic event, calculate initial ENd at that time, and then discuss the interpretation of that initial Nd value. (2) To extrapolate the Nd isotopic evolution back through the latest and any
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other major geologic events to the intersection with DM or CHUR curves to produce a model age, which is then the basis for interpretation. The first approach involves no assumptions about the isotopic evolution except that there has not been isotopic disturbance since the dated major geologic event. The age of that event, however, must be’known to enable calculation of an initial Nd isotopic value. In the Proterozoic and Archean, where no fossils control sediment age, the ages of all rocks must be determined isotopically. Ages of metamorphism can sometimes be determined reliably by Sm-Nd, Rb-Sr or Ar dating. Igneous events, however, are usually obtained by U-Pb zircon geochronology. Many of the more reliably dated times of metamorphism and sedimentation are also due to age brackets derived from pre-dating and postdating igneous rocks. The interpretation of initial ENd values is only simple for values that lie close to the DM evolution at the time of magmatism (or other event). This occurs in the case of “juvenile” mantle-derived igneous rocks in the Proterozoic and later. Usually, however, initial 6Nd values lie below the DM curve, and then the interpretation is complex and often equivocal, as explained in the next section. The importance of approach number 1 however is that the main isotopic parameter that is calculated, initial ENd, refers to a time when some dated and reasonably well-characterized geological event occurred. The interpretation comes afterwards. This is not the case at all in approach number 2, where considerable assumption and interpretation are built into the Nd model age that is calculated. Sometimes this may be unavoidable, as in the case of rocks without adequate geochronology of geologic events; for such cases, the Sm-Nd data only give a rough indication of crustal age, and this may indeed have been the only purpose of the analyses. In general, the TDMmodel age only has a simple interpretation as an actual mantle separation age for the case of an igneous rock derived (whether in one partial melting step or several) from the mantle and crustally stabilized over a time period
Isotopic studies of Proterozoic crustal growth and evolution
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TDMof the resulting magma is only an average of the materials which contributed. (b) Sm/Nd fractionation is practically certain between magma and crustal residues. Usually, the magma will acquire lower Sm/Nd than the starting source. There may be fractionation of the magma by phenocryst precipitation also. For change of Sm/Nd at the time of a magmatic event significantly postdating the original crust stabilization (say by 200 Ma), the TDMage is in error. This error becomes larger as the difference between magmatic and TDMage increases.
2. Sediments (a) The source of the sediments may be of two or more ages, meaning that the TDMis only an average of all contributions. (b) Diagenesis might conceivably alter Sm/Nd ratios or Nd isotopic composition (Ohr et al., 1991, and references therein). (c) Pansport and sedimentation might involve unmixing of detrital mineral species and size fractions having different provenances, leading to unrepresentative Sm/Nd ratios and Nd isotopes in individual samples (McLennan et al., 1989). 3. Metamorphic rocks (a) Sm/Nd might be changed in certain metamorphic processes, leading to incorrect model ages, if metamorphism significantly postdated crust stabilization. (b) High-grade metamorphic rocks may have had felsic melt extracted from them, leading to fractionated (usually higher) Sm/Nd in the residue. This leads to incorrect model ages if melt extraction significantly postdated crust stabilization.
In the above scenarios, it must be stressed that the problems are entirely associated with the application of the model age concept. That is, the isotopic systematics are extrapolated back in time beyond the dated major geologic event. Before that time, the evolution is either incorrect because of Sm/Nd fractionation at the dated event, or in the case of a mixed sediment or granitoid source the mixture did not exist before the event, so its “evolution” is meaningless. It must be stressed that an initial ENd value, calculated in the normal way for the time of the latest major geologic event, is still valid for all the problem scenarios listed above. The only exception would be a later metamorphism or metasomatism that remained undetected. Thus, it is always preferable to work on well-characterized and well-dated rock suites (see also Bridgwater e t al., 1989). Initial Nd parameters can then be calculated at the times of known major events in the history of the rocks, and the interpretation of the Nd values discussed accordingly. Nd model ages can be used as a simple way of determining the gross crustal residence age of poorly understood rocks, or listed as a shorthand means of comparing samples in a more detailed study. The shortcomings of the model age approach have a tendency to be forgotten and buried behind the repetition of a suite of numbers, however. Therefore model ages have to be used with caution, and in the opinion of this reviewer, avoided if at all possible.
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Interpretation of ENd values of Proterozoic rocks Even in cases where all the previously described requisites for reliable and meaningfully presented Sm-Nd isotopic data are fulfilled, unique interpretations of the results can often still not be made. The framework of interpretation is shown in Figs. 3 and 4. A simple conclusion is derived for a sample whose initial 6Nd plotted on the DM curve (sample number 1 in Fig. 4): it must have been a material derived from the mantle shortly before its dated stabilization. This interpretation is still subject to uncertainty over the position of the DM curve globally and regionally during the events which produced the studied rock sample. A higher DM evolution for example, would allow some older crustal influence in the genesis of sample number 1 in Fig. 4. To the extent that the mantle must predominate in the origin of sample 1 however, the interpretation of the data in that case remains simple. Another possible simple case concerns older crustal signatures where the initial ENd values correspond closely to the evolution of some geographically close older crustal terrane. In this case it seems reasonable to conclude that the younger rocks were derived from the older (sample number 2 in Fig. 4). This occurred for example in the much younger case of the Paleozoic Acatlan Terrane of Southern Mexico (Yafiez et al., 1991). Here the Paleozoic shales and greywackes had very similar Nd isotopic evolution to granulites of the nearby Grenville-age Oaxaca Rrrane. Even though the presently seen Oaxaca Terrane is insufficient in size to have yielded the very thick Paleozoic sedimentary pile in the Acatlan Terrane, the correspondence of Nd isotopic evolutions is so close that it seems reasonable to infer a direct relationship. The Oaxaca Terrane is logically part of more extensive +lo
0 ‘Nd -10
-20
I
P
INTERPRETATIONS
2.0
1.5
1.o AGE Ga
0.5
0
Fig. 4. The range of results for initial CNd in old rocks and their interpretation. 1 = a crustal rock with an entirely juvenile mantle (DM) origin; 2 = a crustal rock originating entirely by recycling (e.g. melting or erosion) of older continental crust; 3 = a rock with initial CNd between contemporary depleted mantle and available older crust, but derived by reprocessing (e.g. remelting) of materials (“protolith”) produced from DM at an earlier time; 4 = a rock with similar intermediate €Nd produced by mixing of juvenile mantle products and available older continental crust.
Isotopic studies of Proterozoic crustal growth and evolution
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Grenville crust either to North or South (Yafiez e t al., 1991). For all other cases of crustal Nd isotopic signatures, the interpretation of rock origins must be ambiguous. It is particularly in such cases that the TDMmodel age information can be most misleading. The elements of two basically different interpretations have been discussed widely (e.g. Patchett and Bridgwater, 1984; Bennett and DePaolo, 1987; Bowring and Podosek, 1989; Patchett, 1989). The two main approaches are illustrated by sample numbers 3 and 4 of Fig. 4. (1) Older protolith model (sample 3). It is presumed that the dated rocks were derived by remelting of some older crustal protoliths, which have now been removed from the geologic record. (2) Mixing model (sample 4). The reduction of t N d below D M is explained by mixing of variable quantities of older crustal material with negative ENd into mantle-derived magmas. Both the interpretations have strong arguments for their plausibility. Older protoliths, for example, make sense in terms of primitive intra-oceanic island arc terranes accreted to an orogenic zone 200-300 Ma after formation and remelted into granitoids or reprocessed into sediments at that time. Mixing scenarios seem very plausible because the plate tectonic system delivers a minimum of -400 m thickness of pelagic sediments constantly to all orogenic zones. More locally derived sediments of greater volume, such as trench turbidites, or sequences on passive continental margins that become transformed to active margins, only increase the older crustal material available for mixing processes. Both of the models also have an additional ambiguity associated with them. In the case of the older protolith model, it is not clear what Sm/Nd ratio should be used before the time of the dated event. Note that the assumption of the TDMage is that the Sm/Nd in the protolith was the same as that in the rock studied a t the present day, This is usually unlikely to be true. More plausibly, the protolith could be regarded as a more mafic, maybe arc-like primitive crust. Such a material would have higher Sm/Nd than for instance a granitoid produced by melting it, but the amount by which the Sm/Nd would have been higher is usually unknown. Figure 4 shows a plausible protolith with Sm/Nd midway between CHUR and typical upper crust. This would give 147Sm/144Nd= ca. 0.15 for the protolith. The mixing model incorporates the ambiguity that the age of the older crustal component is usually unknown. Because older crust has more negative EN& this means that a contamination of a mantle-derived magma with 5% of Archean crust has the same effect isotopically as a greater percentage of Proterozoic crust. Although remnant zircons might be used to determine the age of crustal components, this information is usually not available. To construct maps of crustal age or cumulative frequency curves, some assumption has to be made. DePaolo et al. (1991) assumed contamination by sediments of average crustal age in this context. To come now to the approaches that have been used in interpreting Sm-Nd results from Proterozoic rocks, it is no surprise that the full range of possibilities has been used, including approaches described above as suspect. Few workers
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I 3 Patchett
would dispute that ENd values lower than -2 argue for an effect from pre-existing continental crust, usually Archean in the case of Proterozoic orogenic belts. Thus Patchett et al. (1987), Claesson (1987), Chauvel et al. (1987) and Barovich et al. (1989) identified either Archean rocks or Proterozoic sediments containing large Archean crustal components in Proterozoic belts. Additionally, several studies were able to document origins of Proterozoic crust that lay entirely in the mantle shortly before the final crustal stabilization. Examples are from 1.9-1.7 Ga crust in North America (DePaolo, 1981; Nelson and DePaolo, 1984; Chauvel et al., 1987), from 0.9-0.7 Ga crust of the ArabianNubian Shield (Bokhari and Kramers, 1981; Duyverman et al., 1982; Claesson et al., 1984) and from 2.1 Ga crust of West Africa (Abouchami et al., 1990). The interpretation of these results is, like the ones with large Archean crustal components, uncontroversial. It is samples and terranes whose 6Nd values lie between -2 and the DM evolution curve that have generated most variety of interpretations. Thus in situations where very wide areas were studied in a preliminary manner, or where ages of rock units were not very well known, a Nd model age approach was used (e.g. Harris et al., 1984; Nelson and DePaolo, 1985; Hegner and Jackson, 1990). This approach is valid in the cited cases, where an estimate of gross crustal residence age was required. The model age approach was also used, however, in much more detailed study of the Grenville Belt in Ontario (Dickin and McNutt, 1989, 1990). In these studies it was more or less assumed that all rocks were derived from DM at some time, and that the Nd model age usually gave the time of that event (Dickin and McNutt, 1990). This is precisely the kind of approach that has been most strongly criticized by Arndt and Goldstein (1987), Patchett (1989), in this paper, and elsewhere in the literature. It is dangerous because even if the isotopic researchers involved are aware of the limitations of the Nd model age treatment, the impression given to non-specialists is that a rigorous age determination is being presented. In fact, the Nd model age is a very assumption-dependent parameter, as detailed in this chapter. The unreliability of this approach is further underlined by the fact that in all other Proterozoic Sm-Nd data, entirely DM-derived rock units account for only about 25% of the database. This is in spite of the fact that studies have definitely been directed preferentially towards regions of apparent juvenile crust, likely to contain rocks with large DM components. One interpretation of 6Nd values lying between -2 and the positive values of the DM curve is that the crustal production was predominantly from mantle less depleted than DM. This would seem to be reasonable given the range of -3 to +14 in ENd of recent oceanic basalts (e.g. White, 1985). Thus Patchett et al. (1981) found EHf values around +3 for 1.9-1.7 Ga crust in Finland, and contrasted those to EHf = +10 for the Colorado 1.9-1.7 Ga crust (this was the same terrane for which DePaolo (1981) found ENd = +3.5 to +4). Patchett et al. (1981) argued that the Finnish crust was derived from a mantle source less depleted than the one for crust of the same age in Colorado. More recently, Scharer (1991) has
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proposed an undifferentiated mantle (i.e. ENd = 0) origin for 1.7 Ga crust in Labrador based on Pb, Nd, Sr isotopes and other geochemical data. The present author revoked the interpretation of Patchett et al. (1981) when Skiold and Cliff (1984) and Patchett and Kuovo (1986) found mafic or arc-like igneous rocks with DM-like Nd signature in 1.9-1.7 Ga crust of Sweden and Finland. DePaolo et al. (1991) have summarized the arguments which make it unlikely that mantle sources like those of present-day ocean island basalts were ever the main participant in Proterozoic or Phanerozoic crustal genesis episodes. Essentially, ocean island sources seem not to represent a large enough or homogeneous enough volume of mantle to be the main participant in a plate-tectonic cycle lasting 100 Ma or so, and involving of necessity a volume of mantle and mafic crust hundreds of times greater than the continental crust ultimately produced. It does not seem impossible that a major upwelling of some undepleted lower mantle could at any time contribute to crustal genesis, perhaps even to the exclusion of other mantle sources in the manner suggested by Scharer (1991). However, it has to be judged a low-probability event when essentially all other synchronously produced Proterozoic terranes show mafic rocks with ENd values of +3 to +6. An ENd value of +4 was found, for example, in ca. 1.8 Ga basaltic rocks of the Ketilidian crustal region of South Greenland (Patchett and Bridgwater, 1984), only slightly older than, and in pre-drift geography very close to, the 1.7 Ga crust studied by Scharer (1991). The discussion between mixing and older protolith models for generating ENd values between -2 and the DM evolution curve has been described above. Models for contaminating Proterozoic juvenile mantle products with small percentages of Archean crustal materials have been applied by Patchett and Bridgwater (1984), Patchett and Kuovo (1986), Huhma (1986, 1987), Patchett et al. (1987), Ohlander et al. (1987), Chauvel et al. (1987), Patchett and Arndt (1986), Bennett and DePaolo (1987), Skiold et al. (1988), Barovich et al. (1989) and others. Patchett and Ruiz (1987, 1989) gave equal weight to both mixing and older protolith models in Grenville crust of Mexico and Texas. The older protolith model was applied to Australian Proterozoic terranes by McCulloch (1987) and to 1.9-1.7 Ga crust of northern Canada by Bowring and Podosek (1989). The reasons for both the mixing and older protolith models being plausible geologically have been stated earlier. If the reader wishes to digest the arguments in detail, then Patchett and Bridgwater (1984) and Bowring and Podosek (1989) present justifications for the mixing and older protolith interpretations respectively. Thus the interpretation in detail of Proterozoic Nd isotope data have spanned the range of available options. It must be stressed that the detailed interpretation of ENd values near zero for a crustal terrane does not affect the main conclusion to be drawn from the results. That conclusion is that the terrane consists mainly of juvenile mantle-derived materials or their partial melts or differentiates, produced at or a few hundred Ma before the time of crustal stabilization of the terrane. This is definitely the most important conclusion of regional Nd isotopic studies. The detailed interpretation of Nd isotopic parameters and the cause of their shift from
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contemporaneous depleted mantle values is important as well, however. It affects concepts of orogenic process and tectonics, as well as regional geologic history on a time scale of 50-300 Ma. Quite a lot of the geological and tectonic impact of Nd and other isotopic studies in the Proterozoic turns out to depend on the detailed interpretation of E values between -2 and +4. It is unfortunate that such a degree of circularity is built into the arguments on this point.
Origins of Proterozoic crusl Origins identified for continental crust of Proterozoic age span the entire range of environments of orogenic processes seen in the recent and present-day Earth. With the identification of true ophiolites in 1.9-1.7 Ga terranes, Finland (Kontinen, 1987) and QuCbec (St-Onge et al., 1988), it is clear that plate tectonic processes at least similar in gross aspects to the present day were in operation in the Proterozoic. The main area for discussion is if and to what extent continental crust production was more rapid in Proterozoic than in Phanerozoic time, a question reviewed in the last section of this chapter. The present section considers some examples of Proterozoic crustal terranes, the environments identified for their production, and compares those to Phanerozoic counterparts. Figure 5 shows some representative examples of well-known Proterozoic and Phanerozoic orogens in terms of their initial ENd values. The data set is not intended to be complete, but to illustrate the range of observations and studies. A point which emerges from Fig. 5 is that terranes with positive 6Nd values are quite common in Proterozoic crust, whereas in the Phanerozoic only the Canadian Cordillera and New Zealand have this characteristic. Terranes with very negative 6Nd occur in the Proterozoic, but become predominant among Phanerozoic orogenic belts. This is a feature of the Nd crustal data base that has been mentioned repeatedly in the literature, although as data accumulates it is becoming clear that the distinction is not as black and white as was formerly believed.
Proterozoic terranes consisting mainly of recycled older continental crust Many of the regions of crust affected by Proterozoic orogeny that have a high proportion of Archean crustal materials consist mainly of basement rocks reactivated by Proterozoic tectonism. It is clearly a semantic problem as to whether these rocks are called Archean or Proterozoic crust on a tectonic map. Also, new investigations progressively unravel the history of such regions better and better. Thus, in the huge region of northern Canada and Greenland considered to be mostly reactivated Archean crust by Patchett and Arndt (1986), belts of juvenile Proterozoic material appear in new databases (Kalsbeek et al., 1987; Hoffman, 1988, 1989; Hegner and Jackson, 1990). The same reviewers (Hoffman, 1988, 1989) are much more straightforward about classifying rocks as Archean when they were simply tectonically reactivated in the Proterozoic without significant igneous or sedimentary processing. This enables better classification and is more realistic tectonically.
497
Isotopic studies of Proterozoic crustal growth and evolution
and PHANEROZOIC
2.5
2.0
1.o AGE (Ga)
1.5
0.5
0
Fig. 5. Fields of initial eNd values for crustal terranes. DM = LILE-depleted mantle evolution; CHUR = undepleted mantle evolution. NZ = New Zealand (older terranes only shown). For Australia, only samples published with known U-Pb zircon ages are shown (McCulloch, 1987). The “1.9-1.7 Ga North Atlantic” field incorporates many terranes, including several in North America. Not shown, however, are examples of North American Proterozoic crust strongly influenced by mixing with Archean materials in the SW U.S.A. (Ball and Farmer, 1991), north-central U.S.A. (Barovich et al., 1989) and westerncentral Canada (Chauvel el al., 1987). All these results would plot in a region similar to the Swedish “Bothnian Basin” data, o r at even more negative €Nd values. In addition to references mentioned in the text, results are also from Hamilton et al. (1980), Vidal et al. (1984), Bernard-Griffiths et al. (1985), Frost and O’Nions (1985), GariCpy et al. (1985), Wilson et al. (1985), Liew and Hofrnann (1988), Norman et a]. (1987), Frost and Coombs (1989) and Samson et al. (1989, 1990, 1991a, b). Related results not specifically shown here, but documenting the same general relationships, were published in McCulloch and Chappell (1982), Ashwal and Wooden (1983), Basu and Pettingill (1983), Fletcher el al. (1983, 1985), Halliday (1984), Pettingill et al. (1984), Liew and McCulloch (1985), Dickin et al. (1990), Marcantonio et al. (1990) and Daly and McLelland (1991).
Proterozoic crustal belts consisting mainly of older materials recycled via sediments or melting, as opposed to simple reactivation, are not all that common. The best example studied, the Damaran Belt of southern Africa, is actually only just Proterozoic in age: 0.75-0.46 Ga (McDermott et al., 1989), thus extending into the Paleozoic. In Fig. 5 it is seen that the ENd data from this Damaran crust are quite typical of many later orogenic belts in being almost entirely negative (Hawkesworth et al., 1981; Hawkesworth and Marlow, 1983; Harris et al., 1987;
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McDermott et al., 1989). Other major crustal volumes of Archean materials recycled into Proterozoic terranes are mainly confined to the margins of Archean crustal masses (e.g. Huhma, 1987; Skiold et al., 1988; Barovich et al., 1989; Ball and Farmer, 1991). One exception to this is a prominent belt of sediments and partly derivative granitoids lying between two arc-like 1.88 Ga terranes in the Swedish Proterozoic. This is the “Bothnian Basin” of Fig. 5 (Patchett et al., 1987; Claesson, 1987). The abundance of terranes in the Phanerozoic that are dominantly metasediments originally derived from erosion of older crust means that the Proterozoic examples need no special analogy. In fact, the sedimentary sequences resting on visible Archean basement are among the best understood elements geologically of Proterozoic provinces.
Proterozoic terranes consisting mainly of subduction-related igneous rocks or their derivatives Xrranes consisting mainly of subduction-related igneous rocks or their derivatives apparently constitute the most abundant type of Proterozoic terrane. In Fig. 5 are shown several examples in terms of Nd isotopic data. The figure shows a significant part, but not all, of the Nd isotopic data available. Note that the “1.9-1.7 Ga terrane” field includes data from the U.S.A., Canada, Greenland and Europe, and thus represents a large number of individual terranes. The subduction-related tectonic setting of these complexes has often been inferred only from general geological associations. There are however, in addition, many chemical studies that document a chemistry of volcanic and/or intrusive rocks that is like present/day island arcs or continental-margin arcs (e.g. Garrison, 1981; Kahkonen, 1987; Kerr, 1989; Sims et al., 1989). While it is possible that in some regions, somewhat different tectonic environments have been subsumed under the general designation “~ubduction-related”,enough chemical data have been collected on the terranes that such misidentifications cannot be a general problem. The subduction-related volcanism concept is particularly relevant to most of the examples shown in Fig. 5, where large amounts of volcanic rocks, typically calcalkaline and many explosive in eruption style, were produced over comparatively short time intervals, with positive ENd. They are accompanied by derivative granitoids (and sediments in some cases). There has been discussion of the arc crust production rates exemplified by these terranes (Reymer and Schubert, 1986; Patchett and Amdt, 1986; Pallister et al., 1990). Essentially, even in geologic models that allow maximum crust production while remaining realistic in a platetectonic context, the subduction-related complexes can only be produced in the time available if multiple subduction zones are involved. This is so both at the continent scale and on the scale of individual volcano-plutonic belts (Patchett and Amdt, 1986). Arc complexes formed individually in intra-oceanic or near intraoceanic settings would be accreted to the margin of the continent, with attendant metamorphism, remelting and upper/lower crustal differentiation. Very noticeable from Fig. 5 is that most Phanerozoic orogens have negative ENd, and are quite different from these abundant Proterozoic terranes. Whatever
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their tectonic parallels may or may not be, they thus have a totally different (recycled crustal) origin, and constitute a poor model for the juvenile Proterozoic terranes. Patchett and Samson (1990) and Samson and Patchett (1991) show how the Canadian Cordillera provides a reasonable modern analog of Proterozoic crustal additions, geochemically and geologically. Important is that the Canadian Cordillera appears to have had accreted to it arc-like terranes with quite long Phanerozoic histories. These terranes were apparently swept up from a significant part of the paleo-Pacific basin by the plate tectonic system. This is precisely the kind of accretion of arc complexes produced elsewhere that appears to be required in the Proterozoic. Proterozoic terranes consisting of oceanic plateaux and their derivatives? Following the lead of Campbell e t al. (1989), discussing mantle plumes as a source of Archean basalt-komatiite sequences, several authors have developed mantle plumes as sources of potential new crust in Proterozoic and Phanerozoic times. In order to be a source of accreted juvenile material, a plume would have first yielded abundant mafic magmas forming an oceanic plateau. On accretion of the plateau to the continent, the usual melting and consequent upper/lower crustal differentiation would take place, possibly with loss of mafic lower crust to the mantle. The appeal of the plume model lies in (1) the large amount of basaltic crust that can be made over a short time, and (2) that the production of this crust is essentially independent of the plate tectonic cycle. Thus in the Phanerozoic, Richards et al. (1989) modeled continental flood basalts as mantle plume products. While this was not a new idea in itself, the new aspect of the problem became evident when Storey et al. (1991) modeled much of the Caribbean crust as the product of the Galapagos mantle plume. Also Richards et al. (1991) considered the thick Biassic basalts of the Wrangellia Terrane of the Canadian Cordillera to be plume products. The fact that the Triassic basalts are both preceded in the Devonian and postdated in the Jurassic by arc volcanic rocks (e.g. Jones et al., 1977) does not affect the plume interpretation, because the plume-related basalts are essentially imposed from below the near-surface tectonic rkgime. In the Proterozoic, an oceanic plateau model has been employed to explain the extensive 2.1 Ga crust of West Africa (Abouchami et al., 1990; Boher et a]., 1992). Here the mafic volcanic rocks which predominate appear to have virtually no older crustal contribution, and have a very restricted age range of less than 50 Ma. The ENd values are shown in summary form in Fig. 5, and Fig. 1 gives an idea of the geographical extent of this crust in West Africa. During the process of accreting the mafic oceanic plateau to the continent, arc volcanism took place and extensive granitoids were intruded. These granitoids also have the Nd isotopic signature of juvenile crust (Boher et a]., 1992). These examples, particularly the Proterozoic one, underline the potential ability of plume-related volcanism to generate large masses of juvenile, more or less basaltic materials in short time periods. If the plate-tectonic regime locally favored
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development of arcs on such plateaux, or they were internally differentiated during collisions, then they could become difficult to subduct and thus make large contributions to crustal growth at any time or place in Earth history.
More rapid crustal genesis in the Proterozoic? A much discussed question in studies of crust generation is to what extent there is hard evidence for faster global rates of crust production at various times in the Precambrian. A related phenomenon is clearly the heat production budget of the Earth (Fig. 6), whereby heat from radioactive decay was around double the present value at the start of the Proterozoic 2.5 Ga ago, declining to ca. 1.1 times the present value 0.6 Ga ago. Given mantle convection, it seems reasonable that part, maybe most, of the higher heat production in Precambrian time was reflected in higher heat loss at the surface by whatever means. Most workers have assumed that in a mantle convective regime, higher heat loss would result in higher mafic or ultramafic magma production (e.g. Chase and Patchett, 1988). Less clear is in what manner and to what extent the heat production curve is linked to production of continental crust. More mafic crust production presumably must mean total longer oceanic ridge length globally and/or faster mafic crust production rates for a given ridge. The other part of the background to the discussion is that there is evidence for rapid crustal growth in the Proterozoic, whether on a continent scale or on the scale of individual volcano-plutonic belts (e.g. Nelson and DePaolo, 1985; Patchett and Arndt, 1986; Dixon and Golombek, 1988; Pallister et al., 1988, 1990; Abouchami et al., 1990).
I
1 HEAT PRODUCTION OF EARTH
10 -
1oi3w
5-
0
1
2
3
4
AGE Ga
Fig. 6. Heat production of the Earth through time, from Christensen (in James, 1989). This model is based on 21 ppb U in the bulk Earth, chondritic Th/U and K/U = 10,000. The important point for the discussion is the nearly 2 x greater heat production at 2.5 Ga compared to the present day.
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However, the linkage between arc volcanism and continental crust production on the one hand and heat loss/mafic crust cycling on the other is by no means clear. The loss of heat a t subduction zones is a trivial part of the heat budget, and the thermal requirements are met even if no arc volcanism takes place. Melting a t subduction zones, with its high potential for adding to the continental crust, takes place only because of the subduction of wet rocks and the retention of a t least some water down to the 100-150 km depths of magma genesis (e.g. Campbell and ’Eiylor, 1983). Here the amounts of subducted wet sediment, degree of hydration of oceanic crust, and particularly the angle of subduction play roles. Arc volcanism and continent generation by subduction-related magmatism are not necessary to the Earth’s heat budget, and are thus a by-product of the Earth’s main processes. Another very open question in the arc magmatism method of crust generation is how stable subduction zones were with respect to the existing plate framework. This is important because in the recent past, most arc complexes up to 50 Ma old consist almost entirely of basalt and hence are certainly still subductable on density grounds. Mature arc and continent production may require subduction to persist at one site for longer time periods. Another connection that is very loose is that between orogeny and juvenile crust creation. Faster plate motions and more collisions do not necessarily mean more crust creation. As in the Phanerozoic, many orogenic belts could consist mainly of recycled older crustal materials. Thus the connection between heat budget and continent growth is tenuous, for all the reasons stated above. Workers who have considered the question of crust production rates have mostly recognized a need for a t least locally increased growth (Nelson and DePaolo, 1985; Patchett and Arndt, 1986; Dixon and Golombek, 1988; Pallister et al., 1990). Condie (1990) recalculated crust production rates for Laurentia and found a declining rate from Archean to present, consistent in general with the Earth’s heat production curve. This might imply that in spite of all the uncertainties listed here, and the tenuous nature of the link between heat and crust differentiation, the two are in fact quite well correlated. Of other workers, some have argued for concentrating of arc-like terranes by the plate-tectonic system as a means of creating apparent very rapid growth episodes (e.g. Patchett and Arndt, 1986; Pallister et a]., 1988; Samson and Patchett, 1991). Others have regarded subduction-related volcanism as inadequate to produce the rapid growth seen at certain times and in certain regions, and have implied that other means of crust production must have operated (e.g. Reymer and Schubert, 1986). This is clearly the point where the mantle plume-oceanic plateau model may provide at least a partial solution. This model is totally different to the arc magmatism model, not least because the plume contribution can be very large locally and is not related to the near-surface plate tectonic regime. The plume is a Straightforward product of deep mantle heat and as such the frequency of plumes should be related in a more straightforward manner than arc volcanism to the heat production curve. The main question is how often basaltic oceanic plateaux can escape subduction, become accreted and internally differentiated. The West Africa
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Proterozoic crust is so extensive (Boher et al., 1992), that if the oceanic plateau model is the correct mode of origin, then this model is already an important Proterozoic crust generation mechanism. Four or five crustal regions the size of the West Africa one over geologic time would make oceanic plateau production a major source of new crust generally. This would then make discussion of the adequacy of arc production rates to generate Precambrian crust (e.g. Reymer and Schubert, 1986; Patchett and Arndt, 1986) unnecessary. The above questions are still open because of theoretical considerations partly. However, a more important reason is that much basic data on Proterozoic crust origins remains to be gathered (Fig. 1). Even the global extent of juvenile Proterozoic crust cannot be more than crudely estimated with the current state of knowledge. Much work remains for future studies.
ACKNOWLEDGEMENTS
K. Geiger, I. Bambach and G. Feyerherd helped prepare this chapter. Reviews by N.T. Arndt, D. Bridgwater, K. Condie and L. Farmer were useful in revising the paper. The financial support of the Alexander-von-Humboldt Foundation is acknowledged. Researches of the author or close colleagues on which this review is based were supported by the Max-Planck-Gesellschaft, the Deutsche Forschungsgemeinschaft and through National Science Foundation grants EAR8615844, EAR-8616473, EAR-8903764, EAR-8609364, EAR-8804552 and EAR9017330.
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Richards, M.A., Duncan, R.A. and Courtillot, VE., 1989. Flood basalts and hot-spot tracks: Plume heads and tails. Science, 246: 103-107. Richards, M.A., Jones, D.L., Duncan, R.A. and DePaolo, D.J., 1991. A mantle plume initiation model for the Wrangellia flood basalt and other oceanic plateaus. Science, 254: 243-267. Rosing, M.T., 1990. The theoretical effect of metasomatism on Sm-Nd isotopic systems. Geochim. Cosmochim. Acta, 54: 1337-1341. Samson, S.D. and Patchett, P.J., 1991. The Canadian Cordillera as a modern analogue of Proterozoic crustal growth. Aust. J. Earth Sci., 38: 595-611. Samson, S.D., McClelland, W.C., Patchett, P.J., Gehrels, G.E. and Anderson, R.G., 1989. Evidence from neodymium isotopes for mantle contributions to Phanerozoic crustal genesis in the Canadian Cordillera. Nature, 37: 705-709. Samson, S.D., Patchett, P.J., Gehrels, G.E. and Anderson, R.G., 1990. Nd and Sr isotopic characterization of the Wrangellia terrane and implications for crustal growth of the Canadian Cordillera. J. Geol., 9 8 749-762. Samson, S.D., Patchett, P.J., McClelland, W.C. and Gehrels, G.E., 1991a. Nd and Sr isotopic constraints on the petrogenesis of the west side of the northern Coast Mountains batholith, Alaskan and Canadian Cordillera. Can. J. Earth Sci., 28: 939-946. Samson, S.D., Patchett, P.J., McClelland, W.C. and Gehrels, G.E., 1991b. Nd isotopic characterization of metamorphic rocks in the Coast Mountains, Alaskan and Canadian Cordillera: ancient crust bounded by juvenile terranes. Tectonics, 10: 770-780. Scharer, U., 1991. Rapid continental crust formation at 1.7 Ga from a reservoir with chondritic isotope signatures, eastern Labrador. Earth Planet. Sci. Lett., 1 0 2 110-133. Shirey, S.B. and Hanson, G.N., 1986. Mantle heterogeneity and crustal recycling in Archean granitegreenstone belts: evidence from Nd isotopes and trace elements in the Rainy Lake area, Superior Province, Ontario, Canada. Geochim. Cosmochim. Acta, 50: 2631-2651. Sims, P.K., Van Schmus, W.R., Schulz, K.J. and Peterman, Z.E., 1989. Tectonostratigraphic evolution of the Early Proterozoic Wisconsin magmatic terranes of the Penokean Orogen. Can. J. Earth Sci., 26: 2145-2158. Skiold, ' I and Cliff, R.A., 1984. Sm-Nd and U-Pb dating of early Proterozoic mafic-felsic volcanism in northernmost Sweden. Precambrian Res., 26: 1-13. Skiold, T., Ohlander, B., Vocke, R.D. and Hamilton, P.J., 1988. Chemistry of Proterozoic orogenic processes at a continental margin in northern Sweden. Chem. Geol., 69: 193-207. St-Onge, M.R., Lucas, S.B., Scott, D.J., Begin, N.J., Helmstaedt, H. and Carmichael, D.M., 1988. Thinskinned imbrication and subsequent thick-skinned folding of rift-fill, transitional crust and ophiolite suites in the 1.9 Ga Cape Smith Belt, northern Quebec. Current Research, Geol. Surv. Can., Pap., 88-1C: 1-18. Storey, M., Mahoney, J.J., Kroenke, L.W. and Saunders, A.D., 1991. Are oceanic plateaus sites of komatiite formation? Geology, 19: 376-379. Tourpin, S., Gruau, G., Blais, S. and Fourcade, S., 1991. Resetting of REE and Nd and Sr isotopes during carbonitization of a komatiite flow from Finland. Chem. Geol., 90: 15-29. Vidal, P., Bernard-Grifiths, J., Cocherie, A., LeFort, P., Peucat, J.J. and Sheppard, S.M.E, 1984. Geochemical comparison between Himalayan and Mercynian leucogranites. Phys. Earth Planet. Inter., 35: 179-190. White, W.M., 1985. Sources of oceanic basalts: radiogenic isotopic evidence. Geology, 13: 115-118. Wilson, M.R., Hamilton, P.J., Fallick, A.E., Aftalion, M. and Michard, A., 1985. Granites and early Proterozoic crustal evolution in Sweden: evidence from Sm-Nd, U-Pb and 0 isotope systematics. Earth Planet. Sci. Lett., 72: 376-388.
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509
REFERENCES INDEX
Abbott, D.H., 168,174 Aberg, G., 291 Abouchami, W., 36, 46, 435, 438, 440, 494, 499, 500,502, 503 Ackermand, D., 352, 353,442 Adams, J.A.S., 369,380 Admou, H., 91, 94 Afifi, A.M., 462, 471 Aftalion, M., 54, 310, 312, 313,347,378,445, 507 Agena, W.E, 137, 140,292 Ahmad, S.N., 397,416 Ahmad, T, 38, 46,46, 160, 161,174 Aitken, J.D., 132-134,136 Ajibade, A.C., 126, 127,136 Akerblom, G.V., 273,299 Al-Shanti, A.M.S, 65, 71, 88, 436, 440 Alapieti, TT, 100,136, 184-187, 208, 209 Albarkde, E , 46,416,440, 502, 503 Albee, A.L., 475 Albin, A.A., 450,471 Aleinikoff, J.N., 450,452,459,460,469,471, 476, 477,479 Alford, D.E., 463, 471 Allaart, J.H., 430, 432, 440 Alkgre, C.J., 504 Allen, M.B., 427,440, 446 Allen, P., 131, 136 Andersen, D.J., 294 Anderson, A.T., 235, 252,255 Anderson, C.A., 454,471 Anderson, J.L., 4, 31, 47, 97, 136, 219, 251, 252, 255,263-266,268,269,271-274,282-286,288290, 291-294, 298, 299, 359, 434, 438, 440, 447, 450,471, 478, 480 Anderson, P., 454, 456, 471 Anderson, R.E., 91 Anderson, R.G., 477, 507 Anderson, R.R., 138 Anderson, TH., 449, 465, 471, 478 Anderton, R., 131,136 AndrCasson, I., 147 AndrCasson, P.-G., 131, 132, 136, 147 Andreopoulos-Renaud, U., 89 Andrews, A.J., 100,139
Andriambololona, R., 379 Andriessen, P.A.M., 307, 311, 313,347 Anhaeusser, C.R.,51,144 Anonymous, 55, 83,85,88 Anovitz, L.M., 303, 306, 309, 319, 323, 325, 327, 333,334,344,346,347, 354 Aoki, K.-I., 366, 377 Arculus, R.J., 10, 47, 52, 205, 209, 364, 365, 367, 377-380 Arima, M., 305, 326, 347 Armbruster, Th.,336, 339,347 Armstrong, D.G., 471 Armstrong, R.A., 87,88, 197,209, 213,417 Armstrong, R.L., 134,136, 293, 455, 474 Amdt, N.T,7 , 11, 26-28, 47, 51, 89, 102, 103, 136, 206, 209, 293, 361, 312, 375, 376, 377, 380, 417, 427, 444, 484, 485, 489, 494-496, 498,500-502, S02, 503, 506 Arndt, P.J., 291, 297 Aro, K., 50,442 Asami, M., 356 Ashwal, L.D., 215, 217, 229, 245, 246, 254, 255, 259,265, 267,278,291,292,299,497,502 Aspen, P., 378, 381 Atkin, B.P., 122, 123,138 BABEL Working Group, 429, 440 Badger, R.L., 128,136 Baer, A.J., 7 , 47, 56, 75, 88, 443 Bailey, D.G., 103, 104,148 Bailey, J.C., 8, 9, 18, 29, 30, 47 Baillie, P.W., 145 Baker, A.J., 336, 338, 347 Baker, B.H., 97,136 Baker, J.H., 429,440 Bakor, A.R., 59,68,88 Baldwin, D.A., 29, 47 Ball, TT., 497, 498, SO2 Ballhaus, C., 305, 347 Banks, P.O., 298 Baragar, W.R.A., 25-27, 47, 75, 76, 86, 88, 103, 137, 139,188,189,191,192,209,211,267,295 Barbey, P., 302, 304, 307, 309, 312, 316, 317, 325, 340,341,343,344,347, 348,421,440
5 10 Barbosa, J.S.E, 312, 347 Barker, E , 264, 267-269, 271, 272,275,277, 279, 281, 284,291,292, 298,449,452, 453,471, 478 Barley, M.E., 417 Barnes, R.G., 147,357 Barnes, S.J., 64, 89, 125, 126,137, 203, 209 Barnicot, A.C., 7 , 4 7 Barovich, K.M., 427, 434,440, 486, 487, 494,495, 497, 498,502 Bartel, A., 297 Barton, C.N., 357 Barton, J.M., 111,112,137,197,209, 210 Bassot, J.P., 36,47 Basu, A., 49 Basu, A.R., 497,502 Bauer, P.W., 463, 468, 471, 477 Baur, M.E., 402, 408,412 Baur, N., 353 Beams, S.D., 293 Beard, L.S., 454,471 Beccaluva, L., 61, 89 Becker, R.H., 408,412 Becker, S.D., 442 Bedard, J., 90 BCdard, J.H., 152,174 Beeson, J., 305,348 BCgin, N.J., 76,89, 94, 95,444,507 Behr, S.H., 213 Behrendt, J.C., 116,117,137, 138, 140,290,292 Behrens, K., 440 Belevtsev, R.Ya., 388, 412 Belevtsev, Ya.N., 388, 412 Bell, R.T., 121, 132, 133, 135, 137 Bence, A.E., 52,176 Bender, E.E., 47, 252, 255, 263, 265, 269, 271274, 282, 283, 285, 286, 288, 290, 292,471 Bender, J.E, 176 Bennett, VC., 84,89,263,272,275,277,292,450, 451, 468,471, 472, 493, 495, 503 Berg, A.H., 189, 190, 209 Berg, J.H., 118, 119,142, 219, 220, 223, 230, 231, 237,253,256,309,348 Bergeron, R., 76,89 Bergh, S.G., 107, 108,137 Berhe, S.M., 65, 89 Bernard-Griffiths, J., 421,440,497,503, 507 Bernasconi, A,, 125, 137 Berner, R.A., 404,412, 416 Berry, R.E, 305,306,347, 355 BerthC, D., 317, 348 Berthelsen, A,, 420, 421,440, 441 Bertrand, J.M.L., 89, 126, 127,137, 138 Beswick, A.E., 8,47
References index Bethune, K.M., 319,348 Betton, P.J., 48 Beukes, N.J., 5, 384, 385, 388, 392-399, 401-409, 411,412, 415-417 Beunk, EE, 148 Bevier, M.L., 77, 83, 91 Bhattacharya, P.K., 356 Bhowmick, S., 356 Bickford, M.E., 51, 95, 219, 256, 263, 265, 269, 271, 273, 277, 286, 290, 290, 292, 293, 297 298, 445, 449, 453, 459-462, 467, 472, 475477, 479, Bickle, M.J., 7 , 47, 56, 57, 77, 80,89, 94, 97, 143, 166, 167,175, 176,336,338,351,416 Billaud, P., 61, 92 Binns, R.A., 309,329,348 Bird, J.M., 128, 137 Birkett, TC., 145 Bishop, D.T, 149 Bjorlykke, K., 131,137 Blacet, P.M., 471 Black, L.P., 51, 154, 159,177, 306, 313, 348, 356, 357 Black, R., 89,126,137, 138 Blackburn, W.H., 33,47 Blais, S., 507 Blake, D.H., 110,137 Blake, TS., 97,137 Blakely, R.J., 478 Blight, D.E, 307,348 Blockley, J.G., 385, 398, 402, 417 Bloomer, S.H., 93 Blount, J.G., 459,472 Blusson, S.L., 414 Boardman, S.J., 449, 453, 459, 462, 472, 477 Bobnow, D.J., 174 Bodinier, J.B., 127,137 Bodinier, J.L., 61, 63,89,379 Boeuf, M.G.A., 90 Bogatikov, O.A., 146 Boher, M., 46,440,499,502,502, 503 Bohlen, S.R., 301, 306, 309, 310, 320, 321, 323, 325, 327-329, 336, 344, 348, 355, 358, 361, 362,376,377 Bokhari, EY., 494,503 Bolivar, S.L., 297 Bolsover, L.R., 253,256 Bolton, W.R., 293 Bond, G.C., 99,127,128,135,137, 139 Bondeson, E., 432,440, 446 Bonnichsen, B., 412 Bonnot-Courtois, C., 396, 397, 412 Borg, G., 123,124,137, 138
References index Bornhorst, TJ., 141 Bose, M.K., 38,47 Boulanger, J., 218, 256 Boullier, A.M., 302, 309, 317,348 Bowes, D.R., 93,347, 351,444 Bowling, G.P., 457, 472, 473 Bowring, S.A., 30, 31, 47, 49, 50, 84, 92, 423, 427, 442, 447-451, 453-458, 463, 464, 467, 467469,471-475, 477,487,493,495,503 Boyd, D.M., 145 Boyd, ER., 171,174 Bradbuty, H.J., 129-131,140 Braddock, W.A., 297, 459,472 Bradley, G.M., 147 Braile, L.W., 141 Brandt, E.L., 478 Brannon, J.C., 117, 119, I38 Breitkopf, J.H., 64, 65,89, 98, 125, 126,138, 144, 411,412 Bressler, 122 Brewer, T.S., 11, 13, 15-21,52, 106, 108, 122,123, 138, 145,429,444 Bricker, O.P., 402, 415 Briden, J.C., 7, 51 Bridgewater, D., Bridgwater, D., 31, 51, 106, 145, 219, 220, 256, 264, 278, 292, 297, 430, 432, 440, 444, 487, 491, 493,495,503, 506 Briegel, J.S., 230, 256 Bristow, J.W., 199, 213 Brock, M.R., 460,472 Brock, TD., 412 Brookins, D.G., 271, 277, 293, 297, 363,379, 486, 503 Brooks, C.K., 167,174 Brophy, J.A., 91 Brown, B.A., 442 Brown, G.E, 65, 87,89 Brown, G.M., 181, 213, 243, 261 Brown, L., 146, 147 Brown, P.E., 349, 353, 358, 434,442 Brown, R.E., 147,357 Brown, S.K., 260 Brown, W.L., 218,256 Brueckner, H.K., 139 Briigmann, G.E., 47,136 Brun, J.P., 317,348 Brunfelt, A.O., 140 Bryant, B., 449, 454,472, 475 Budding, A.J., 271, 293, 463, 473 Buddington, A.F., 239,256, 309, 310, 329,348 Buick, I.S., 324, 327, 328, 336, 339, 351 Buick, R., 418
511 Buisson, G., 61, 61, 89 Bunting, J.A., 414 Burger, A.J., 192, 197,209, 213 Burke, K., 7 , 47, 56, 56, 57, 59, 63, 64,71, 75, 89, 90,97,138 Burns, L.K., 474 Burton, K.W, 307, 311,333,336,343,348 Butler, P., Jr., 412 Button, A., 402, 406, 412 BVSP (Basaltic Volcanism Study Project), 9, 47, 116-118,138 Byerly, G.R., 406,412 Caby, R., 63,64,89,126,127,137, 138 Cadman, A.C., 152, 153,155,174 Caen-Vachette, M., 36,47 Cairns-Smith, A.G., 406, 412 Caldwell, G., 18, 47 Callaham, D.M., 458,468,472,478 4Calvez, J.Y., 92 Cameron, E.M., 408,413 Cameron, K.L., 363, 372, 374, 379, 380 Cameron, M., 372,379 Camp, YE., 65-67, 69, 70,89, 94, 436, 440 Campbell, I.H., 100, 114,138, 143, 169,174, 184, 187,188,208,209, 211,499,503 Campbell, L.M., 501, 503 Cann, J.R., 98, 129,145, 178 Cannon, W.E, 116,117,137, 138,140, 142,292 Capdevila, R., 440 Card, K.D., 147,174 Carlson, J.E., 271, 298 Carlson, K.A., 422, 445 Carlson, R.W., 220, 225, 253,256, 378 Carlson, W D . , 328,353, 472, 479 Carmichael, D.M., 95,507 Carney, J.N., 357 Carswell, D.A., 378 Carter, B., 267, 293 Carter, J.L., 363, 364, 366, 367, 376, 379 Cartwright, I., 7 , 47, 156, 161, 162,174, 333, 339, 340,348 Casati, C., 356 Cavin, WJ., 463, 472 Cawthorn, R.G., 51,144,200,203,209-211 Chacko, T, 307, 309,311,313,322,323,325,327, 348, 350 Chakraborti, M.S., 38, 47 Chakraborti, S., 356 Chamberlain, K.R., 449, 450, 454, 464,472 Chandler, VW., 116,138, 141, 143 Chapman, H.J., 100,138, 154,174 Chapman, N.A., 381
512 Chappell, B.W., 47, 136, 176, 261, 288, 293, 298, 379, 497,505 Chapperll, B.W., 377 Charlou, J.L., 416 Charsley, TJ., 353 Chase, C.G., 117, 138, 333, 334, 344, 346, 347, 488,500,503 Chauvel, C., 28, 47, 75, 89, 136, 291, 293, 494, 495, 497,503 Chauvel, J.J., 398, 409, 413 Check, G., 349 Chengzao, J., 40, 48 ChevC, S.R., 384,413 Chiarenzelli, J., 219, 220, 259, 268, 296, 333, 354 Chivas, A.R., 176 Chorlton, L.B., 87,90 Chou, I-Ming, 402, 413 Choudhari, A,, 53 Christiansen, R.L., 356,444 Christie-Blick, N., 139 Church, W.R., 60, 63, 65, 90 Claessens, W, 142 Claesson, L-A., 18, 53, 429, 445, 506 Claesson, S., 109, 131,138, 494, 498, 503 Clark, M.D., 454,473 Clarke, G.L., 304, 305, 309, 315, 316, 323, 325, 333,334,348, 355, 358 Clarke, H.S., 298, 299, 306, 324, 325, 328, 359, 478, 480 Clarke, J.W, 128,146 Clayton, R.N., 353, 408, 412 Clemens, J.D., 348, 366,377, 381,445 Clendenin, C.W., 141, 197, 210 Cliff, R., 20, 53 Cliff, R.A., 304, 325,348,363,377,507 Clifford, P.M., 495,504 Cloud, P., 402, 406,413 Clubley-Armstrong, A.R., 209 Cocherie, A,, 507 Codding, D.B., 463,474 Coertze, EJ., 197, 209 Cohen, A.S., 178 Cohen, R.S., 365,377 Coish, R.A., 128, 129,138 Cole, J.C., 506 Cole, J.S., 93 Coleman, D.S., 479 Coleman, R.G., 55, 56, 58, 65, 87, 89, 90 Collerson, K.D., 52, 223, 251, 256, 319, 356, 487, 503 Colley, H., 19,48, 429, 440 Collier, J.D., 272, 293 Collins, W.J., 288, 293, 340, 341, 348, 358
References index Compston, W., 88,353, 417,443 Condie, K.C.,1, 6, 7 , 8, 31-33, 45, 46, 48, 50, 52, 56,84,87,90,97-101,111,112,138,139,162, 174, 203, 205, 209, 263, 264, 269, 271, 288, 290, 293, 447, 449, 451-460, 462, 463, 463, 467, 467, 469, 472, 473 475, 475-479, 484, 503, 506 Coney, P.J., 437,441, 447, 470,473, 475 Connelly, TJ., 473 Connolly, J.R., 472 Conradie, J.A., 293 Conte, J.A., 128,144 Convert, J., 440 Conway, C.M., 92,454, 456, 457,473, 475, 478 Cook, FA., 128,138 Cook, P.J., 412 Coolen, J.J.M.M., 302, 307, 310, 325, 335, 336, 347, 349 Coornbs, D.S., 497, 504 Cooper, D.J.W., 444 Cooper, J.A., 436,441 Cooper, J.R., 457, 473 Cooray, P.G., 349 Copeland, P., 449, 457, 458, 464, 473 Corfu, E , 100,139 Corks, J.B., 394, 413 Corrigan, D., 349 Costa, M., 356 Courtillot, VE., 507 Courtney, R.C., 114,139 Coward, M.P., 439,441, 444 Cox, K.G., 15.5, 164,172, 173,174, 175, 206,209 Craddock, C., 193, 209 Crawford, A.J., 155, 174 Crawford, M.B., 430,441 Creaser, R.A., 288, 289,293 Creasey, S.C., 454, 471 Crittenden, M.D., 133,139 Crow, C., 100, 101, 111,112,139, 203, 205,209 Crow, M.J., 357 Cudzillo, TE, 271, 277, 292 Cullers, R.L., 263, 265, 268, 269, 272-274, 283, 284,286-290,292, 293,434,440,476 Culotta, R., 146 Culshaw, N.G., 319,349 Cuney, M., 54 Cuney, M.K., 421,440 Currie, K.L.,261, 298, 305, 319, 323, 326, 328, 331,333,349 Czamanske, G.K., 378 D’Agrella-Filho, M.S., 146 Dall’agnol, R., 264, 268, 293, 296
References index Dallmeyer, R.D., 355,464,465,469,473, 474, 478 Dallwitz, W.B., 349 Daly, J.S., 263, 267, 287, 293, 313, 315, 349, 356, 497,503 Dalziel, I.W.D., 99, 127, 128, 135,139 Damon, P.E., 295, 297 Daniel, C.G., 478 Daniel, C.J., 463, 469, 474 Daniels, L.R.M., 155,175 Danielson, A., 396, 416 Dann, C.J., 33,48 Dann, J.C., 57,58,84,90,439,441, 455, 456,474, 475 Dasgupta, S., 356 Dash, B., 340,347, 351,355 Davidson, A., 140, 302, 303, 306, 318, 319, 322, 323,325,331,349, 356,427,443-445 Davies, G., 200, 201, 203, 209, 210 Davies, G.R., 177, 356, 365, 366, 381 Davies, J.H., 166, 175 Davis, D.W., 116,139, 145 Davis, G.A., 277, 293 Davis, G.L., 220, 258, 364, 365, 377 Dawoud, A. S., 443 Dawson, J.B., 377, 378 Day, W.C., 100,147 Dazhong, S., 39, 40,48 De Beer, J.H., 203, 210 De Laeter, J.R., 41 7 Delaney, G.D., 120,139 Delevaux, M.H., 294 Delhal, J., 50 Delor, C.P., 348 Demaiffe, D., 221, 232, 239, 245, 246, 251, 252, 256, 261,294, 297 DeMalas, J.P., 457, 458,473 Demarchi, G., 364, 381 Dempster, TJ., 304, 317, 325, 347, 349, 442 Dennison, R.E., 297, 459, 464,474 DePaolo, D.J., 84, 89, 151, 175, 220, 256, 263, 275, 277, 278, 287, 288, 291, 292-294, 297, 361, 379, 450, 451, 455, 468, 470, 471, 472, 474, 476, 484-486, 488, 493-495, 500, 501, 503, 504, 506, 507 Derrick, G.M., 54 Derry, L.A., 395,396, 413 Desborough, G.A., 277,292 Devlin, W.J., 133, 134, 139 De Waard, D., 216, 223,256,310, 329,349 Dewey, J.F., 7 , 47, 48, 55-57, 63, 64, 75, 85, 89, 90, 97, 128,137, 138, 433,441 De Wit, M.J., 86, 87, 88, 90 DeWitt, E., 265, 449, 450, 454, 474
513 Dia, A., 36, 48 Dick, H.J.B., 51,143 Dickas, A.B., 117,138, 139 Dickin, A.P., 53,378, 494, 497, 504, 505 Diecchio, R.J., 128,140 Dimroth, E., 103,139,385,398, 409,413 Ding, P., 426,441 Dixon, TH., 6 5 , 9 4 4 4 1 , 500, 501,504 D’Lemos, R.S., 441 Doe, B.R., 273,294, 297, 479 Doe, M.F., 475 Doherty, W., 143 Doig, R., 76, 90 Dollase, W.A., 348 Dorr, J. Van N., Jr., 401, 409, 413 Dostal, J., 89, 121, 137, 139, 205, 210, 368, 377, 379 Downes, H., 363, 375,377 Downes, J., 340, 341,355 Drage, J., 349 Drake, A.A., 146 Drever, J.I., 402, 403, 413 Dreyer, C.J.B., 388,412 Duchesne, J.C., 215-217,219-221, 229-231,236, 237,248, 251,256, 261,264, 294 Ducrot, J., 61,90 Duebendorfer, E.M., 452,464,474 Duncan, A.R., 10,36,48,98,139, 205,210 Duncan, R.A., 507 Dunlop, J.S.R., 418 Dunning, G.R., 87, YO Dupuy, C.,89,137, 139, 205,210,377, 379 Duyverman, H.J., 494,504 Dymek, R.E, 395, 397,413, 414 Dymek, R.G., 227, 234, 237, 238,257 Dymond, J., 413 Easton, R.M., 30, 48, 109, 110,139, 422, 441 Eckstrand, O.R., 188,212 Edwards, B.R., 268,294 Edwards, D.L., 472, 479 Edwards, R.L., 306,309, 310, 329, 336,349 Eggler, D.H., 268, 269, 294 Eglington, B.M., 203, 210, 264, 294 Ehlers, C., 19, 48, 429, 441 Ehrenberg, S.N., 364,377 Eishacher, G.H., 132,136, 139 Ekstrom, H., 267, 294 Elderfield, H., 394-396, 413 Ellam,R.M., 155,164,172,175, 176 Elliott, G., 479 Ellis, D.E., 324, 349
5 14 Ellis, D.J., 171, 175, 301, 305, 306, 309, 321, 322, 324,326,326-328,343,346,349, 356, 359 Ellis, M., 141 Elston, D.P., 122, 133, 139 Elthon, D., 254, 257 Elvsborg, A., 137 Embree, G.E, 379 Emerman, S.H., 97,148, 152,175 Emery, J.V., 367,378 Emslie, R.E, 215-217, 219, 220, 228, 230, 231, 234-238, 244, 245, 248, 251-254, 255, 257, 259, 264, 265, 267, 268, 290, 292, 294, 299, 343,349 Engel, A.E.J., 67, 87, 90, 436,441 England, P.C., 329, 334, 344, 349, 356, 444 England, R.N., 349 Erdmer, E, 305,350, 355 Erickson, R.C., 271, 277, 294, 457, 468, 474 Eriksson, K.A., 148, 463,475, 478 Eriksson, P.G., 197, 203, 206, 210 Erlank, A.J., 13, 48, 51, 52,143 Ermanovics, I., 103,139 Ernst, W.G., 470, 474 Escher, A,, 350,432,440, 441 Esperanca, S., 364, 366, 373,378 Essence, E.J., 347 Essene, E.J., 303, 306, 309, 310, 319-321, 323325, 327, 329, 336, 339, 348-350, 354, 358, 380, 381 Etheridge, M.A., 45, 48, 113-115, 139, 172, 175, 316, 329, 344-346,353, 426, 439,441 Ethridge, EG., 463, 474 Eugster, H.P., 402, 412, 413 Evans, C.R., 154,175 Evans, J.A., 357 Evans, P.D., 136 Evans, S.H., Jr., 471 Evenchick, C.A., 133, 140 Ewart, A,, 8-10, 43, 49 Ewers, W.E., 402,413 Fabre, R., 36, 49 Fahrig, W.F., 103, 106, 133, 139, 140, 152, 153, 156,175, 290,294 Fairbairn, H.W., 244, 258 Falkum, T, 123,140 Fallick, A.E., 336, 338, 347, 445, 507 Falloon, TJ., I74 Farmer, G.L., 275, 277, 294, 486, 497, 498, 502, 504 Farmer, H., 147 Faulhaber, S., 307, 312, 323, 350 Fenglan, Z., 295
References index Fengqing, Z., 48 Ferguson, J., 377, 379 Fichter, L.S., 128,140 Field, D., 353 Filen, B.A., 136, 209 Fink, R.P., 385,415 Fiorentini, E., 336, 339,350 Fischer, L.B., 93, 443, 506 Fitches, W.R., 94 Fitz, TJ., 118, 141 Fitzsimons, I.C.W., 306, 315, 317, 326-328, 343, 350, 351 Flawn, P.T, 458, 467,474, 475, 479 Fleck, R.J., 91,441 Fleming, ES., 138 Fletcher, I.R., 497, 504 Floran, R.J., 413 Flower, M.EJ., 254,257 Flueh, E.H., 443 Foley, S., 171,175 Forbes, B.G., 132,133,135,146 Foster, C.T, Jr., 452, 479 Fountain, J.C., 265, 268, 294 Fourcade, S., 507 Fowler, C.M.R., 7 , 5 1 Fox, TP., 148,445 Frakes, L.A., 411, 413 Fralick, A.E., 54 Fram, M.S., 234, 254, 257 Francis, D.M., 75, 76, 86, 90, 92, 102, 140 FranGois, L.M., 406, 413 Franklin, J.M., 116,140 Freeman, M.J., 145 French, B.M., 101,148,199,213, 386,413 Frey, FA., 51,143 Friend, C.R.L., 49,175 Frimmel, H., 91 Fripp, R.E.P., 137 Frisch, T, 304, 322, 323, 325, 328, 332, 350 Frost, B.R., 252,257, 258,285,294,322,327,335, 350 Frost, C.D., 338,350,497,504 Frost, C.S., 257 Frost, E.G., 293 Flyer, B.J., 384, 395, 396, 413, 414 Fuhrman, M.L., 219,251, 252,257, 283, 285,294 Fukuoka, M., 356 Fullagar, P.D., 128,144 Fumerton, S., 92 Furnes, H., 131,140 Fyfe, W.S., 340, 350 Fyson, W.K., 87,90
References index Gaal, G., 11, 15, 46, 49, 106, 107, 123, 140, 184, 186,206,210,419-422,430,441 Gabelman, J.L., 452, 463, 474 Gabrielse, H., 140, 401, 414 Galer, S.J.G., 168,175, 488,504 Gallego, M.D., 149 Garde, A.A., 304,350 Garitpy, C., 497,504 Garrels, R.M., 402, 414 Garrison, J.R., 267, 294, 422, 441 Garrison, J.R., Jr., 498, 5#4 Garvey, O.C., 363, 367,378 Gasgarth, J.W., 28, 29,49 Gaskarth, J.W., 75, 90 Gass, I.G., 65, 71, 88, 90, 93 Gastil, G., 263, 294, 456, 474 Gehrels, G.E., 477,507, 508 Geist, D.J., 219, 252, 257 Gelinas, L., 8,49, 51 Ghazi, A.M., 259 Ghisler, M., 432, 441 Ghosh, P.K., 49 Gibb, R.A., 75, 90 Gibbs, A.K., 35, 49, 52 Giblin, P.E., 212 Gibson, I.L., 121, 140,177 Gibson, R.G., 475 Gierth, E., 417 Giles, C.W., 43,49 Gill, J., 294 Gill, J.B., 9, 49 Gilmer, T H., 117,138 Giovenazzo, D., 103,140 Girnis, A.V., 146 Gittins, J., 305, 319, 323, 326, 328, 331, 333, 349 Glassley, W.E., 304, 319, 350 Gleason, J.D., 265, 294 Gledhill, A.R., 141 Glen, R.A., 354 Glennie, K.W., 55, 90 Glickson, A.Y., 56, 91 Glover 111, L., 146 Glukhovsky, M.Z., 271,296 Goldberg, S.A., 265, 295 Goldsmith, J.R., 366,379 Goldsmith, R., 146 Goldstein, S., 375,380 Goldstein, S.L., 206, 209, 372, 375, 376, 377, 488, 489,494,502, 504 Gole, M.J., 385, 397, 406, 414 Golombek, M.P., 65, 90, 500, 501, 504 Golubev, A.I., 106, 108, I40 Golyshev, S.E., 408, 414
5 15 Goode, A.D.T, 302,305,354,396,409,4I4 Goodge, J.W., 475 Goodwin, A.M., 409,412, 414 Gorbatschev, R., 11, 15, 46, 49, 106, 107, 123, 140,422, 441, 444,506 Gorbatsev, R., 19,50 Gorbunov, G.I., 430, 444 Gordon, M.B., 119,140, 423,441 Gordon, TM., 47,305,350 Gorman, B.E., 145 Gould, D., 197,210 Gower, C.F., 106, 111, 140, 142, 147, 303, 305, 310,313,350, 355, 356,423,441, 444 Gower, R., 349 Graham, C.M., 129-131,140 Graham, R.H., 94 Grambling, J.A., 295, 305, 313, 350, 461, 463465,468,473, 474, 477-479 Grant, J.A., 324,325,340,341,350 Grant, S.M., 306, 322, 323,350 Grauch, R.I., 486, 504 Gray, C.M., 348 Greaves, M.J., 394-396,413 Green, A.G., 119,137, 138, 140,292 Green, D.H., 174,324-326,350, 352 Green, J.C., 3, 97, 105, 116-120, 138, 140, 141, 264,295 Green, TH., 254,257 Greenberg, J.K., 434, 442 Greenwood, P.B., 177 Greenwood, W.R., 65,91 Grew, E.S., 305,306,310,313,326,350,364,377 Griffin T , 293 Griffin, W.L., 304, 313, 325, 326, 328, 345, 351, 361,364, 365,367, 368,377, 378 Griffiths, N.H., 51 Griffiths, R.W., 114,138, 169, 170,174, 175,503 Griggs, A.B., 141 Grissom, G.C., 295 Grocott, J., 351 Groenewald, P.B., 306, 351 Gromet, L.P., 237,238, 257, 395, 414 Gross, G.A., 398,414 Gross, S.O., 252, 257 Groves, D.I., 97, 137 Gruau, G., 507 Guevara, A.N., 167,175 Guilbert, J.M., 454, 471 Guiraud, M., 348 Gulson, B.K., 432, 442 Gupta, A,, 38,49 Gurney, J.J., 86, 91, 155, 175 Gust, D.A., 52,377
516 Haapala, I., 268,271,295-297,430,430,438,442, 443 Haase, C.S., 385, 385, 386, 414 Hackman, B.D., 353 Hadley, D.G., 91 Haggart, M.J., 349 Haggerty, S.E., 239,257,381 Haines, S.M., 376, 382 Hakkinen, J.W., 351 Hakli, TA., 213 Halden, M.N., %-28,49 Halden, N.M., 75, 91, 93,351 Hall, A,, 10, 49 Hall, J., 379 Hall, L.M., 146,359 Hall, R.P., 45, 49, 109, 141, 153, 155, 156, 159, 161, 171,175, 208,210 m i , WD.M., 414 Hallett, R.B., 460, 467,474, 475 Halliday, A.N., 378, 497, 504, 506 Halls, N.C., 100,138, 147, 154, 156,174, 175 Hameurt, J., I 4 0 Hamilton, M.A., 246, 257 Hamilton, P.J., 54, 145, 202, 210, 378, 445, 497, 504, 506, 507 Hamlyn, P.R., 203, 210 Hammarstrom, J.M., 283, 299 Hammond, J.G., 267,295 Hammond, P., 252,257 Han, T.M., 415 Hanninen, E., 213 Hansen, E.C., 335, 336,351, 353,378 Hansen, W.R., 271, 277,295 Hanski, E., 13, 16,49 Hanski, E.J., 16, 49 Hanson, G.N., 176, 243,258, 260,296, 488,507 Harbour, R.L., 464,375 Hargraves, R.B., 251, 258 Harley, S.L., 4, 301, 304-307, 309, 310, 313, 315, 317-324, 326-328, 331, 332, 334, 336, 338, 339,343-345,348, 350, 351, 357 Harmer, R.E., 3,97,148,181,184,197,199-201, 203-206,214 213 Harmer, R.E.J., 101,148 Harmon, R.S., 379 Harms, U., 264, 296 Harper, G.D., 86,91, 134,141 Harris, C.N., 464, 475 Harris, L.B., 348 Harris, N.B.W., 52, 212, 297, 336, 338, 351, 356, 381, 494, 497,504, 505 Harrison, J.E., 121,141 Harrison, TM., 285, 286, 298
References index Harrison, 'IN., 349 Harry, W.T, 432,442 Hart, R.A., 90 Hart, R.J., 90 Hart, S.R., 51,143, 152, 164, 165,176, 363, 364, 373,378-380,477 Harte, B., 365, 373,378 Hartel, T.H.D., 331, 336, 338, 357 Hartman, H., 406, 414 Hartnady, C., 64, 65,91, 126,141 Hartnady, C.J.H., 65, 91, 94 Haskin, L.A., 239, 258, 260, 414 Hatcher, R., 128,141 Hatton, C.J., 154, 156, 159, 176, 181, 194, 200, 202,203, 206,208,210, 213 Hauser, E., 146 Hawkes, H.E., 380 Hawkesworth, C.J., 48, 126, 141, 152, 166, 176, 363-365,367,373,379-381,497,504, 505 Haydoutov, I., 86,91 Hayes, J.M., 412, 415, 417 Heaman, L.M., 28, 49, 151-154, 156, 169, 176, 279,281,282,295,505 Heath, S.A., 244, 258 Hebeda, E.H., 347 Hedge, C.E., 292, 297-299, 436, 445, 471, 475, 506 Hedlund, D.C., 265, 295, 452, 462, 475 Hefferan, K.P., 60, 63, 91, 93, 94 Hegner, E., 77, 83,91, 494,496,505 Heier, K.S., 328,351, 368, 378 Heikkineen, P., 440 Heintz, G.M., 478 Hellingwerf, R.H., 429, 440 Helmers, H., 148 Helmstaedt, H., 3, 52, 55, 58, 86, 87, 90, 91, 94, 95, 364,381, SO7 Hempton, M.R., 119,140,423,441 Henderson, C.E., 378 Henderson, J.B., 406,414 Henderson, J.R., 305, 351 Henderson, M.N., 351 Henderson, P., 371,378 Hendry, G.L., 53,212 Henriksen, N., 353, 432,442, 446 Henry, G., 125,141, 401,414 Hensel, H.D., 377 Hensen, B.J., 304, 309-311, 313, 323, 324, 326, 328,332,341,351, 352, 357, 358 Herd, R.K., 306, 326,352 Hergt, J.M., 152, 161, 172, 173,176 Hermans, G.A.E.M., 357 Hertogen, J., 232, 239, 246, 256
References index Hem, N., 218,235,258,264,295 Hetherington, E.A., 459, 464,474 Higgins, A.K., 446 Hildebrand, R.S., 30,49 Hildreth, W., 152,176, 206,210, 289, 295 Hill, J.D., 222, 223, 250, 251, 258 Hill, R.I., 170, 176, 503 Hillhouse, J., 505 Hills, A X , 294, 298 Hills, D.V, 381 Hinze, W.J., 116,138, 140, 141, 145, 148, 149 Hirdes, W., 443 Hiroi, Y., 307, 309, 329, 352 Hirschleber, H., 440 Hjelt, S.E., 429,442 Hobday, D.K., I48 Hodge, D.S., 218,258, 294 Hoefs, J., 347, 397, 417 Hoernes, S., 350 Homauer, R., 350 Hoffman, A.W., 381 Hoffman, P.F., 22, 29, 30, 49, 50, 56, 72, 74-76, 83,86, 87,91,99,103,109-111,127,128,141, 253, 258, 263, 290, 295, 384, 414, 422424, 427, 428, 432, 434, 438, 439, 442, 467, 475, 496,505 Hoffren, V, 146 Hofman, H.J., 402, 414, 418 Hofmann, A.W., 497,505 Holcombe, R.J., 455, 471 Hole, M.J., 172, 176 Holland, H.D., 402, 414 Holland, TJ.B., 320, 321, 324, 352, 355 Hollis, J.D., 366, 382 Hollister, L.S., 283, 295 Holloway, J.R., 340, 343,358, 366,381, 438, 445 Hollta, P., 302, 304, 322, 325, 329, 331, 334, 345, 352 Holm, D.K., 106, I41 Holm, P.E., 8-10, 19, 50, 53, 98, 122, 123, 141, 147 Holst, TB., 141 Holzl, S., 311, 313, 352 Honkamo, M . , 17,50, 106, I41 Hooker, P.J., 145 Hopgood, A.M., 313,315,315-317,352 Hoppe, W.J., 265, 271, 295 Hormann, P.K., 302,304,336,352, 421,442 Horrocks, P., 137 Houseman, G.A., 171,176,352 Housh, TB., 503 Houston, R.S., 84, 92, 452, 464,474, 475 Howard, K.A., 449,475, 479
5 17 Howell, D.G., 475 Hoy, T, 137 HOy, Tiygve, 120,141, I42 Huang, C.H., 47 Hughes, C.J., 184, 210 Hughes, D.J., 49, 109, 241, 155, 156, 159, 161, 171,175,208,210 Hughes-Clarke, M.W, 90 Huhma, H., 49, 50,74,91,295, 296,348,428,442, 495,498,505 Huiqin, F., 295 Hulbert, L.J., 201, 203,212 Humphreys, H.C., 325,352 Hunt, PA., 47, 217, 219, 220, 257, 267, 294, 343, 349 Hunter, D.R., 148, 194, 203,210, 211, 306,352 Hunter, R.H., 364, 378 Huntington, H.D., 242, 244,258 Huppert, H.E., 151,176 Husler, J.W., 479 Hutchinson, D.R., 116, 119, 120, 137, 138, 140, 142,292 Hutton, D.H.W., 349, 432,442 Hynes, A.J., 56, 75, 76, 86, 90-92, 240, 486, 505 IanelIi, TR., 132, 133, I42 Iden, I.K., 351 Indares, A., 302, 305, 306,317, 319, 323,352 Irifune, T , 166, 177 Ilvine, TN., 184, 189-192, 202, 203, 208, 211, 212,231,258,267,295 Irving, E., 7, 50, 102, 142 Isachsen, C.E., 503 Isachsen, Y.W, 296,315, 318,354, 359 Ishihara, S., 268, 284, 295 Ishizuka, H., 356 Jaanus-Jarkkala, M., 48 Jachems, R.C., 478 Jackson, D.H., 356 Jackson, G.D., 132,133,142, 494, 496,505 Jackson, I., 376,377, 378 Jackson, M.P.A., I48 Jackson, N.J., 264,295 Jackson, P.M., 378 Jackson, TA., 47 Jacobsen, S.B., 396,413, 414 Jacobsen, S.G., 390, 396,413 Jahn, B., 35,50 James, D.E., 364,378, 500,505 James, H.L., 383-385,398,412,414, 415 James, P.R., 426, 441 James, S.D., 43,50, 114,142
518 Jan, M.Q., 176, 441, 442, 444 Janardhan, A.S., 335, 351, 353 Jansen, H., 111, 112,142 Jansen, J.B.H., 357 Javoy, M., 246,256 Jefferson, C.W., 121,132, 133,135,137 Jensen, S.B., 446 Jiang, J., 336, 353 Jianhua, Yu, 264, 271,295 Jolo, X.J., I77 Jocelyn, J., 378 Johanson, B., 211 Johansson, A., 123,142 Johansson, L., 123,142 John, B.E., 475 Johnson, C.D., 328,353 Johnson, C.M., 363,373,378 Johnson, M.C., 283,295 Johnson, Y.A., 21, 22, 50, 109,142 Johnston, A.D., 438,444 Jolly, W.T., 22, 50 Jones, A.P., 365, 367, 378 Jones, D.L., 441, 447, 475, 499, 505, 507 Joron, J.L., 54,178 Joubert, P., 91, 141 Jyothender Reddy, Y., 234, 258 Kahkonen, Y., 19,50, 429, 439, 442, 498,505 Kalamarides, R.I., 256, 260 Kalsbeek, E, 31, 50, 152, 176, 302, 313, 318, 319, 353, 432,442, 496,505 Kamo, S.L., 130,142 Kaplan, I.R., 417 Kapp, H.E., 190,212 Karlstrom, K.E., 31, 47, 84, 92, 427, 442, 447457, 464, 465,467, 467-469, 471, 472, 475 Karson, J.A., 91, 94 Kasch, K.W., 64, 92 Kasting, J.F., 406, 415 Kaufman, A.J., 401, 402, 412, 415 Kaufman, S., 146,147 Kay, R.W., 365, 366, 378, 379 Kay, S.M., 365, 366, 378, 379 Keays, R.R., 203,210 Keil, K., 381 Keith, D.W, 211 Keith, S.B., 271, 277, 295 Kemp, J., 68, 92 Kempton, P.D., 176,363,364,375,379, 381 Kennedy, WQ., 59, 65, 92 Kent, S.C., 452, 463, 472, 475 Kerr, A., 7, 45, 46, 50, 264, 268, 294, 296, 498, 505
References index Kemch, R., 279,281,282,299,347 Kesola, R., 441 Kesse, G.O., 443 Kessler, E.J., 271, 277, 296 Key, R.M., 302, 353 Khan, M. Asif, 432,442 Khan, M.A., 156,176 Kidd, W.S.F., 47, 89 Kielinczuk, S., 47, 89,293, 503 Kim, S.-J., 281, 282, 296 Kimbell, G.S., 210 King, E.R., 116,142,193,211 King, J.E., 503 King, P.B., 458, 467,475 Km, J., 296 Kirschvink, J.L., 411,415 Kizaki, K., 352 Kleemann, G., 202,211 Klein, C., 5, 384-397, 399, 401-408, 411, 412415, 418 Klerkx, J., 124,142 Klewin, K.W., 118,119,142 Knoper, M.W., 31,50,452, 453, 459, 460, 462,475 Knutson, J., 377 Koch, R.J., 293 Kohler, H., 352 Koistinen, TJ., 71, 92, 93 Kojima, H., 356 Kojima, S., 356 Koljonen, T, 146 Kolker, A., 248,25&252,257-259,265,278,284, 296 Kominz, M.A., 135,137 Kontinen, A., 17, 50, 57, 58, 73, 74, 86, 92, 106, 108,142, 428, 439,443, 496,505 Korotev, R.L., 414 Korsch, R.J., 143 Korsman, K., 304,353,429,443 Kouvo, O., 297,428,444,506 Kozlovsky, E.A., 15, 50 Kramers, J., 302,305,312,357 Kramers, J.D., 494,503, 504 Kranck, S.H., 415 Krane, K., 413 Krill, A.G., 132,142 Krill, K.O., 421, 443 Kroenke, L.W, 177,507 Krogh, E.J., 351 Krogh, TE., 49,142,220,258,432,445 445 Kroner, A., 56, 59, 60, 64, 65,85, 92, 97, 124, 125, 141, 142, 170, 176, 311, 353, 401, 415, 436, 437,443 Kruger, EJ., 202, 211, 213
References index Krummenacher, D., 293 Kudo, A.M., 381 Kukla, PA., 126,142, 443 Kulikov, VS., 146 Kulikova, V.V.,146 Kumar, S., 444 Kumpulainen, R., 131,143 Kuno, H., 366,379 Kuovo, O., 495,506 Kushiro, I., 254, 258 Kusky, TM., 87,92 Kuster, D., 264, 296 Kuzmin, M.I., 149 LaBerge, G.L., 398, 402, 415, 417 Labotka, TC., 449,475 Lafon, J.M., 264, 293, 296 Lagerbland, B., 19, 50 Lahtinen, J.J., 136, 185-187, 209, 211 Laing, W.P., 354 Lal, R.K., 304, 326, 328,353 Lamb, R.C., 305, 313, 329,353 Lamb, W.J., 358 Lamb, W.M., 301, 331, 336, 338, 339,353, 358 Lamothe, D., 52, 58, 76, 92,146 Lancelot, J.R., 61, 90 Langmuir, C.H., 152,176 Lanphere, M.A., 475 h n z i r o t t i , A., 460, 465, 476 Larsen, L.M., 174 Larsen, M., 138 Larson, R.L., 167, 169,176 Larue, D.K., 148,433,434,443, 445 Lassiter, J., 381 Lavreau, J., 142 Lavrov, M.M., 136,209 Leat, P., 177 LeBlanc, M., 55, 58-61, 63,89, 92, 127,143 Le Breton, N., 340, 353 LeCheminant, A.N., 143, 151-154, 169,176 LeCheminant, G.M., 110, 111,143 Ledent, D., 36,50 Lee, M., 137,138 Lee, M.W, 137,292 Lee, N.W., 140 Lee-Huh, C.N., 292 Leelanandam, C., 215, 218, 234,258, 264,296 Leeman, W.P., 363,379 Lees, G.J., 9, 51 LeFort, P., 507 Legg, J.H., 182, 213 Lehnert, K., 47, 136 Lehtonen, M.I., 441
5 19 Leonardos, O.H., 417
Le Roex, A.P., 10, 51, 134,143 Lesher, C.M., 385,416 Leube, A., 435,443 Leventhal, J.S., 404, 416 Lewry, J.E, 21, 24, 28, 29, 51, 56, 7 5 9 2 , 95, 101, 143,437,443 Leyleroup, A., 365,368,377, 379 Leyleroup, A.F., 377 Li, T-D., 444 Libby, W.G., 504 Lidiak, E.G., 265,296 Liegeois, J.P., 117, 142 Liew, TC., 497,505 Lightfoot, PC.,119,143 Lindberg, PA., 454,476 Linder, D., 354 Lindqvist, G., 440 Lindroos, A., 48 Lindros, A., 429,441 Lindsay, J.F., 133, 135,143 Lindsley, D.H., 239, 248, 251-253, 256-258, 284, 294, 296,417 Lindstrom, D.J., 258 Lindstrom, M.M., 258 Link, P.K., 134,139, 141 Linn, A.M., 503 Liprnan, P.W, 378 Lippard, S.J., 93 Livesey, C., 51 Livingston, D.E., 295 Lochhead, A,, 296 Liifgren, C., 19, 51 Loiselle, M.C., 267, 296 Long, C.B., 54,149 Long, L.E., 294 Longhi, J., 229, 234, 254, 257, 259 h o c k , G., 366,379 Loosveld, R.J.H., 316, 329, 344-346,353 Lowe, D.R., 406,412, 416 Lowry, P.H., 478 LUGIS, S.B., 52, 75-78, 87, 93-95, 102, 147, 444, 507 Ludden, J., 8, 51 Ludden, J.N., 90,140 Ludwig, K.R., 456,476, 478 Lugmair, G.W., 381 Lull, J.S., 476 Lumbers, S.B., 147 Lund, C.E., 440, 443 Luosto, U., 421,443 Lyle, M., 413
520 Maboko, M.A.H., 305, 313, 322, 325, 326, 333, 349, 353 Macambira, M., 293, 296 Macaudiere, J., 347 Macdonald, R., 51,93,95 MacDougall, J.D., 86, 95 Machado, N., 49,384,413 Maczuga, D.E., 177,259 Madrid, R.J., 479 Mahoney, J.J., 177, 507 Maiden, K.J., 64, 65, 89, 123-126,138, 141, 414 Maijer, C., 357 Malni, O., 351 Mangini, A,, 405, 416 Manton, W.I., 305,310,313,326,350 Mantovani, M.S.M., 176 Maquil, R., 237,256, 259, 294 Marcantonio, E, 497, 505 Margulis, L., 412 Marjoribanks, R.W., 315,316,318,320,354 Marker, M., 419-421,440, 441, 443 Marlow, A.G., 497,504 Marques, J., 356 Marrett, R., 152,175 Marriner-177 Marsh, J.S., 8-10, 13, 15, 16, 39, 48, 51, 98, 143, 202,205,211 Marsh, N.G., 178 Martell, C., 459,473 Martignole, J., 217, 219, 259, 302, 305, 306, 317, 319,323,352, 355, 356,444 Martin. A,, 406, 416 Martin, H., 64, 93 Martin, S.E, 356 Mason, TR., 406,414 417 Massey, N.W.D., 119,143 Matsubara, S., 354 Matsueda, H., 354 Mattey, D.P., 177, 178, 356 Matty, D.J., 379 Mauer, R., 443 Mawer, C.K., 350,474, 477, 478 Max, M.D., 54,149 Maynard, J.B., 506 Mazzucchelli, M., 381 McCall, G.J.M., 56, 93, 209 McCallum, I.S., 260 McCallum, M.E., 459 McCarthy, T.S., 51,144, 209 McClay, K.R., 100,143, 187, 211 McClelland, W.C., 477, 507 McClenaghan, M.P., 145 McClennan, S.M., 7,51
References index McCourt, S., 197, 209 McCrink, TP., 33,48, 452, 453,473, 476 McCulloch, M.T, McCulloch, M.T, 45, 51, 114, 143, 176, 177,178, 348, 356, 365, 377, 379, 380, 382, 426, 443, 446,484,485,488,495,497,505, 506 McDermott, E, 497,498,504, 505 McDonough, WE, 166,178,380 McDougall, I., 176, 353 McDowell, S.D., 475 McGlynn, J.C., 7 , SO, 102,142 McIlwaine, W.H., 140 Mcintyre, R.M., 378 McKee, C.G., 455,476 McKee, E.H., 133,139 McKenzie, D.P., 7,51,97, 114,143,148, 153,166, 167,169, 170,176, 178,352 McKinney, C.R., 271,298 McLellan, E.L., 354 McLelland, J.M., 219, 220, 259, 263, 267, 267, 268, 287, 293, 296, 313, 315, 318, 328, 333, 343,349, 354, 358, 359, 497,503 McLennan, S.M., 369, 370,379-381, 491,506 McMechan, M.E., 120,122,143 McNaughton, N.J., 417 McNutt, R.H., 295,494,504, 505 McSwiggen, P.L., 116,138, 143, 147 Medaris, L.G., 298 Meisheng, G., 295 Meixner, H.M., 182,211 Melezhik, V.A., 408, 416 Mellinger, M., 49 Mendelson, C.V., 402,416 Mengel, E , 302, 305,313, 319, 325, 33,345,354 Menuge, J.E, 245,259 Menzies, M.A., 177,378, 379, 381 Merilainen, K., 421,443 Meschede, M., 98, 129,143 Metzger, K., 361,362, 376,377 Meyer, H.O.A., 363,379 Meyer, M., 51 Meyer, R.P., 116,144 Meyers, R.E., 253,259 Mezger, K., 333,354 Michard, A., 44 54, 394, 395, 416, 440, 445, 502, 503, 507 Michard, G., 394,395,416 Michot, J., 256, 297 Michot, P., 220, 259 Milkereit, B., 137, 138, 140, 292 Millard, H.T, 292 Miller, A.R., 143 Miller, C.E, 295
References index Miller, D., 479 Miller, D.M., 272, 299, 449, 450, 468,479 Miller, E.L., 479 Miller, J.D., 119,143, 194, 211, 220, 259 Miller, J.D., Jr., 119, 143 Miller, R. McG., 48, 125, 126,141 Miller, R.M., 504 Miller, R.McG., 144 Millward, D.G., 167,176 Min, T., 48 Minster, J.E, 416 Minter, W.E.L., 148 Misra, K.C., 128,144 Mitchell, A.H.G., 436, 440 Mitchell, J.N., 248, 259 Miyano, T, 385,416 Miyashiro, A., 108,144, 296 Moecher, D.P., 304, 324, 336, 339,354 Mocller, F?, 396, 416 Moench, R.H., 455,477 Moinet, E., 381, 445 Mokshantsev, K.B., 148 Molnar, P., 176, 352 Monger, J.W.H., 441,475 Montgomely, C.W., 295 Moody-Stuart, M., 90 Moorbath, S., 131, 144, 152, 176, 206, 210, 289, 295,379 Moore, A.C., 302, 305,354 Moore, J.M., 422, 443 Moorehead, J., 92 Moores, E.M., 55-57, 59, 66, 76, 83, 85, 86, 93, 99,144 Mora, C.I., 305, 313, 354 Moralev, V.M., 271, 296 Moreau, B., 440 Moreau, C., 256 Morey, G.B., 104, 106, 117, 138, 143-145, 147, 193,213, 398,416 Morgan, G.E., 7 , 51 Morgan, P., 97,136 Morris, R.C., 402,416 Morris, S.L., 379 Morrison, D., 177 Morrison, D.A., 144, 259 Morrison, J., 4, 97,136, 246, 259, 263, 294, 358 Morrison, M.A., 53, 212 Morse, S.A., 184, 189, 190, 211, 215, 217-219, 221-224, 235, 237, 243, 2h4, 246, 253, 254, 257, 259,290,297 Morton, M., 210 Mose, D.G., 271, 277,292, 297 Mosher, S., 459,476, 479
521 Motoyoshi, Y., 307, 309, 315, 329, 331, 352, 354, 357 Moyes, A., 52 Mudrey, M.G., Jr., 117,139, 141 Muehlberger, W.R., 267,271, 277,297 Mueller, R.E, 416 Muhling, J.R., 304, 313, 316,354 Mukhopadyay, B., 277,297 Mullen, E.D., 98,108, 129, 134,144 Miiller, P.J., 405, 416 Murphy, D.M.K., 307,359 Murphy, J.B., 438,443 Murray, C.G., 145 Murray, D.P., 146 Mutanen, ' I 187,211 , Myers, J.S.,86, 93, 145, 424-426, 443 Myers, P.E., 141 Myers, R.E., 45,51,98, 100,144 Naidoo, D.D., 61, 93 Naldrett, A.J., 193, 208, 209, 212 Nance, R.D., 438,443 Nance, WB., 371,379 Narsimha Reddy, M., 218, 258 Nash, J.T., 454, 471 Nassief, M.O., 69, 93 Natapov, L.M., 149 Nealson, K.H., 412 Neary, C.R., 88 Nelson, B.K., 263, 275, 277, 278, 287, 288, 297, 361, 379, 451, 455, 468, 470, 476, 484, 494, 500,501,506 Nelson, D., 146 Nelson, D.O., 100,144 Nelson, D.R., 417 Nesbitt, R.W., 53,178 Nesse, W.D., 459, 476 Newton, R.C., 301, 321, 335, 348, 351, 353, 354, 366,379 N G o m , P.M., 36, 51 Nicholas, R.L., 465, 476 Nichols, G.T., 306, 355 Nicholson, S.W., 118-120,144 Nickeson, P.A., 137 Nicollet, C., 305, 307, 310,311, 323, 325,326,355 Nielsen, K.C., 478 Nielsen, TED., 153, 155, 156, 162,17$ 176 Nielson, D.L., 471 Nielson, H., 417 Niernela, R., 353 Nimz, G.J., 364, 379 Nironen, M., 429,443 Nisbet, E.G., 7 , 51,416
522 Nixon, P.H., 177,362,378, 379 Noblett, J.B., 460, 476 Nokleberg, W.J., 476 Nolan, K., 235, 259 Noll, P.D., Jr., 456, 473, 476 Norman, A.R., 304,316, 333,355 Norman, D.I., 464, 476, 497, 506 Norry, M.J., 8,52, 134,145,177 Nriagu, J.O., 412 Nunn, G.A.G., 445 Nurmi, P.A., 429, 430,443 Nuter, J.A., 462, 473 Nutman, A.P., 350 Nystuen, J.P., 131,140, 143, 144 Nzenti, J.P., 347 Obradovich, J.D., 122, 144, 149 O’Brien, S.J., 147 Ocola, L.C., 116,144 Odom, A.L., 128,144 Oen, I.S., 108,144 Oen, O.E., 432,442 Offe, L.A., 356 O’Hara, M.J., 178 O’Hara, N., 145 O’Hara, P.E, 454,476, 478 Ohlander, B., 54, 495,506, 507 Ohnenstetter, D., 89 Ohnenstetter, M., 89 Ohr, M., 491, 506 Ojakangas, R.W., 117,145,409,416 Okrusch, M., 366,379 Oliveira, E.P., 164, 170, 177, 178 Oliver, J., 146, 147 Oliver, J.E., 128, 138 Oliver, R., 52,146 Oliver, R.A., 50, 142 Oliver, R.L., 177, 304, 307, 348, 355 Oliviera, E.P., 53 Olnisted, J.E, 119,145 Olsen, K.I., 351 Olson, J.C., 452, 462, 475 Olson, K.E., 253,254,259 O’Neil, J.R., 219, 260, 292, 333, 336, 339,358 O’Nions, R.K., 109,145, 166,176, 178, 307, 311, 333,336, 343,348,377,497,504 Onstott, TC., 146, 302, 304, 317, 333, 357 Oray, E., 116, 145 O’Reilly, S.Y., 326,351, 361, 378 Ormaasen, D.E., 351 Orrell, S.E., 298, 478 Ortega, G.E, 381 Ortega-Gutierrez, E , 380,477,508
References index Osanai, Y., 356 Owen, J.V., 305,355 Owen, L.B., 258 Owen, V., 106,111,140 Oxburgh, E.R., 333,334,345,346,355 Pacca, I.G., 146 Paces, J.B., 116, 117, 119,139, 145 Packer, B.M., 406, 417 Padalko, N.L., 414 Padgham, W.A., 91 Padovani, E.R., 363,364,366,367,373,376,378380,477 Page, B.M., 149 Page, R.W., 43,51, 54,149 Palacz, Z.A., 176,503 Pallister, J.S., 65-71, 87, 93, 437, 443, 452, 471, 476, 477,498, 500,501,503, 506 Palmer, H.C., 116,145 Pankhurst, R.J., 504 Papike, J.J., 413, 416, 417 Papunen, H., 213,430,444 Parchman, M.A., 472, 479 Park, A X , 17,51,56,71-74,86,93,106,145,355, 427-429,444, 447,476 Park, R., 316,355 Park, R.G., 7 , 17,50, 51,142,174 Parker, A.J., 54, 116,145, 149 Parr, J., 108,145 Parrish, R.R., 76, 77, 83, 93, 101, 140, 145 Parry, S.J., 212 Parslow, G.R., 28, 29, 49, 75, 90 Parsons, I., 181, 212 Pasteels, P., 264, 297 Patchett, J., 5, 11, 31, 51, 427, 428, 444 Patchett, P.J., 51, 99, 106,145, 277, 278, 287, 291, 297, 361, 380, 428, 432, 440, 444, 458, 476, 477,484-489,493-496,498-502,502, 503, 506508 Paterson, M.S., 234,260 Patifio Douce, A.E., 438, 444 Pattison, D.R.M., 303, 343,355 Peacock, M.A., 267, 297 Peacock, S.M., 166,177 Peacor, D.R., 380,506 Pearce, J.A., 8-10, 12, 13, 15, 18-20, 29, 39, 45, 50-53, 62, 75, 84, 93, 95, 98, 106, 108, 129, 134,142,145, 203-205, 212,273,297 Pearce, T.H., 98,145 Pearson, D.G., 168,177 Peart, R.J., 182, 211 Peate, D.W., 176 Pechenkin, S.A., 414
References index Peck, D.C., 29, 30,52 Pedersen, L.B., 444 Pegram, E., 295 Pellaton, C., 92 Perera, L.R.K., 353, 356 Perfit, M.R., 8, 52 Perkin, D.J., 54 Perkins, D., 320-322, 327,354, 355 Perreault, S., 306, 351, 355 Perry, E.C., 397, 402, 408,416 Pesonen, L.J., 141 Peterman, Z.E., 53, 104, 106, 144, 147, 149, 271, 277, 295, 297, 427,440, 444,502, 507 Peters, E.K., 295 Petersen, E.U., 366,380 Petersen, J.S., 123,140 Peterson, N.P., 457, 476 Petterson, M.G., 432, 444 Pettingill, H.S., 267, 297, 497, 502, 506 Peucat, J.J., 440,503, 507 Pharaoh, T.C., 8, 11-21, 45, 52, 102, 106, 108, 145, 429,444 Phillips, G.N., 304, 309, 316, 329, 355 Philpotts, A.R., 253, 259 Philpotts, J.A., 371, 380 Phinney, W.C., 144, 155,177, 218, 259 Piboule, M., 52,146 Pieard, C., 26, 27, 52, 92, 102, 103,140, 146 Pidgeon, R.T., 353,505 Pietsch, B.A., 145 Piirainen, T , 184, 186, 187, 209 Pilaar, W.F.H., 90 Pimentel-Klose, M.R., 397, 414 Pin, C., 89,381,445 Piper, J.D.A., 7 , 52 I’lalker, G., 470, 476 Podmore, E, 100,146,183,212 Podosek, F.A., 493,495,503 Porada, H., 64, 93, 125,146, 355 Postaire, B., 440 Poulsen, K.A., 140 Powell, R., 304-307,320,321,323-325,328,331334, 340, 341,346,348, 352, 355-357 Powers, R.E., 336,355 Poyner, R., 138 Prame, W.K.B.N., 351 Pratt, T, 119, 146 Preiss, W.V., 132, 133, 135,146 Premo, W.R., 93,443,452,467,472, 476, 477,506 Prendergast, M.D., 100,149,182-184,212, 213 Presper, T, 362, 367, 369,375,377,380 Price, R.A., 120, 122, 135,143, 146, 147 Price, R.C., 293
523 Pushkar, P.D., 295 Pye, E.G., 181,212 Quick, J.E., 69, 70, 93,506 Raase, P., 304, 323, 336, 352, 355, 356, 441, 442 Raczek, I., 381 Rahaman, M.A., 136 Raiswell, R., 404,416 Raith, M., 302, 304, 307, 309, 312, 316, 317, 323, 336,344,347, 350, 352, 355, 421,440, 442 Rajamani, V., 38, 46,46 Ramberg, I.B., 217,260 Ramo, O.T, 268, 271,277,295, 297 Ramo, T , 430,438,442 Rankin, D.W., 105,128,129,146 Ransome, EL., 457,477 Ransome, I.G.D., 91 Ranson, W.A., 223,259, 245, 297 Rasmussen, T M . , 429,444 Rast, N., 105,146 Ratcliffe, N.M., 146 Rathbone, P.A., 210 Rattray, G., 381 Ravindm Kumar, G.R., 348, 351 Read, J.E, 146 Reddy, M.N., 264,296 Reed, J.C., Jr., 128,146, 447, 449, 452, 453, 459, 471, 472, 477 Register, M.E., 271, 277, 297 Reid, A.M., 51,243 Reid, D.I., 36-38, 52 Reid, M., 465,477 Reid, M.R., 363, 364, 366,380 Reinhardt, B.M., 90 Reino, J., 213 Reischmann, T, 443 Renne, P.R., 126,146 Renner, R., 35,52 Retief, E.A., 213, 417 Reuber, I., 91, 94 Reuss, R.L., 449, 460,477 Rex, A.J., 444 Rex, D., 441, 444 Rex, D.C., 48 Reymer, A., 498, 501,506 Reymer, A.P.S., 168,177 Reynolds, M.W., 120,146 Reynolds, P.H., 130,148 Reynolds, S.J., 295 Ribbe, P.H., 235,259 Richards, M.A., 499,507 Richardson, S.W., 329,349
524 Richman, D.L., 294 Rickard, D., 108, 145 Rickwood, P.C.,145 Ridley, J.R., 478 Ries, A.C., 94, 504 Ringwood, A.E., 166,173,177,324, 325,350 Rivalenti, G., 381 Rivers, T, 302, 303,305,310,313, 315-319,325, 328, 344,345,354-356,422, 423,444 Roach, R.A., 51 Robbins, E.I., 402, 415, 417 Roberts, J.L., 207, 212 Roberts, R.G., 444 Roberts, S., 93 Roberts, S.J., 361, 364, 366-368, 375, 376, 380 Robertson, J.M., 31, 33, 52, 449, 455, 463, 464, 477 Robinson, D.N., 363, 367,378 Robison, L.C., 473 Roddick, J.C., 131,138, 141 Roddick, J.H., 414 Rogers, J.J.W., 369, 380 Rogers, N.W., 176, 365,367,367,373,380, 381 Romer, R.L., 428,444 Rondot, J., 352 Roobol, M.J., 58, 67, 68, 94 Roots, C.E, 134,146 Roper, H., 409,417 Rose, A.W., 370,380 Rosing, M.T., 487,503,507 Rosman, K.J.R., 504 Ross, G.M., 134,146 Roths, P.J., 463,477,478 Rozendal, R.A., 464,465,476 Rudnick, R.L., 361-363, 365-369, 372, 374, 375, 377,380,458,477 Ruiz, J., 4, 287, 297, 361, 363, 364, 366-369, 372, 373,376,380,381,458,476, 477,495,506, 508 Rundle, C.C., 353 Rutherford, M.J., 283, 295 Rutland, R.W.R., 48,139,175,354,441 Ryabchikov, I.D., 106,146 Ryan, A.B., 427, 441 Ryan, B., 111,137, 140,222,259 Ryburn, R.J., 358 S a , H.S., I78 Sacchi, R., 305, 310,356 SACS (South African Committee for Stratigraphy), 124,146 Sahama, TG., 271,297 Sakiyama, T, 356 Sakko, M., 429,445
References index Salpas, P.A., 239, 258, 260 Samson, S.D., 470,477, 497, 499, 501,506 507 Sanders, I S , 307, 356 Sandiford, M., 312,316,327, 334, 344,346,356 Sangster, D.E, 70, 75,93, 94 Santosh, M., 307, 322, 325, 328, 331, 332, 335, 336,338,339,356 Saquaque, A,, 61,91, 93, 94 Sassarini, N., 293 Sauer, P.E., 462,477 Saunders, A.D., 29, 52, 152, 166, 167, 171, 172, 176-1 78,507 Saunders, A.O., 84,94 Saverikko, M . , 106,146 Savolahti, A,, 271, 297 Sawkins, EJ., 253, 260 Sawyer, E.W., 64, 89, 125,137 Scharer, U., 111,147,494,495,507 Schenk, V., 331,356 Schidlowski, M., 409, 417, 418 Schiffries, C.M., 227, 238, 257 Schiotte, L., 503 Schmidt, D.L., 65, 91 Schmidt, R.G., 417 Schnetzler, C.C., 371, 380 Schnutgen, A., 379 Schoch, A.E., 264,293 Scholz, E.A., 471 Schopf, J.W., 402,406,414-417 Schrank, A., 35, 50 Schreiber, U.M., 210 Schreurs, J., 304, 313, 315, 322, 325, 327-329, 334,336,338,345,356 Schreyer, W., 347 Schrijver, K., 217, 219,259 Schroder, B., 379 Schubert, G., 168,177, 498,501,502,503, 506 Schulz, K.J., 53, 75, 94, 106,142, 147, 444, 507 Schulze, D.J., 86,91, 364,381 Schumacher, R., 307,312,323,325,328,333,356 Schuster, R.D., 298,472 Schwarcz, H.P.,298 Schweitzer, J., 197-200,212 Scoates, J., 226, 236,260 Scoates, J.S., 259 Scoates, R.EJ., 25-27,47, 75,76,86,88,103,137, 188, 189,209, 212 Scott, D.J., 3, 26,52, 55, 57, 58, 77, 79, 80, 82, 83, 94, 95,444, 507 Searle, M.P., 55, 94, 432,444 Sears, J.W., 120, 122, 135,147 Seck, H.A., 381 Secor, D.T, Jr., 146
References index Seeley, J.M., 464, 477 Seguin, M.K., 128,147 Sehlstedt, S., 445 Seibert, J., 138 Seifert, F., 352, 442 Seifert, K.E., 239, 258 Selverstone, J., 364,381 Sengupta, P., 304,323,326,356 Senior, A., 305, 313,358 Serpa, L., 116,147 Setzer, T, 147 Sevigny, J.H., 133, 134,147 Shackleton, R.M., 65,94 Shadel, C.A., 452,473, 477 Shafiqullah, M., 295 Shake], D.W, 271,297 Shand, S.J., 297 Shanti, M., 58, 67, 68, 94 Sharaskin, A.Ya., 53 Sharp, J., 147 Sharp, WN., 292 Sharpe, M.R., 184, 194, 197, 199-203, 207, 208, 210-213 Shastri, L.L., 471 Shaw, D.M., 295, 298 Shaw, R.D., 311,315,316,318,325,356,357 Shegelski, R.J., 409, 417 Sheppard, S.M.E, 503,507 Sheraton, J.W., 45, 52, 154, 156, 159, 161, 163, 177,315, 316,319,349, 356 Shieh, Y.-N., 279, 281, 281, 282, 295, 298 Shiraishi, K., 307,352, 356 Shirey, S.B., 118-120,144,378, 488, 507 Shive, P.N., 260 Shonk, K.N., 452,462,477 Shufflebotham, M.M., 51 Sial, A.N., 45, 53 Sides, J.R., 292 Sighinolfi, G.P., 368, 381 Sigurdsson, H., 152,177 Silberling, N.J., 505 Silver, L.T., 267,271,273,274,293, 297, 298,449, 454,456-458,465,471, 473, 477-479 Simmons, E.C., 220, 235,243-246, 251,260,417 Simmons, G., 379 Simmons, K.R., 220,251,260 Simonen, A,, 71, 94 Simonson, B.M., 398, 409, 417 Simpson, C., 475 Simpson, R.W., 456, 465, 478 Sims, P.K., 24, 53, 104, 106, 147, 427, 433, 434, 440, 444,498,502, 507 Singewald, Q.D., 460, 472
525 Sinha, A.K., 128,136,297,506 Sinha, M.N., 140 Sinigoi, S., 381 Siroshtan, R.I., 412 Sisson, V.B., 295 Sivell, W.J., 41-43, 45, 53 Skiold, T , 20,53, 495, 498, 506, 507 Slatt, R.M., 454, 478 Smalley, P.C., 353 Smellie, J.A.T, 445 Smith, B.M., 472 Smith, C.H., 190-192, 211, 212 Smith, D., 364, 377 Smith, E.I., 434,444 Smith, H.S., 94 Smith, J.V, 354, 378 Smith, R.E, 350 Smith, R.W., 416, 476,506 Smith, YE., 2, 7 , 10, 18, 29, 30, 47, 52, 53, 122, 123,147 Smith, VC., 478 Smithson, S.B., 217,218,258, 260 Smolkin, VE, 49, 209,136 Snyder, D., 231,260 Snyder, G.A., 246,260 Snyder, G.L., 49,175, 452,478 Snyman, J.A., 194,207,212 Soegaard, K., 458,463, 468,472, 478 Solheim, S., 140 Solyom, Z., 129, 131,147 Sonder, L.J., 356, 433, 444 Sorensen, K., 302, 304,319,350 Soucie, G., 8, 47 Southwick, D.L., 100,106,147 Spall, H., 55, 90 Sparks, R.S., 151,176 Speer, J.A., 219,260 Spencer, C., 137,138,140,292 St-Onge, M.R., 52, 57, 75-78, 83, 86, 87, 93-95, 102,147,439,444,496,507 Stacey, J.S., 93, 436, 441, 443, 445, 475, 479, 506 Stdlhos, G., 445 Stanistreet, I.G., 126,141, 142, 414, 439,443 Stanley, R.S., 146 Stauffer, M.R., 75,92, 94,101,143,437,443 Steiger, R.H., 271, 277, 298 Stephens, W.E., 378 Stephenson, N.C.N., 305,357 Stern, C.R., 364, 381 Stern R.J., 90 Stern, R.J., 298,441, 443 Stern, TW., 471 Stevens, B.P.J., 147,313,315, 318,357
526 Stevens, R.D., 47 Stevens, R.K., 55,94 Stevenson, D.J., 166,175, 418 Stewart, A.D., 131,147 Stewart, A.J., 309, 313, 318, 323, 325, 329, 356358 Stewart, J.H., 120, 133, 134,147, 271,298 Stoeser, D.B., 65-67, 69, 70,93, 94,441, 506 Stolz, A.J., 365, 366,381 Stone, J., 293 Storey, M., 161, 170,177,499,507 Stosch, H.G., 366, 368,381 Stowe, C., 91,141 Stowell, H.H., 295 Strachan, R.A., 441 Streckeisen, A,, 216, 260 Strobell, J.D., Jr., 471 Strong, D.F., 128, 130,147 Strong, P.G., 147 Stroud, W.J., 114,147, 357 Stuckless, J.S., 294 Studley, S.A., 412 Stukas, V., 130,148 Stuwe, K., 306, 307, 325, 328, 331-334, 340,357 Subbarayudu, G.V., 277,298 Subramanian, A.P., 218, 260 Suddaby, P., 146 Sudo, A., 172,177 Sumner, W., 472 Sun, S.S., 8, 9, 45,53,152,155,156,166,178,348 Sutcliffe, R., 138 Sutcliffe, R.H., 116, 119, 139, 143, 148 Sutter, J.F., 474 Sutton, J., 136, 263,298,440 Svetov, A.P., 106, 108,140, 148 Swan, M.M., 271,277,298 Swanepoel, D.J., 213 Swapp, S.M., 302,304,317,333,357 Swift, P.N., 452, 478 Syme, E.C., 29,47, 53 Tait, R.E., 343, 357 Takahashi, Y., 356 Tan, F.C., 397, 416 Tankard, A.J., 97, 100, 101, 111, 112, 125, 126, 148 Tanner, J.G., 218, 260 Tarney, J., 3, 7 , 29, 45, 45, 52, 53, 84, 94, 97, 100, 109,148,151-161,164,172,174-179,441,442 Tatsurni, Y., 172,177 Tatsumoto, M., 297, 503, 506 Taylor, EC., 189, 212 Taylor, H.P., 246, 260, 279-281, 298
References index Taylor, P.N., 31, 50, 152, 176, 313, 351, 353, 432, 435,442,505 Taylor, S.R., 7 , 51-53, 176, 365-371, 377, 379381,501,503, 506 Taylor, S.W, 147 Teixeira, W., 146 Tewksbury, B.J., 464,478 "hacker, M.S., 461, 462,478 Thakur, V.C., 444 Thirlwall, M.E, 18,174 177,441,442 Thom, A., 47,89,293,503 Thomann, W.E, 464,476, 478,506 Thomas, A., 347, 427,445 Thomas, J., 354 Thomas, J.J., 263, 265, 269, 271, 282, 298 Thomas, R.J., 268, 296 Thomas, W.A., 128,148 Thomas, W.M., 272, 283-285, 292, 298, 299, 359, 478, 480 Thompson, A.B., 334, 340, 344, 349, 353, 357, 469,478 Thompson, A.G., 469,473, 474, 478 Thompson, R.N., 9,10,53,177, 204, 205,212 Thomson, J.W., 52 Thorensen, K., 368,378 Thornber, C.R., 137 Thorne, A.M., 417 Thost, D.E., 306,350, 357 Thy, P., 254, 260 Tilley, B.J., 142 Tilley, C.E., 155,179 Tindle, A.G., 52, 212, 297 Tingey, R.J., 356 Tobi, A.C., 309,357 Todd, S.G., 211 Todt, W., 352,444,506 Toft, P.B., 381 Topley, C.G., 441 Tornroos, R., 211,213 Torske, ' I 107,108,137 , Touret, J., 301, 331, 335, 336, 338, 357 Tourpin, S., 486, 507 Towe, K.M., 406,417 Tredoux, M., 201, 210 Trehu, A., 137 Treloar, P.J., 302, 305, 307, 309, 312, 316, 317, 317,325,328,344,357 Trendall, A.F., 383-385,398, 402,412, 415-417 Treuil, M., 54 Trevena, A.S., 456, 478 Trinquard, R., 50 Trudel, P., 49, 51 Trushkov, Yu.N., 122,148
References index Turcotte, D.L., 97,148 Turner, J.S., 170, I75 lbrner, R.J.W., 470, 479 Tveten, E., 351 Tweto, O., 459, 460, 479 Twist, D., 101,148, 199, 202, 204, 206, 211, 213 vier, N., 474 Tynvhitt, D.S., 209 Ueng, W.C., 106,148, 433, 434, 445 Ueng, W.L., 434,443 Upadhyay, H., 353 Upton, B.G.J., 364,378, 381 Vaasjoki, M., 49, 429, 445 Vachette, M., 36,53 Vail, J.R., 436, 445 Valbracht, P.J., 108, 109, 148 Valley, J.W., 156, 161, 162, 174, 219, 259, 260, 301, 305, 313, 329, 333, 336, 339, 340, 348, 353, 354, 357 358,380 Van Bever Donker, J.M., 325,352 Van Biljon, W.J., 182, 213 Van Breemen, O., 363,364, 373, 381, 427,445 Van Calsteren, P.W., 176, 365, 381, 504 Vance, K.R., 33,47 Vance, R.K., 453,454,473, 479 Van der Molen, I., 234, 260 Van der Neut, M., 210 Van der Pluijm, B.A., 422, 445 Van Kranendonk, M.J., 320,358 Van Schrnus, WR., 51, 53, 75, 95, 116, 148, 263, 265, 292, 295, 298, 422, 440, 444, 445, 452, 467,472, 477,507 Varet, J., 178 Vayrynen, H., 74,95 Veizer, J., 397, 417 Vermaak, C.E, 218,260 Vernon, R.H., 304, 309, 322, 328, 329, 334, 340, 341,345,348, 358 Vidal, P., 440, 497, 503, 507 Vielzeuf, D., 340, 343, 348, 358, 366, 367, 381, 438,445 Visser, D., 305, 313, 358 Vitanage, P.W., 350, 353, 356 Vivallo, W., 18,53, 429, 445 Vocke, R.D., 507 Voitsekhovsky, S.N., 136, 209 Volpe, A.M., 86, 95 von der Borch, C.C., 133, 135,148 Von Brunn, V., 406,416, 417
527 Von Gruenewaldt, G., 3, 97, 148, 154, 156, 159, 176, 181, 182, 194, 196, 197, 202, 203, 205, 206,208,210, 212, 213 Voshage, H., 376,381 Vry, J.K., 339,358 Vuorelainen, Y., 186, 213 Vyhnal, C., 296 Wager, L.R., 181,213 Wager, R.L., 243,261 Walcott, R.I., 75, 90 Walde, D.H.G., 409, 417 Walker, C., 168,178 Walker, J.C.G., 384,385,402,406,408,409,417, 418 Walker, N., 449, 459,479 Walker, N.W., 459,472, 479 Wall, V.J., 304, 309, 316, 329, 348, 355 Wallace, P., 349 Walraven, E, 197, 213 Walsh, J.N., 295 Walsh, K.L., 211 Walsh, M.M., 412 Walsh, N.J., 52, 145 Walter, M., 411, 418 Walter, M.R., 402, 406, 412, 418 Wandless, G.A., 380 Wang, G-R., 440, 446 Wanless, R.K., 140 Ward, D.B., 468,474 Ward, P., 428, 445 Wardle, R.J., 103, 104, 148 Warnke, D.A., 449,479 Warren, A., 52,145 Warren, R.G., 304, 309, 311, 318, 323-326, 328, 341,358,364,381 Wasenius, P., 353 Wass, S.Y., 366, 382 Wasserburg, G.J., 271, 275, 277, 294, 298, 458, 479, 488,505 Waters, D.J., 305, 309, 313, 315, 317, 323-329, 331,339-341,343,345,358, 359 Waters, F.G., 154,156, 161, 168,169,178 Watson, E.B., 285, 286, 298 Watson, J.V., 136 Watson, S., 169,178 Watters, B.R., 29, 53, 75, 95 Watterson, J., 319,350, 440 Weaver, B.L., 44, 45,53, 100, 109,148, 151-153, 155-159,161, 165,166, 172,178, 179, 503 Weaver, S.D., 177 Webb, J.S., 380 Weber, W., 49
References index
528 Wegmann, C.E., 72, 95 Weiblen, P.W., 119, 143, 148, 193, 194, 211, 213, 220,259,416 Weidner, J.R., 253, 261 Weis, D., 220, 245, 252,256, 261 Weiss, N., 7 , 51 Weixing, H., 48 Welin, E., 17, 54, 430, 445 Welke, H.J., 52 Wells, J.D., 141, 459, 479 Wells, P.R.A., 345, 359 Wendlandt, E., 372,374,381, 382 Wenner, D.B., 279,298 Wernicke, B.P., 356,444 Wessels, R.L., 475 Westra, L., 19, 48, 304, 313, 315, 322, 325, 327329,334,336,338,345,356,429,440 Wetherill, G.W, 292,479 Whalen, J.B., 252,261, 273, 274, 298 Whales, C.L., 339, 341, 343,359 Wheeler, E.P., 221, 223, 251,261 White, A.J.R., 288, 293 White, C.A., 452,479 White, D.L., 271, 277, 299 White, R.S., 114,139, 142,148,153, 167, 169,178 White, W.M., 494,507 Whitney, P., 296 Whitney, P.R., 250,261, 328,354 Wiebe, R.A., 4, 97,148, 215-217, 221, 223-225, 227, 229-231, 233-238, 241, 244-248, 250254,255, 256, 260, 261,299 Wiener, R.W., 316, 318,359 Wilband, J.T, 141, 148,445 Wild, T, 224, 246,261 Wilde, S.A., 307,359,504 Wiles, J.W., 183, 213 Wilford, J.R., 143 Wilhelm, H.J., 213 Wilkinson, A.F., 353 Williams, H., 127, 130,147, 149 Williams, I.S., 88, 298, 353, 363, 365, 372, 380, 417 Williams, M.L., 295, 350, 463, 471, 474, 475, 477, 479 Willis, I.L., 147 Willis, I S . , 357 Wilmart, E., 252, 261 Wilson, A.H., 100,146, 149,182-184,212, 213 Wilson, E.D., 456, 479 Wilson, I.H., 11, 41, 43-45,54, 116,149 Wilson, J.T, 55, 56, 95 Wilson, M.R., 18,54,273,299,428,429,445,497, 507
Wilton, D.H., 147 Winchester, J.A., 9, 50, 54, 131,142, 149 Windley, B.E, 5, 7, 48, 53, 56, 57, 59, 86, 95,176, 218-220, 256, 261, 264, 267, 292, 299, 352, 354, 419, 420, 422, 423, 427, 430-433, 438, 440-442, 444-446 Windrim, D.P., 382, 426, 446 Wirth, K., 159,178 Woakes, M., 136 Wobus, R.A., 271,284,299 Wodicka, N., 349 Wold, R.J., 116,149 Wones, D.R., 267,292, 296 Wood, B.J., 304, 325,359 Wood, D.A., 8, 54, 98, 129, 149, 152, 156, 178, 179 Wooden, J., 449,454,479 Wooden, J.L., 47, 245, 246, 255, 265, 267, 267, 272, 292, 295, 299, 449, 450, 467, 471, 472, 479, 497,502 Woodhead, J.A., 298 Woodward, L.A., 455,467,472, 479 Wormald, R.J., 293 Worst, B.G., 182, 183, 213 Wortman, G.L., 451, 462, 468, 470,479 Wright, L.A., 449,480 Wrucke, C.T., 473 Wu, 'l-W., 279,281, 282,299 Wyborn, L.A.I., 7, 40, 41, 45, 48, 54, 113, 114, 116,139, 149,175,426, 427,441, 446 Wyllie, P.J., 288, 289, 299 Wyman, W.E, 455,480
Xiao, X-C., 444 Xu, R.H., 504 Yamazaki, M., 356 Yanai, K., 352 Yaiiez, P., 492, 493, 508 Yelovskikh, VV, 148 Yeo, G.M., 401, 411, 418 Yoder, H.S., 155,179 Yoshikura, S., 356 Young, D.N., 343,359 Young, E.D., 283, 298, 299, 305, 325, 359, 450, 478, 480 Young, G.C., 467,480 Young, G.M., 120, 122, 132-134, 149, 401, 401, 418 Zajac, I S . , 398, 414 Zartman, R.E., 120,144, 149,298,376,382,474
References index Zeitler, P.K., 353 Zen, A-En, 283, 299 Zhang, C., 440, 446 Zhao, 2-H., 446 Zhou, ZY., 440 Zietz, I., 116,142, 292, 298, 445, 472
529 Zietz, J., 193, 211 Zindler, A., 152, 164, 165,176 Zonenshain, L., 122,149 Zonou, S., 36,54 Zwaan, B., 132,142 Zwanzig, H.V., 47
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53 1
SUBJECT INDEX
Acatlan terrane, 492 Accretion, 33 Adirondack Mountains, 309, 316, 318, 320, 323, 327, 329, 333,335, 339 Akaitcho Group, 30 Alder terrane, 33, 455, 457 Group, 456 Alice Downs ultrabasics, 42 Ameralik dike suite, 45 Anatectic melt, 335, 340-343 Andean arc, 20 Andes, 30, 38 Andesite, 45, 454 Phanerozoic, 9 Proterozoic, 14, 34 Animikie Group, 105, 106 Animikie rift, 104-106 Anorthosites, 4, 215, 253 Archean, 218 geochemistry, 239 mineralogy of, 235 modes, 228 origin, 253 oxygen isotopes, 246 REE, 239,241 xenoliths, 242, 244 Aravalli Supergroup, 38 Arc accretion, 31, 40 continental margin, 20 high-K calcalkaline, 20 intraoceanic, 22 mature oceanic, 20 Norbotten Porphyry, 18,19 oceanic, 453, 455 Archean, 7 anorthosites, 218 cratons, 45 oceanic crust, 7 Proterozoic boundary, 7 Argon cooling ages, 469 Argylla Formation, 43, 44 Arjeplog porphyries, 18, 20 Arunta Block, 42
Arunta Complex, 309, 313, 315, 316, 323, 325, 326,328,329, 333,340,341 Ash Creek terrane, 31, 453, 457, 466, 468 Aulacogen, 122 Australian basins, 113-116 Australo-American trough, 132-135 Back-arc basin, 461, 463, 466 Bakkilvarri Formation, 13, 20 Baltic Shield, 10-21 Hallandian orogeny, 11 Kviby Group, 20 Lower Holmvatn, 20 Svecofennian orogeny, 11 Svecononvegian-Grenvillian orogeny, 11 Baltic Shield, 107 Banded iron formations (BIF), 5 (see also Ironformations) Basalt, 33 Albin, 13 calcalkaline, 10 continental flood, 10 Phanerozoic, 9 Proterozoic, 14, 34 spinifex, 27 Basin back-arc, 20, 40, 456, 460, 469 Bothnian, 17 Chaibasa, 38 continental-margin, 33, 39, 40 foreland, 458 Glengarry, 42 intra-arc, 33 Lau, 24 marginal, 35 transpressional, 464 Basins, South Africa, 100, 101, 111-113 Beauparlant Sub-group, 27 Belt-Purcell Basin, 120-122 Bergslagen field, 17, 19, 20 Bimodal volcanic suite, 452, 455, 457 Biscay Formation, 42 Bjerkreim-Sokndal lopolith, 230 Boninite, 18, 44
532 Bothnian Basin, 17 Broken Hill block, 43 Bushveld Complex, 3, 194-206 Canadian Shield, 110, 111 Cape Smith Belt, 102, 103 Carbonic metamorphism, 335 Carrizo Mountains terrane, 458, 459 Chaibasa Basin, 38 Cheyenne Belt, 452 Cheyenne orogeny, 33 Cheyenne shear zone, 452, 464 Chukotat Group, 26-28 Churchill-Superior boundary zone, 27 Circum-Superior Belt, 24, 25, 28, 29, 101-104 Cochetopa-Salida terrane, 32 Cochetopa terrane, 461 Collisions, arc-continent, 466, 467 Continental CNSt, 7 flood basalts, 10 growth rates, 5 rifts, 3 Continental-margin arc, 33 Continental-margin back-arc basin, 32 Continental rifts, 362 Cook Gap Schist, 42 Cooling rates granulites, 333-335 Cordilleran terranes, 469 Craton, 469 Cratonic sedimentation, 463 Cratonization, 469 Cree Lake zone, 29 Crustal extension, 345 formation age, 447, 470 genesis rate, 500, 501 province, 447 reworking, 318-320, 361 xenoliths, 4, 465 Cumulate, 235, 244, 249 Cumulus layering, 231 Dahomeyan-Pharusian Belt, 126,127 Dalma Belt, 38 Dalma Epidiorites, 38 Damaran Belt, 497 Deformational history, 466 Dehydration melting, 340-343 Depleted mantle, 454, 492, 494 Detrital zircons, 469 Ding Dong Downs volcanics, 42
Subject index Diorite, 216, 223, 246-248 Dos Cabezas-Pinal terrane, 33, 457, 468 Dubois terrane, 31, 33, 452 Dullstroom Formation, 198 Duluth Complex, 192-194 Dykes, 151-174 chronology of, 154 composition of, 154-164 form of, 152-154 Late Proterozic, 162-164 Nd isotopes in, 161, 164, 165 norite, 159 origin, 169-173 oxygen isotopes in, 161,162 petrology of, 154-156 Sr isotopes in, 161, 164, 165 trace elements in, 156-164 Dyke swarms, 44,151-174 Early Proterozoic island arc sequences, 45 Early Svecofennian volcanics, 20 Eau Pleine suture, 24 Egersund Massif, 217 Egersund-Ogna intrusion, 229 Elliot Lake Group, 22 epsilon-Nd values interpretation, 492 Proterozoic rocks, 492 Eskimo Formation, 26, 28 Etendeka Suite, 10 Etendeka volcanics, 22 Ethiopian rift, 16 Felsic volcanics, 490 Ferropicrites, 16 Fe-Ti oxide deposits, 252 Fiery Creek volcanics, 44 Flaherty Formation, 26 Flin Flon Greenstone Belt, 28, 29 Flood basalts, 161, 172, 173 Fluid buffering, 339 Fluid inclusions, 338 Fox River Belt, 24 Fox River Group, 28 Fox River sill, 188 Franklin Mountain assemblage, 464 Furua Complex, 313, 319,323,326, 327 Gabbro, 17 Garell Park pluton, 461 Gawler Craton, 42 Gawler Range volcanics, 43
Subject index Geochemistry anorthosites, 239 Geothermobarometry, 320,324 Gerowie tuff, 42 Glacial deposits and iron-formations, 401 Glengany sub-basin, 42 Granites, 4, 223, 251, 252, 460 anorogenic, 4, 252 -anorthosite association, 1 Granitoids, 490 Granulite-facies xenoliths, 361 Granulites, 4, 302, 375, 376, 460 Antarctica, 319 carbon isotopes, 338 deformation of, 315 Early Proterozoic, 312 geothermobarometry, 324 high-pressure, 325-327 histories, 320 horizontal structures in, 316, 317 Lapland, 317, 323 Late Proterozoic, 313 metamorphic features, 304-308 Mid-Proterozoic, 313 mineral assemblages, 324 occurrence, 302-313 oxygen isotopes, 338 P-T conditions of, 320 role of fluids, 335-340 sapphi-'qe, 326 structural styles of, 313 tectonic setting, 344 textures, 324 thrusting of, 317 Gravity anomalies, 218 Great Bear magmatic zone, 30 Great Dyke, 182-184 Green Mountain terrane, 31, 452 Greenstone Belt Belomoride, 15 Bergslagen, 20 Cheyenne, 452 Damaran, 497 Early Proterozoic, 10, 29 Flin Flon, 28, 29 Fox River, 24 geochemistry of, 37 Karasjok, 20 Kautokeino, 20 Kiruna, 20 Lynn Lake, 29 Mazaruni, 35 Tampere Schist, 19, 20
533 Vittangi, 20 Greenstones, 2, 452 Grenville Front Tectonic zone, 302 Grenville Province, 219, 302, 303, 309, 310, 313, 317, 318, 320, 323, 325-327, 331, 334, 346, 458 Guyana Shield, 34, 35 Haimarka Formation, 35 Halls Creek Inlier, 42 Hamersley Group, western Australia, 384, 398 Haparanda Series, 18,20 Harp Lake Complex, 230 Harly Anorthosite Complex, 42 Harts Range meta-igneous complex, 42 Heat production, 500 Hepburn intrusive suite, 30 High-Al orthopyroxene megacrysts, 230,237 Highland Complex, 333 Highland Series, 131, 310, 323, 325 Hottah magmatic arc, 30 Hualapai terrane, 32, 449, 454 Huronian Supergr-uup, 22 Iapetus rift, 127-132 Intracontinental rifting, 45 Irindina supracrustal assemblages, 42 Iron-formations, 383-418 abundance of, 384 Archean, 384 bulk chemistry of, 393 carbon-Al203 relation in, 405 carbon-sulfur relations in, 405 carbonate isotopic composition in, 405 cesium anomalies in, 394 deep marine origin of, 407 definition of, 383 distribution in time of, 384 europium anomalies in, 394 formation in density-stratified ocean system, 406 glacial aspects of, 401, 411 hydrothermal input into, 396 metamorphism of, 385 origin of, 401, 402 oxidation state of, 393, 395 paleoenvironmental interpretation of, 403,409 rare earth element (REE) chemistry of, 394396 sedimentology of, 397, 398 South African, 398-400 stratigraphy of, 397, 398 trace element chemistry of, 395, 396 transgressive sequence in, 407,408
534 Iron, hydrothermal input of, 395, 396 Island arc, 38 Isobaric cooling, 301, 328 Isothermal decomposition, 301, 327 Isotopic ages, 220 of anorthosites, 220 Issineru Formation, 35 Ivanpah orogeny, 33,450 Jatulian meta-arenite, 1 7 Jatulian volcanic suites, 20 Jimberlana intrusion, 182-184 Jorma Ophiolite, 16, 20 Jorn Group, 18 Jotunites, 221 Juvenile mantle, 490, 495 Kaapvaal Craton, South Africa, 400 Kaapvaal volcanic sequences, 45 Kalahari Copper Belt, 123, 124 Kalevian suites, 16 Karasjok Greenstone Belt, 20 Karelian Province, 13 Karoo Province, 20 Karoo volcanics, 45 Kautokeino Greenstone Belt, 20 Kemio-Orijarvi-Lohja volcanic belt, 19, 20 Kerala Khondalite, 309, 311, 313, 323, 338 Ketilidian orogen, 22 Ketilidian Province, 495 Ketilidian terrane, 31 Keweenawan Rift, 116-120,192-194 Kibaran Belt, 124 Kiglapait intrusion, 189, 190, 223, 244 Kiiminki Belt, 16 Kirnberlite, 363 Kiruna Greenstone Belt, 18-20 Kissynew Domain, 29 Komatiite, 14, 16, 18, 27, 28, 32, 35, 38, 45, 46 La Ronge Belt, 29 Labrador trough, 103, 104,384,387 Lake Superior region, 384, 386 Lamarche Sub-group, 26, 27 Lamprophyre, 363 Lapland Granulite Belt, 15 Lapland Granulite Complex, 302 Late Proterozoic dykes, 162-164 Late Proterozoic rifts, 125-135 Lau Basin, 24 Layered igneous complex, 3 Layered intrusions, 181-208, 230 Australian, 187, 188
Subject index Fennoscandian, 184-187 origin of, 206-208 parental magmas, 182-194 Lebombo Suite, 10 Lebowa Granite Suite, 202, 203 Leichhardt Metamorphics, 43,44 Leptite, 17, 19 Lesotho Suite, 10 Leucogabbro, 216 Leuconorite, 216 Leucotroctolite, 216 Jiwisian Complex, 21, 22 Lithologic association, 1 Lithosphere, 344 Lithospheric thinning, 345 Llano uplift, 459 Loch Maree amphibolites, 22 Loch Maree Group, 21 Luostarin Series, 15 Lutzow-Holmvatn Complex, 329, 331 Lynn Lake Greenstone Belt, 29 Mafic dykes (see dykes) Mafic dyke swarm, 3 Magna Lynn Metabasalt, 43 Mantle evolution, 164-169 plumes, 38, 46, 169-172 primitive, 167-169 Manzano assemblage, 463, 464 Marshfield terrane, 24 Mazaruni Greenstone Belt, 35 Mazatzal Group, 456 Mazatzal orogeny, 33,468 Mazatzal Province, 31,33, 455, 465, 470 Metadiabase, 33 Metagabbro, 33 Metamorphic minerals in iron-formations, 392 Metamorphism, 8 Metasomatism, 8 Michikamau intrusion, 230 Microcontinent, 22 Middle Proterozoic rifts, 116-124 Migmatites, 340-343 Mojave Province, 33,448, 466, 468 Mojave-Sonora megashear, 464,465 Molson Dykes, 28 Moore Gulch fault, 465 MORB-normalized element plot, 8 Mount Bonnie Formation, 42 Mount Isa Inlier, 43, 44 Muskox intrusion, 190-192 Mylonite, 315, 452
Subject index Nagssugtoquidian mobile belt, 21, 22 Nain Anorthosite Complex, 221-224, 230 Namaqualand Metamorphic Complex, 36 Namaqua Province, 37,131, 309,323-328,340 Narracoota volcanics, 42 Naru Plateau, 36 Neodymium isotopes, 244,485 global coverage, 481 interpretation, 490 mixing model, 493 model ages, 485, 488 Niagara suture zone, 22 Nikel Series, 15 Norbotten Porphyry arc, 18-20 Nordkalott Volcanic Province, 15 Norite dykes, 159 North Atlantic Craton, 22, 31 North Svecofennian subprovince, 17,19 Oaxaca terrane, 492 Oceanic arc, 470 Oceanic plateau, 499, 501 Oklahoma Province, 33 Oonagalabi gneiss complex, 42, 43 Ophiolite, 2, 3, 24, 455, 456 Ophiolite complex, 45 Orange River Group, 36, 38 Orogenic belt Cape Smith Belt, 102, 103 Circum-Superior, 101-104 Grenville, 122, 123, 492, 494, 495 Kemio-Orijatvi-Lohja Kiiminki, 16 Labrador trough, 103, 104 Lapland Granulite, 15 Nagssugtoquidian, 21, 22 Nagu-Korpo, 19 Pechenga-Varzuga, 15 Skellefte, 20 Svecokarelian, 106-109 Thompson Nickel, 24 Wopmay, 109,110 Orogenic collapse, 345 Outokumpu Allochthon, 16, 20 Overlap assemblages, 32,447,459-461 Oxygen isotopes, 246
P-T, 321,326,328 conditions, 321 data, 322 estimate, 321 histories, 321, 328 P-T paths, 321,322,327-331
535 clockwise, 329-331 counter clockwise, 329-331, 344 decompressional, 326, 328 retrograde, 331-333 Panton sills, 42 Partial melting, 301, 344-343 Payson Ophiolite, 33 Pb isotopic composition, 244 Pechenga-Varzuga Belt, 15 Pecos terrane, 31,33,455,461,466 Pembine-Wausau terrane, 24 Penokean orogen, 22 Petrogenetic grid, 320 Phanerozoic ophiolites, 33 sediments, 482, 483 Pilgujatvi Suite, 16 Pine Creek Inlier, 42 Piumhi Massif, 35 Plutons, 223 diapiric, 225, 228 layered, 225 massive, 225 Port Manvers Run intrusion, 230 Povungnituk Group, 26 Primitive mantle, 167-169 Project 217 (IGCP), 2 Proterozoic andesites, 14, 34 basalts, 14, 34 crust, 481,496 crustal growth, 497 greenstone belts, 10 juvenile crust, 498 mantle plumes, 499, 501 of the Canadian Cordillera, 499 orogens, 5 recycled crust, 496 rifts, 97-135 subduction, 20 terranes, 496, 498 volcanic suites, 8, 10 xenoliths, 361 Proterozoic greenstones Africa, 36-38 Australia, 40-44 China, 34-40 India, 38, 39 South America, 34, 35 SW North America, 31-33 Proterozoic orogenesis, 45 Proterozoic rifts, 97-135 Proterozoic subduction, 36
536 Province boundaries, 464, 465 Purtuniq Ophiolite, 26, 27, 77-83 Quartzwacke, 457,458 Quinnesec Formation, 24 Rapitan Group, iron-formations in, 384, 401 Rare earth elements (REE), 156-158,486 Rauer Group, 339,343 Rayner Complex, 131, 136, 319 Red Rock Group, 456 Reindeer Lake zone, 29 Richtersveld Sub-province, 37 Rifts and plate tectonics, 135, 136 Animikie, 104-106 Australian, 113-116 continental, 97-135 Early Proterozoic, 99-116 geochemistry of, 108, 118, 119, 129, 130, 133, 134 Proterozoic, 97-135 volcanism in, 117-119 Rogaland Massif, 220 Rooiberg Group, 198, 199 Ruin granite, 457 Rustenburg Suite, 201, 202 Sandstone provenance, 458,459 Scourie dykes, 21,156, 157 Serpentinites, 17 Shear zones, 450,452 Singhbhum Craton, 38 Sinclair Formation, 36 Skellefte Belt, 17, 20 Skellefte Group, 18 Slave Craton, 30 Sm-Nd isotopic system, 481, 486 alteration, 487 effect of hydrothermal alteration, 486 “Snowball Earth”, 411 South African basins, 100,101, 111-113 Southern Cape conductive belt, 36 South Svecofennian Sub-province, 17 Soutpansberg Group, 111, 112 Sr isotopic composition, 244 Stay Creek volcanics, 42 Strangways Metamorphic Complex, 42 Subcontinental lithosphere, 45 Subduction zones, 501 Superior Craton, 22, 24, 26, 29 Supertermne, 447,453 Supracrustal rocks, 449, 454
Subject index Svecofennian Province, 16 Svecokarelian Belt, 106-109 Tampere Schist Belt, 19, 20 Tectonic models collisional, 344 extensional, 344 Tectonic setting, 1 Temporal distribution of iron-formations, 384 Terranes Acatlan, 492 Alder, 33, 455, 457 Ash Creek, 31, 453, 457, 466,468 Marshfield, 24 Oaxaca, 492 Pembine-Wausau, 24 Wrangellia, 499 Terranes, 319, 447,461 Thermobarometric studies, 219 Thompson Nickel Belt, 24 Tonto Basin Supergroup, 456 Trans-Hudson orogen, 22,24, 29 Transcandinavian granites, 21 Transvaal Supergroup, 384 Trondhjemite, 17 Trondhjemitic dykes, 33 Troodos ophiolite, 24 Underplating, 45 Urucum region, Brazil, 384 Vioolsdrif Batholith, 37 Vishnu Complex, 454 Vittangi Greenstone Belt, 20 Volcanic rocks Cas’kejas Group, 13 geochemistry of, 10 Kalevian, 20 Kiruna, 13 Kviby Group, 13 Lower Holmvatn Group, 13 Nussir Group, 13 Petsamo Group, 13 Upper Holmvatn Group, 13 Upper Lapponian, 13 Vittangi Greenstone Group, 13 Wasekwan Group, 29 Waterberg Group, 112, 113 Wathaman Batholith, 29 Watts Group, 24, 26 Wet Mountains terrane, 32
Subject index Willyama Complex, 43, 313, 315, 317, 340, 341 Wilson Cycle, 30 Widgiemooltha Dyke Suite, 41 Woodward Dolerite, 42 Wopmay orogen, 22, 30, 109, 110 Wrangellia terrane, 499 Wyoming Craton, 33, 452, 466 Xenoliths ages, 363 alteration, 367 anorthositic, 242, 244 crustal, 4 geochemistry, 367-371 geochronology of, 371-375
537 high-pressure, 242, 244 lead isotopes, 374,375 lower crustal, 361, 362 mineralogy, 363-366 neodymium isotopes, 372 Proterozoic, 361 thorium in, 369 transport, 366, 367 uranium in, 369 Yavapai Province, 31, 33,451, 465, 468-470 Yavapai Supergroup, 31, 453, 454 Zamu Dolerite, 42 Zoned iridescent plagioclase, 233, 235
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